Ph.D. Thesis - IS MUNI

MASARYK UNIVERSITY

FACULTY OF SCIENCE

Department of Geological Sciences

Ph.D. Thesis

Sudipta Tapan Sinha Brno, 2014

MASARYK UNIVERSITY

FACULTY OF SCIENCE

Department of Geological Sciences

FACTORS CONTROLLING A MICROCONTINENT

DEVELOPMENT DURING CONTINENTAL BREAKUP-

THE ELAN BANK CASE STUDY

Ph.D. Thesis

SUDIPTA TAPAN SINHA

Supervisors

Doc. RNDr. Rostislav Melichar, Dr.

Prof. Michal Nemčok, DrSc. (Consultant)

Brno, 2014

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Bibliografický záznam

Autor: Sudipta Tapan Sinha Přírodovědecká fakulta, Masarykova univerzita Ústav geologických věd

Název práce: Faktory řídící vývoj mikrokontinentů při kontinentálním rozpadu na příkladu Elan Bank

Studijní program: Geologie, doktorský

Studijní obor: Geologické vědy

Školitel: Doc. RNDr. Rostislav Melichar, Dr.

Školitel-konzultant: Prof. Michal Nemčok, DrSc.

Academic Year: 2013-2014

Pages: 305

Keywords: kontinentální rozpad, mikrokontinent, hyper-extendovaný okraj, plášťový chochol, riftové zóny, horizontální posuny, Východní Indie, Elan Bank

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Bibliographic record Name of author: Sudipta Tapan Sinha Faculty of Science, Masaryk University Department of Geological Sciences

Title of dissertation: Factors controlling a microcontinent development during continental breakup- The Elan Bank case study

Degree programme: Geology, Doctorate

Field of study: Geological Sciences

Supervisor: Doc. RNDr. Rostislav Melichar, Dr.

Supervisor Consultant: Prof. Michal Nemčok, DrSc.

Academic Year: 2013-2014

Pages: 305

Keywords: Continental breakup, microcontinent, hyper-extended margin, mantle plume, competing rift zones, ridge-jump, strike-slip margin, East India and Elan Bank,

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Candidate’s Declaration

I hereby declare that the work, which is being presented in the thesis entitled “Factors controlling a

microcontinent development during continental breakup-The Elan Bank case study” in partial fulfilment

of requirements for the award of the degree of Doctor of Philosophy, submitted in the Department of

Geological Sciences, Masaryk University, Brno, Czech Republic is an authentic record of my own work

carried out for the period of 4 years from 2010 to 2014 under the supervision of Prof. RNDr. Rostislav

Melichar and Prof. Michal Nemčok.

The matter embodied in this thesis has not been submitted by me for the award of any other degree. To

the best of my knowledge and belief, it contains no material previously published or written by another

person nor material which to a substantial extent has been accepted for the award of any other degree

or diploma of the university or other institute of higher learning, except where due acknowledgement

has been made in the text.

Date:

Place: (SUDIPTA TAPAN SINHA)

This is to certify that the above statement made by the candidate is correct to the best of our knowledge. Doc. RNDr. Rostislav Melichar, Dr. Department of Geological Sciences, Masaryk University Brno, Czech Republic

Prof. Michal Nemčok, DrSc. The Energy & Geoscience Institute University of Utah Bratislava, Slovakia

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Abstrakt

Mikrokontinent Elan Bank, který byl spojen s východní Indií až do spodní křídy, v současné době leží v

jižní části plateau Kerguelen v jižní části Indického oceánu nedaleko Antarktidy. Je to východozápadně

směrovaný výběžek, který na mapě vytváří pozitivní batymetrickou a gravitační anomálii.

Tato práce se zabývá dvěma hlavními hypotézami vysvětlujícími vytvoření mikrokontinentu Elan Bank.

První je hypotéza přesunu plášťového chocholu, která vyžaduje plášťový chochol k iniciaci obnoveného

riftingu, čímž se postupně uvolní mikrokontinent. Druhá hypotéza známá jako hypotéza konkurenčního

riftu vysvětluje složitost "soutěží" mezi propagujícími se riftovými zónami v extrémně extendovaném

kontinentálním okraji, takže dojde k štípání litosféry za vzniku mikrokontinentu.

Datové soubory použité v této studii zahrnují reflexní seismiku a data z vrtů z východní Indie a

mikrokontinentu Elan Bank, hodnoty potenciálních polí měřená družicově a pomocí vlečných lodí a

petrologické informace z publikovaných souborů dat. Použité metody jsou založeny na interpretaci

seismických dat a dat potenciálních polí, techniku seismického profilování, určování času tektonických

aktivit, syntetické rekonstrukce oblasti mezi mikrokontinentem Elan Bank a východní Indií a spasovávání

dráhy mikrokontinentu Elan Bank se stopou horké skvrny v Bengálském zálivu.

Stavby extrémně extendované kůry spřažených okrajů východní Indie a mikrokontinentu Elan Bank se

vyznačují různými fázemi deformace, jako je extenze, ztenčování, exhumace a rozšiřování oceánského

dna. Proximálně-distální stavba zahrnující obě spojené a oddělené domény byla interpretována na obou

spřažených okrajích. Oba spřažené okraje jsou segmentovány tak, že obsahují segmenty kolmé, kosé i

horizontálních posunů. Ve východní Indii jsou kolmé okraje buď nevýrazně propojeny prostřednictvím

stavby ramp-flat-ramp nebo výrazně spojené horizontálními posuny. Přítomnost domény exhumace byla

doložena ve východní Indii i na severozápadním okraji mikrokontinentu Elan Bank. Stavba ekvivalentních

okrajů ukazuje, že okraj východní Indie se nachází v nadložní kře vyklenutého hlavního odlepení (horní

deska). Mikrokontinent Elan Bank pak leží v podložní kře stejné poruchy (spodní deska).

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Stratigrafické korelace vrstev v mělkovodních a hlubokovodních částech extenzních riftových zón

ukazují, že při kontinentálním rozpadu nedochází k diskordanci z rozšiřování, ale že konformní povrch

může být korelován přes riftovou zónu. Je navrženo, aby první hlubokomořská sekvence, která obaluje

celou syn-kinematickou sekvenci od proximální k distální části okraje okraji, byla horní hranicí

skutečného rozpadu kontinentu. Interpretace seismického a tektonického časování ukazují následující

historii rozpadu podél východoindického okraje. K prvnímu štěpení Indie a Elan Bank a Antarktidy podél

východní části Cauverské příkopové zóny došlo kolem valanginu. Druhé nastalo mezi mikrokontinentem

Elan Bank a Indií podél riftové zóny Krishna-Godavari v raném aptu. Rozpad je nejmladší podél

Coromandalské zlomové zóny (střední apt), která je spojena s druhým eventem rozpadu v Cauverské

pánvi, protože mikrokontinent Elan Bank byl odsunut od indického okraje podél Coromandalského

systému horizontálních posunů. Cauverská oblast prodělala dva izostatické výzdvihy v důsledku dvou

rozpadových událostí, zatímco v oblasti Krishna-Godavari proběhl pouze jeden izostatický výzdvih.

Ukazuje se, že hypotéza konkurenčního riftu lépe vysvětluje uvolňovací mechanismus mikrokontinentu

Elan Bank. I když raný rozpad mezi Indií a Antarktidou začal ve východní části Cauverské příkopové zóny

a na jižním okraji mikrokontinentu Elan Bank, propagace rozpadu západním směrem pokračovala podél

zóny Krishna-Godavari i po oddělení v překrývající se části Cauverské zóny. Obě kolmé extenze řídily

rozpad dokud nakonec byly zcela propojeny prostřednictvím Coromandalského horizontálního posunu.

Efektivní kinematické propojení ustavené mezi odumřelou Cauverským riftem a částí riftu Krishna-

Godavari tak podlehlo kontinentálnímu rozpadu. Následkem toho oddělování Indie a Antarktidy nakonec

přeskočilo z jeho počáteční do konečné polohy podél severního okraje mikrokontinentu Elan Bank a na

rift Krishna-Godavari. To znamená, že konkurence mezi dvěma rifty v rámci kontinentálního rozpadu za

asymetrické propagace hřbetu vede k přeskočení hřbetu do nové pozice. Coromandalský transformní

pohyb se nezastavil, dokud mikrokontinent Elan Bank a centrum rozpínání oceánského dna na sever od

něj nebyly odsunuty od Východoindického pobřeží.

Alternativní model přesunu plášťového chocholu byl odmítnut, protože tento model nevysvětluje ani

kinematiku riftingu za vzniku extrémně extendovaného pasivního okraje, ani časové omezení pro různé

tektonické události. Navíc jediným kandidátem pro vulkanismus vázaný na plášťový chochol jsou útvary

vytvořené zcela zřetelně až po rozpadu kontinentu.

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Abstract

The Elan Bank microcontinent, which was attached to East India until early Cretaceous, is currently

located inside the southern Kerguelen Plateau in southern Indian Ocean close to Antarctica. It is

characterized by an east-west trending promontory and a positive feature in the bathymetry and gravity

anomaly maps.

The thesis discusses two major candidate hypotheses to explain the Elan Bank microcontinent release

mechanism. The first one is the plume refocusing hypothesis, which requires a mantle plume to initiate

renewed rifting to successively release the microcontinent. The other one, known as competing rift zone

hypothesis, suggests that the complexity in the competition between the propagating rift zones in a

hyper-extended continental margin, to host the lithospheric breakup, contributes to the release of the

microcontinent.

The datasets used in this study include reflection seismic and borehole data from East India and Elan

Bank, satellite and ship-track potential field data and petrological information from published datasets.

The applied methods include seismic and potential field data interpretation, seismic profile marriage

technique, tectonic timing determination, synthetic reconstruction between the Elan Bank and the East

India and matching of Elan Bank displacement track with hot-spot track in Bay of Bengal.

The hyper-extended crustal architecture of the East India and the Elan Bank conjugate margins are

characterized by different phases of deformation, like stretching, thinning, exhumation and sea-floor

spreading. The proximal-distal margin architecture including both the coupled and decoupled domains

has been interpreted at both conjugate margins. Both conjugate margins are segmented

containingorthogonal, oblique and strike-slip segments. In East India, the orthogonal margins are either

soft-linked through ramp-flat-ramp architecture or hard-linked though strike-slip faulting. The presence

of an exhumation domain is interpreted in the East India and north-western margin of the Elan Bank.

The conjugate margin architecture shows that the East India margin is located on the hanging wall side

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of the main convex-up detachment fault, which the upper plate. The Elan Bank is located on the foot-

wall side of the same fault or the lower plate.

The stratigraphic correlation of strata in shallow-water and deep-water parts of the extensional rift

zones indicate that the continental breakup is not represented by a breakup unconformity, but a

conformable surface, which can be correlated across the rift zones. It is proposed that the first deep

marine sequence that envelopes the entire syn-kinematic sequences from the proximal to the distal

margin is the upper limit of the true continental breakup. The seismic and tectonic timing

interpretations suggest the following breakup histories along the East Indian margin. The first breakup

between India-Elan Bank and Antarctica along the eastern portion of the Cauvery rift zone occurred

around Valanginian. The second one took place between the Elan Bank and India along the Krishna-

Godavari rift zone in Early Aptian. The breakup is youngest along the Coromandal fault zone (Mid

Aptian), which is associated with the second breakup event in Cauvery Basin as the Elan Bank was

clearing the Indian margin using Coromandal strike-slip system. As a consequence of two breakup

events, Cauvery region experienced two isostatic uplifts while the Krishna-Godavari region experienced

only one isostatic uplift. It is proposed that the competing rift zone hypothesis better explains the Elan

Bank release mechanism. Although early breakup between India and Antarctica was initiated in the

eastern portion of the Cauvery rift zone and southern margin of Elan Bank, the westward breakup

propagation continued along the Krishna-Godavari rift zone even after the breakup in the overstepping

portion of the Cauvery rift zone. Both orthogonal extension controlled breakups got eventually hard-

linked through the Coromandal strike-slip fault. The effective kinematic linkage got established between

the failed Cauvery rift zone and the portion of the Krishna-Godavari rift zone that underwent a

continental breakup. As a result, the breakup between India and Antarctica eventually jumped from its

initial to its final location along the northern margin of the Elan Bank and the Krishna-Godavari rift zone.

Thus, the competition between the two rift zones to capture continental breakup and asymmetric ridge

propagation resulted in a ridge jump. The Coromandal transform movement did not stop until the Elan

Bank microcontinent and the sea-floor spreading center, to the north of it, laterally cleared the East

Indian margin.

The alternate plume refocusing model was rejected because this model explains neither the kinematics

of rifting and formation of neither the hyper-extended passive margin nor the age constraints for

different tectonic timing events. Additionally, only candidate for the plume related volcanism has been

clearly determined to represent a post-breakup event.

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© Copyright by SUDIPTA TAPAN SINHA, Masaryk University, 2014

All Rights Reserved

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Acknowledgements

At the end of this long voyage, I would like to take this opportunity to express my sincere appreciation

to those who have contributed to this thesis and supported me in one way or the other during this

amazing journey. It would not have been possible to carry out the study without the help and support

of those nice people around me, to only some of whom it is possible to give a particular mention here.

Foremost, I must thank my organization Reliance Industries Ltd and its management for allowing me to

carry out this research, while I am actively working with them. They have been kind enough for

permitting me to use some of their data in this research. I am proud to work in this organization, which

unconditionally extended its support throughout the entire period of the study.

My deep gratitude also goes to Masaryk University, Brno, Czech Republic and its administration.

Without their help and support, it wouldn’t be possible to finish this thesis for me as an external foreign

student.

I would like to express my sincere gratitude to my advisor Professor Michal Nemčok, EGI Laboratory,

Bratislava, who has been not just an advisor but helped in all respect like a friend, philosopher and a

genuine guide. I am deeply solicited by all the useful discussions and brainstorming sessions, especially

during the conceptual development stage. His deep insights and continuous support helped me at

various stages of my research. I thank him for his patience, motivation, enthusiasm, and the vast

knowledge that was kept me alive even in the most difficult times. His valuable advice, edits and critical

comments while writing this thesis aided me a lot during last 4 years. To add more, he has been kind

enough to host me several times during my visits to Bratislava.

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My earnest gratitude is reserved for my advisor, Professor Rostislav Melichar, Masaryk University,

Brno, without whom this study would have been remotely possible. From the inception, he has been a

great support in the University, whenever required. I am indebted to him for his invaluable help and

support, which he had provided throughout the study period. His valuable technical advice and guidance

was a quintessential part of this research. I also thank him to kindly host me during my yearly visits at

Brno.

A very special appreciation is also reserved for my senior colleague Mr Neeraj Sinha, Head Exploration,

Reliance E&P. He has been a great source of support and encouragement. He always stressed on the

fact that this research should have a direct impact not only on our organization but the overall

exploration industry. In a way, he always kept me focused. Apart from his regular administrative support

within the organization, I fondly recollect all the critical technical discussions with him during this

tenure.

I would also like express my honest thanks to my thesis committee chairman, Prof Jiří Kalvoda, Masaryk

University, Brno, and other distinguished thesis committee members for their earnest help to carry out

this research. A very special mention is reserved for Prof Josef Zeman for his timely help as well. My

thanks also go to the all other faculty members in the Department of Geology at the university. I must

thank Mr Petr Bures, the foreign student coordinator at the University, for his administrative support. It

also requires a special mention about his timely communication and apt responses for all official

communications.

I, solemnly, convey my appreciation to the PhD review committee members in Reliance Industries Ltd.,

namely Dr RJ Singh, Mr Palakshi Karjigi and Dr Lalaji Yadav. Their inspiring comments and reviews

always helped me a lot. I would like to also convey my sincere thanks to Dr G. Srikanth, CIO, and

member of E&P technical publication review committee, Reliance Industries for his help in internal

publication review process for articles in peer reviewed journals and international conferences. His

critical reviews have been always helpful.

I express my profound sense of reverence to Prof Gianreto Manatschal, University of Strasbourg,

France. A part of this thesis is very well shaped by his valuable suggestions, deep knowledge and

meaningful comments. I learnt a lot on the technical aspects of hyper-extended margins during many

discussion sessions and personal communications. I would like to thank Dr Dale Bird, Bird Geophysical,

US. Dale has been instrumental to support part of this work by potential field modeling. He also helped

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me to learn some basics of the study of potential field methods. It was always a pleasure to work and

discuss with him about the gravity and magnetic modeling, during which my learning has been improved

significantly. Similarly, I convey my sincere thanks to Mark Longacre, MBL, US, who also actively helped

me to learn few basics on potential field modeling and interpretation of different gravity datasets.

A thesis remains incomplete without the active support of the colleagues and fellows, both present and

the past. I remain sincere to those great people around me in different countries in different continents.

In this short space, only a few names are possible to refer. A special mention goes to Mainak Choudhuri,

my fellow PhD student and team member in Reliance Industries Ltd. The wide verities of discussions

with him have been a great stimulus. Similarly, I thank my team member and colleague Achyuta Ayan

Misra for his constructive opinions and knowledge sharing. A special mention is reserved for my senior

colleague Suraj P Sharma, Reliance as well. I am also grateful to my fellow co-workers in Reliance for

their helps and support in their respective fields of expertise even at odd times. I remain sincere to Dr

Nishikanta Kundu for his comments in subsidence analysis, Ms Vincy Yesudian for helping me learning

biostratigraphic analysis, Ms Deblina Bhawal for her help in seismic mapping, and Dr Pundarika

Dhanukonda, who helped me to prepare maps in ArcGIS. A special mention goes to Sidhartha Sinharay

and Ms Somali Roy, whom I consider few very special people around me.

My special thanks go to all my colleagues in EGI, Bratislava laboratory. I am obliged to Cameron Sheya,

a former colleague at EGI, for helping me to do the official paper works at the beginning of the study. I

am also thankful to Ms Marina Matejova for keeping me updated all the time. My honest appreciation

is reserved for all other lab-mates, who always have been supportive during my stay over there. I am

also thankful to my all colleagues and fellow students in Department of Geology, Masaryk University at

Brno. I am personally obliged to Jan, Adam and others for all the help they provided me in the

University.

I would like to take this opportunity to thank ION Geophysical, US and Geoscience Australia. I am

obliged to these two organizations, which provided me some of the most crucial data for this study.

Here, a special mention goes to the then India-Span Manager with ION, Ms Sujata Venkatraman, who

helped me a lot in this regard. I also must acknowledge the Earthbyte Group, University of Sydney,

Australia. I have downloaded few very important datasets from their repository, which is freely

available for research.

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My deepest sense of gratitude is reserved for Nemčok family in Bratislava. Apart from Michal, I am also

indebted to Renata, Evana and Jan. I fondly remember Renata and her kids as the best ever host,

possible in a foreign country. I sincerely remain thankful to them for their wonderful favour, which kept

me comfortable during my many visits to Bratislava.

Words cannot express the feelings I have for my parents, my sister and my in-laws for their constant

unconditional support. Special thanks are also due to my father Santatan Sinha, who always wanted me

to peruse a doctorate study, which I think I have finally fulfilled.

I would like to acknowledge the most important person in my life – my wife Tanima. She has been a

constant source of strength and inspiration. There were times during the past four years when

everything seemed hopeless and I didn’t have any hope. I can honestly say that it was only her

determination and persistent encouragement that ultimately made it possible for me to see this project

through to the end.

Finally, very special appreciations are always there for the sweetest and dearest person in my life-

Ritoshree. My beautiful daughter, without her was it possible that the thousand flowers would bloom?

To all of my family members

for making me who I am

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Contents

Page no

Bibliographic record (Czech) iii

Bibliographic record (English) v

Candidate’s declaration vii

Abstract (Czech) ix

Abstract (English) xi

Copyright xiii

Acknowledgements xv

List of figures xxv

List of tables xxxv

Chapter 1: Introduction 1

1.1 Microcontinents 1

1.1.1 Plume refocussing model 2

1.1.2 Competing rift zones and wound linkage hypothesis 4

1.2. Elan Bank Microcontinent 5

1.3. Plate reconstruction between India-Antarctica-Australia 7

Chapter 2: Plate tectonic evolution and regional geological settings 15

2.1 Study Area 15

2.2 Breakup of Gondwana and Plate tectonic evolution 15

2.3 East Indian passive margin 20

2.3.1 Physiographic Setting 20

2.3.2 Geological and Tectonic Setting 22

2.3.1.1 Precambrian Shield Regions 22

2.3.1.2 Permian-Triassic Intra-cratonic Rift 29

2.3.1.3 Jurassic-Lower Cretaceous passive margin 32

2.3.1.4 Cretaceous flood basalt provinces 38

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2.3.1.5 Late Cretaceous to present day passive margin 40

2.4 Kerguelen Plateau and Elan Bank microcontinent 43

2.4.1. Physiography and plate boundary setting 43

2.4.2 Kerguelen volcanic province 46

2.4.3 Crustal structure of Kerguelen Plateau 47

2.4.4 Tectonic Provinces 50

2.5 East Antarctica Passive Margin 58

2.5.1 Physiographic setting 58

2.5.2 Geological and Tectonic Setting 60

2.5.2.1 Precambrian Shield Region 61

2.5.2.2 Permian-Triassic Intra-cratonic rift setting 63

Chapter 3: Data 73

3.1 Reflection seismic data 73

3.1.1 East Indian margin data 74

3.1.2. Elan Bank data 103

3.2. Potential field data 118

3.2.1 Gravity Maps 118

3.2.2 Gravity Profiles 124

3.2.2 Magnetic Profiles 130

3.3 Borehole data 130

3.4. Petrological data from literature 136

Chapter 4: Methods 141

4.1 Seismic Interpretation 141

4.1.1 Seismic and sequence stratigraphic interpretation 142

4.1.2 Structural interpretation 144

4.2 Potential field data interpretation 148

4.2.1 Qualitative Analysis 148

4.2.2 Quantitative Analysis 152

4.2.2.1 Forward modeling along profiles 152

4.3 Gathering petrological data on Elan Bank/Antarctica/East India for geological correlation 154

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4.4 Tectonic reconstruction between Elan Bank and East India 155

4.5 Seismic profiles marriages from conjugate margins 160

4.6 Tectonic timing 160

4.6.1 Break-up timing based on borehole data 161

4.6.2 Break up timing based on seismic interpretation 162

4.6.3 Break-up timing based on magnetic stripe anomaly data from literature 163

4.6.4 Hot-spot activity timing for exact locations 165

4.7 Match of Elan Bank with hot-spot track in Bay of Bengal 165

Chapter 5: Interpretations 167

5.1. Seismic Interpretation 167

5.1.1 East India 167

5.1.1.1 Crustal Architecture 167

5.1.1.2 Rifting style, fault geometry and nature of continental breakup 172

5.1.1.3 Margin segmentation 177

5.1.1.4. Hyper-extended crustal architecture model 179

5.1.2 Elan Bank 182

5.1.2.1 Crustal architecture and rifting style 183

5.1.2.2 Nature of continental breakup and margin segmentation 186

5.2. Potential field data interpretation 188

5.2.1. Qualitative Interpretation 188

5.2.1.1 East India 188

5.2.1.2 Elan Bank 189

5.2.2 Quantitative Interpretation 194

5.2.2.1 East India 194

5.3 Interpretation of Petrological Data 200

5.3.1 Textural and compositional analysis 200

5.3.2 Geochemical analysis 201

5.3.3 Isotopic and radiometric analysis 202

5.4 Tectonic reconstruction of Elan Bank and East India 204

5.4.1 Structural architecture of different margin segments 206

5.4.2 Crustal architecture of margin segments and geometry of crustal boundaries 208

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5.5 Seismic profile marriages 212

5.6 Tectonic timing 216

5.6.1 Breakup timing 216

5.6.2 Hot spot timing 228

5.7 Geodynamic evolution and release of Elan Bank microcontinent 229

5.8 Match of Elan Bank with hot-spot track in Bay of Bengal 238

Chapter 6: Discussions 241

6.1 Elan Bank microcontinent in India-Antarctica plate reconstruction models 241

6.2 The physiographic boundary and crustal architecture of continental 248

Elan Bank microcontinent

6.3 The importance of the East Indian crustal architecture 251

and ocean-continent transition in reconstruction with conjugate Elan Bank

6.4 Time constraints for tectonic events and rift evolution at hyper-extended margins 258

6.5 Elan Bank microcontinent release mechanism 263

Chapter 7: Conclusions 281

7.1 Conclusions 281

7.2 Future Works 284

References 287

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List of Figures

Page No

Figure 1.1 Conceptual model for plume-related microcontinent formation 3

Figure 1.2 Schematic model of rifting evolution illustrating the formation of the different categories of crustal blocks

4

Figure 1.3 Location Map of the study area and outline of the Elan Bank microcontinent on the satellite free air gravity map

5

Figure 1.4 ODP well-1137drilling result 6

Figure 1.5 The previously published seismic interpretation of line S179-05 7

Figure 1.6 Tight assembly of Eastern Gondwana at 145 Ma, incorporating Precambrian crustal fragments

8

Figure 1.7 Permian-Triassic reconstruction of India-Antarctica by Harrowfield et al, 2005

9

Figure 1.8 Reconstructed Precambrian geology of Africa, India, East Antarctica and Australia

10

Figure 1.9 Plate reconstruction models for early breakup between India and Antarctica by Gaina et al., 2007

12

Figure 1.10 Plate reconstruction models for early breakup between India and Antarctica by Veevers, 2009.

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Figure 2.1 Map of the study area 16

Figure 2.2 Entire Gondwana reconstruction at 250 Ma 17

Figure 2.3 The distribution of Permo-Triassic intra-cratonic Karoo rift basins, 17

Figure 2.4 The four phases of Gondwana breakup and drift history 19

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Figure 2.5 Physiographic setting of East India passive margin 21

Figure 2.6 The onshore geology of East India passive margin 23

Figure 2.7 Precambrian geology of the onshore East India 25

Figure 2.8 Distribution of stable composite, consolidated rigid Archean cratonic blocks

26

Figure 2.9 Simplified geological map of the Southern Granulite Terrain 27

Figure 2.10 Distribution of Permo-Triassic rocks within intra-cratonic rift systems 31

Figure 2.11 Map of the major tectonic segments of the study area and structural features of the adjacent onshore

34

Figure 2.12 Tectonic element map of the Cauvery Basin. 35

Figure 2.13 Onshore and offshore basement trends in the Krishna-Godavari Basin 37

Figure 2.14 Tectonic map of Mahanadi basins 38

Figure 2.15 Mafic magmatic activity in the Mahanadi Basin in Cretaceous 39

Figure 2.16 The distribution map of Deccan Volcanic Province and equivalent rocks in Peninsular India

41

Figure 2.17 Plot of significant geological phenomena recorded in rocks of southern Tibet and elsewhere in the region

42

Figure 2.18 The present day plate settings in the southern Indian Ocean 44

Figure 2.19 Interpretation of gravimetric and magnetic maps associated with bathymetry

45

Figure 2.20 Estimated Kerguelen magma output since 130 Ma 46

Figure 2.21 Locations of geological sampling sites, ODP wells and sonobuoy stations in Kerguelen Plateau

48

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Figure 2.22 Best fit velocity model across the northern Kerguelen Plateau from 1D inversion of travel times and Final model across southern Kerguelen Plateau

49

Figure 2.23 Tectonic provinces of the Kerguelen Plateau and adjacent areas 51

Figure 2.24 Composite logs for ODP Sites 1139, 1140, 736, 737, 1138, 1137 within Kerguelen Plateau

52

Figure 2.25 Detail of seismic data at ODP Site 1137 55

Figure 2.26 Summary of drill sites from the Kerguelen Plateau and Broken Ridge that recovered volcanic rocks

56

Figure 2.27 The interpreted seismo-geological cross section across the Labuan Basin and South Kerguelen Plateau

57

Figure 2.28 Physiographic division of Antarctic continent 59

Figure 2.29 Map of East Antarctica in its reconstructed Gondwana context 60

Figure 2.30 Schematic map of different tectono-stratigraphic and geological provinces of Antarctica

62

Figure 2.31 Map of East Antarctica showing the regions with Precambrian crust of different ages

62

Figure 2.32 Generalized geologic map of the Prydz Bay area showing the location of the Lambert Graben

64

Figure 2.33 Summary of the stratigraphy, measured sections and lithology of the Amery Group in the Beaver Lake area

65

Figure 2.34 Apatite fission track samples from the Lambert graben 67

Figure 2.35 Tectonic elements of the continental margin of East Antarctica 68

Figure 2.36 Seismic profile GA-229/35 through the offshore western Enderby Land (Stagg et al., 2005). Arrow shows an inboard edge of oceanic crust.

69

xxviii

Figure 2.37 Interpreted seismic profiles through the offshore Antarctica near Terre Adélie rift block

70

Figure 2.38 Potential field model for seismic profile through the offshore Antarctica near Terre Adélie rift block

71

Figure 3.1 The reflection seismic Profile-1 through the East Indian margin 75

Figure 3.2 The reflection seismic Profile-2 through East Indian margin 77



Figure 3.5 ION reflection seismic profile 800 through East Indian margin 81




Figure 3.9 ION reflection seismic profile 900 through East Indian margin 88




Figure 3.13 The Refection seismic Profile-11 through East Indian margin 95

Figure 3.14 Map of the study area, offshore East India, showing the location of the IndiaSpan reflection seismic profiles

99

Figure 3.15 ION-Reflection seismic profile 1000 in East India margin 100


xxix


Figure 3.18 The 2D seismic grid in Elan Bank 104

Figure 3.19 The Reflection seismic Profile-1 through the central Kerguelen Plateau 105

Figure 3.20 The Reflection seismic Profile-2 through the north-western margin of Elan Bank.

107

Figure 3.21 The Reflection seismic Profile-3 through the north-western margin of the Elan Bank

110

Figure 3.22 The Reflection seismic Profile-4 through the south-western margin of Elan Bank

111

Figure 3.23 The Reflection seismic Profile-5 through the southern margin of Elan Bank

114

Figure 3.24 The Reflection seismic Profile-6 through the central part of Elan Bank 115

Figure 3.25 The Refection seismic Profile-7 across the Elan Bank through the Kerguelen Plateau and the Labuan Basin

116

Figure 3.26 Free-Air gravity map of the East Indian offshore 119

Figure 3.27 Topography of the Elan Bank and surrounding Kerguelen Plateau 120

Figure 3.28 Free air gravity map of the Elan Bank 121

Figure 3.29 10 km High pass Bouguer gravity map of East India offshore 123

Figure 3.30 Bouguer gravity map of Elan Bank 124

Figure 3.31 Isostatic residual anomaly map of East India offshore 125

Figure 3.32 Bouguer gravity and magnetic anomaly curves along ION IndiaSpan profiles

127

Figure 3.33 Calculated Bouguer anomaly curves along S179 survey profiles in the Elan Bank and the southern Kerguelen Plateau

128

xxx

Figure 3.34 The well data in East India showing a distribution of paleo-environments 133

Figure 3.35 Composite stratigraphic section of ODP Site 1138 134

Figure 3.36 Composite stratigraphic section of ODP Site 1137, which is located on the Elan Bank and tied to seismic Profile-6 (Coffin et al., 2000).

135

Figure 3.37 The fluvial samples from ODP -1137 site. 137

Figure 3.38 Radiometric age dating and comparison of continental basement origin of India and Elan Bank.

138

Figure 3.39 Radiogenic isotopes of selected samples from ODP site 1137 139

Figure 3.40 Field photographs of different lithounits of the Eastern Ghat mobile belt from the Nagavalli shear zones

140

Figure 4.1 Stratal terminations that can be observed above or below a stratigraphic surface in seismic profiles and larger scale outcrops

143

Figure 4.2 Architecture of facies, genetic units (systems tracts) and sequence stratigraphic surfaces

144

Figure 4.3 Segment of the PROBE profile 5 showing the occurrence of the “rift indicators” on the proto-oceanic crust

145

Figure 4.4 A seismic profile showing the continental crust in detail in East India margin

147

Figure 4.5 The gravity data processing workflow. 149

Figure 4.6 Fit of the continents around the Atlantic Ocean 156

Figure 4.7 Sea floor spreading and the generation of magnetic lineations 157

Figure 4.8 Ocean floor age map 159

Figure 4.9 Progressive extension along major basin bounding fault and sedimentary deposition

163

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Figure 4.10 Different magnetic anomaly datasets in Indian Ocean combined from various published sources

164

Figure 5.1 Tectonic element map of East India passive margin 174

Figure 5.2 Hyper extended crustal architecture model of the Krishna-Godavari Basin 180

Figure 5.3 Hyper extended crustal architecture model of the Krishna-Godavari Basin 181

Figure 5.4 Hyper-extended crustal architecture model of the Palar Basin 182

Figure 5.5 Hyper-extended crustal architecture model of Elan Bank 187

Figure 5.6 Qualitative seismic-gravity integrated interpretation through profile-2 showing possible crustal architecture on Elan Bank

191

Figure 5.7 Qualitative-seismic-gravity integrated interpretation through profile-3 showing possible crustal architecture on Elan Bank

192

Figure 5.8 Qualitative-seismic-gravity integrated interpretation showing possible crustal architecture on Elan Bank

193

Figure 5.9 Qualitative-seismic-gravity integrated interpretation showing possible crustal architecture on Elan Bank

193

Figure 5.10 Gravity and magnetic modeling along the reflection seismic section ION-GXT-1000 in East India

195

Figure 5.11 Gravity and magnetic modeling along the reflection seismic section ION-GXT-1000 (without POC) in East India

196


197


198


198

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199


199

Figure 5.17 High resolution seismic-geological interpretation of ODP-Well -1137 site on the Profile-6 at Elan Bank

200

Figure 5.18 Multi element variation diagram for volcanic clast and non-volcanic clast normalized to primitive mantle composition

202

Figure 5.19 Isotopic analysis to compare the granite clasts, sampled from Eastern Ghat, Indian carton and Indian shield region

203

Figure 5.20 India–Elan Bank reconstruction 205

Figure 5.21 The calculated paleo-stress analysis from meso-scale fracture data from outcrops all along the East India passive margin

207

Figure 5.22 East India margin crustal boundary map 209

Figure 5.23 Tectonic element map of Elan Bank 210

Figure 5.24 Map showing the locations of the India –Elan Bank conjugate margin paired seismic section locations

213

Figure 5.25 Elan Bank reconstruction at 123 Ma (Transect A-A’) 214

Figure 5.26 India –Elan Bank reconstruction at 123 Ma (B-B’) 215

Figure 5.27 Interpreted responses and timings at different margin settings in northern Krishna-Godavari rift zone.

218

Figure 5.28 Interpretation of tectonic timings from paleo-environment chart 219

Figure 5.29 A drilled well shows paleo-environments and age controls in the proximal part of northern Krishna-Godavari rift zone.

220

Figure 5.30 Subsidence profiles in northern rift zone calculated at 1D location of wells 221

xxxiii

T2, P1 and D1.

Figure 5.31 The structural element map in north Krishna-Godavari basin with two clear fault trends

222

Figure 5.32 A drilled well shows paleo-environments and age controls in the proximal part of southern Krishna-Godavari rift zone.

223

Figure 5.33 Subsidence profiles in the central Coromandal strike slip system calculated at 1D location of well T9

224

Figure 5.34 Time thickness map K20-K30 is showing relative change in post breakup dynamic topography change as a result of isostatic uplift

225

Figure 5.35 Time thickness map K30-K60 is showing relative change in post breakup dynamic topography change as a result of isostatic uplift

226

Figure 5.36 Hotspot track created in Bay of Bengal as a result of migrating the plate over Kerguelen plume

229

Figure 5.37 Schematic kinematic evolution model for India-Antarctica passive margin and microcontinent release

231

Figure 5.38 The Elan Bank reconstruction at Bajocian (168Ma) 232

Figure 5.39 The Elan Bank reconstruction at Beriassian-Valanginian (140 Ma) 233

Figure 5.40 The Elan Bank reconstruction at Valanginian-Hauterivian (132 Ma) 234

Figure 5.41 The Elan Bank reconstruction at Aptian (123 Ma) 236

Figure 5.42 The Elan Bank reconstruction at Albian-Aptian (112 Ma) 238

Figure 5.43 The relationship between Elan Bank drift path and hotspot track 239

Figure 6.1 125 Ma reconstruction of the Indian Ocean with the Enderby Basin model

243

Figure 6.2 145 Ma tight rigid plate reconstruction of the Indian ocean 244

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Figure 6.3 145 Ma tight rigid plate reconstruction of the Indian Ocean 245

Figure 6.4 120 Ma reconstructions of the southern Indian Ocean by Gibbons et al., 2013

247

Figure 6.5 1 min Satellite-derived free-air gravity map for the Enderby Basin 263

Figure 6.6 Conceptual model of magmatic heating of the lithosphere 266

Figure 6.7 Conceptual model of the Elan Bank microcontinent formation based on plume refocusing hypothesis

268

Figure 6.8 The seismic interpretation shows Moho imaging in offshore Antarctica 270

Figure 6.9 Model of uplift above a plume head 271

Figure 6.10 Competing rift zone hypothesis model for the Elan Bank microcontinent formation

273

Figure 6.11 Crustal architecture and continental breakup of the Jan Mayen Microcontinent

275

Figure 6.12 Kinematic evolution of the Norway Basin and the Jan Mayen microcontinent illustrated by a series of key plate reconstruction stages

276

Figure 6.13 Schematic sequence of panels, showing the evolution of the North Atlantic ridge system and development of the Jan Mayen microcontinent

278

Figure 6.14 Tectonic evolution of the Seychelles microcontinent 276

Figure 6.15 Tectonic map of East Africa 280

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List of Tables

Page no

Table 3.1 Tectono-stratigraphic chart for East coast India used for identification of seismic sequence boundaries.

96

Table 3.1 Tectono-stratigraphic chart for Elan Bank used for identification of seismic sequence boundaries. References

117

Table 4.1 Density used for gravity forward modeling 153

Table 5.1 Width of the different interpreted crustal domains in East India from integrated seismic interpretation and potential field modeling

211

Table 5.2 Width of the different interpreted crustal domains in Elan Bank from integrated seismic and potential field interpretation

212

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1

Chapter 1

Introduction

1.1 Microcontinents

The microcontinents are continental fragments located within deep oceanic regions with an extent of

less than 106 km2 (Rey et al., 2003). Typical characteristics of a microcontinent includes bathymetric

highs, positive free-air anomalies, magnetic quiet zones, prograding sedimentary sequences on the

flanks, continental crustal structures, being older than surrounding crust and different heat flow

signatures (Rey et al., 2003). Their development is fairly common in geologic history of different passive

margins. Typical examples of microcontinents include Jan Mayen in North Atlantic (Müller et al., 2001;

Rey et al., 2003), Elan Bank in Southern Indian Ocean (Borissova et al., 2003; Gaina et al., 2007) and

Seychelles in Indian Ocean (Dyment, 1998; Müller et al., 1998) etc.

The passive margins associated with microcontinent development usually had gone through poly-phase

rifting (Manatschal et al., 2007) and multiple breakup processes (Péron-Pinvidic and Manatschal, 2010).

As a result of that they have different thermal histories and, in many cases, direct influence of mantle

plumes complicates their thermal evolution. The most common phenomenon, which is observed during

microcontinent separation, is the ridge jump. It has been observed that under certain conditions there is

reorganization in sea-floor spreading, where active spreading is gradually overtaken by a newly formed

spreading center. This is known as ridge jump (Goff and Cochran, 2006). The ridge jump is a critical

interplay between plate motion, nature of continental extension, involvement of competing rift zones

and possibly an interaction with hotspots. Hot-spot activity during ridge jump is commonly associated

Chapter-1 Introduction

2

with bulk magmatism. As a result, it leaves a prominent imprint on the thermal history of any passive

margin. Therefore, the study of microcontinents can provide critical insights into many unresolved

problems on conjugate margins, including plate reconstructions, role of competing rift zones in

continental breakup processes, hotspot volcanism and associated thermal history.

There are several hypotheses explaining a ridge jump. The primary ones are plume assisted ridge jump

(Brozena and White, 1990; Mittelstaedt et al., 2008) and ridge jump due to irregular ridge propagation

(Sempéré et al., 1995; Searle et al., 1998). The plume assisted ridge jump may occur as a result of

lithospheric tension induced by buoyant, convective asthenosphere (Mittelstaedt and Ito, 2005),

thermo-mechanical thinning of the lithosphere by the ascending plume (Jurine et al., 2005) and

reheating of the lithosphere due to magmatic penetration through the plate (Kendall et al., 2005).

It is Interesting to note that majority of the researchers distinctly agree that ridge jump is a ‘must’

phenomena, which causes a microcontinent separation. However, the question remains that, what

mechanism halts the existing sea floor spreading process, while causing the creation of another new

spreading center within the stretched continental crust? Few models have been put forward to

understand the mechanisms involved in microcontinent initiation and development. Synthesizing the

observations from different microcontinents, kinematics and numerical modeling, two major

hypotheses for the microcontinent developement were considered.

1. Plume refocussing model (Müller et al., 2001), and

2. Competing rift zones and wound linkage hypothesis (Péron-Pinvidic et al., 2010)

1.1.1 Plume refocussing model

This model was proposed by Müller et al., (2001), which directly require a mantle plume to release a

microcontinent (Figure 1.1). This model is built upon active rift model and a hypothesis by Steckler and

ten Brink (1986). The hypothesis states that minimum yield strength occurs along the edges of a rifted

margin, because the crust there is already weakened and hotter than normal crust, owing to a

conductive heat flow from the adjacent rift (Müller et al., 2001). Therefore, when the newly-formed

passive margin moves over a hotspot, it facilitates renewed rifting at the edge of the margin. The

renewed rifting under the influence of the hotspot causes another continental breakup. Subsequently,


3

the active ridge migrates towards hotspot causing the ridge-jump. This gradually isolates the

microcontinent, which later becomes a part of newly formed plate. The ridge-jump toward hotspot

causes further asymmetries in oceanic crustal accretion, resulting in excess accretion and series of

extinct ridges, even within the microcontinent (see Müller et al., 2001; Gaina et al., 2003 for details).

Figure 1.1: Conceptual model for plume-related

microcontinent formation (Müller et. al, 2001)

Fragmentation of two large plates has caused a

young ocean basin to form (A). A young passive

margin moves into the vicinity of an existing

mantle plume, enhancing the yield strength

minimum landward of the rifted margin by

conductive heating (B, C). This results in renewed

rifting, followed by a jump of the active ridge

towards the hot spot, and the isolation of the

microcontinent (D). Further ridge jumps towards

the hot spot cause subsequent asymmetries in

crustal accretion, resulting in excess accretion

and a series of extinct ridges on the plate

including the microcontinent, if the hot spot

remains under the opposite ridge flank (D). If the

ridge crosses the hot spot, the sign of

asymmetries in crustal accretion would switch,

causing excess accretion on the ridge flank

opposite to the microcontinent.


4

1.1.2 Competing rift zones and wound linkage hypothesis

This hypothesis suggest that during final stages of rifting, several rift zones simultaneously thin the

lithosphere in various places forming several weak zones known as crustal wounds (Collier, et al., 2008;

Péron-Pinvidic and Manatschal, 2010; Nemčok et al., 2012a; Figure 1.2). The kinematic competition

between each crustal wounds that want to host the crustal breakup eventually forms a breakup

trajectory by establishing the linkage between these rift zones or the wounds (Péron-Pinvidic and

Manatschal, 2010; Nemčok et al., 2012a). Once an initial breakup trajectory is established, it is followed

by a discrete early sea-floor spreading. However, depending on the structural, compositional and

thermal inheritance and the strain rate, sometimes the rifting does not stop but further propagates

within some of the competing rift zones (Péron-Pinvidic and Manatschal, 2010). Therefore, this event

causes reorganization in previously attempted breakup locations. This additional competition between

these propagating rift zones establishes a renewed linkage among them and successively forms a final

breakup trajectory. As a result, the continental block, which is trapped between initial and final breakup

trajectories, departs with one of the newly-formed plates becoming a microcontinent.

Figure 1.2: Schematic model of rifting evolution illustrating the formation of the different categories of crustal blocks (Péron-Pinvidic and Manatschal, 2010). The genesis, evolution, final shape and position of each block in the margin are related to the distinct modes of deformation affecting the margin during rifting. The competition between different propagating rifts to host continental breakup and an unsuccessful attempt may generate microcontinents. However, this model is still evolving.


5

1.2. Elan Bank Microcontinent

One of the most conspicuous microcontinents in the Indian Ocean is the Elan Bank (Wise et al., 2000;

Coffin et al. 2000; 2002; Borissova, et al., 2003; Gaina et al., 2007), which is situated inside the southern

Kerguelen Plateau (Figure 1.3). Elan Bank is characterised by an east to west trending promontory within

the Southern Kerguelen Plateau. It is a relatively positive bathymetric feature compared to the

surrounding oceanic basins and a positive anomaly feature in the Free-Air gravity map (see figure 2 in

Coffin et al., 2002). The Elan Bank is currently submerged about 500 m below the sea level. It is about

100 to 200 km wide in north-south direction and trends east-west for about 600 km. The microcontinent

is located about 900 km to the north of the Antarctic margin, and encompasses approximately 140,000

km2 (Borissova, et al., 2003).

Figure 1.3: Location Map of the study area. SRTM topography and bathymetry (Sandwell and Smith, 2009) map shows the geographic distribution of different continents, important ocean basins, plate boundaries and physiographic provinces. KP=Kerguelen Plateau, EB=Elan Bank, SEIR=South-East Indian Ocean Ridge, SWIR=South-West Indian Ocean Ridge, CR=Charlesberg Ridge, ECMI=Eastern continental margin of India. Outline of the Elan Bank microcontinent on the satellite free air gravity map (Coffin et al., 2002), It is a relatively large micro-continent located inside the southern Kerguelen Plateau in the Indian Ocean.


6

The continental nature of Elan Bank has been proven by many studies including potential field

investigations (Coffin et al., 1986, Gaina et al. 2003), seismic imaging (Borissova et al., 2003), velocity

analysis (Charvis et al., 1995; Operto and Charvis, 1996; Charvis, and Operto, 1999; Borissova et al.,

2003) and petrological analysis (Nicolaysen et al., 2001; Weis et al. 2001; Ingle, 2002). Initially, it was

considered to be part of the Kerguelen Large Igneous Province (Royer and Coffin, 1992). Petrological

analyses were carried out on the fluvial conglomerate samples recovered from the ODP well-1137,

which is located on the central part of the microcontinent (Figure 1.4). The textural, compositional and

geochemical analysis of the various clasts found in these intra-formational conglomerates confirms the

close proximity of continental crust, which was the primary provenance for these types of clasts

(Nicolaysen et al., 2001; Weis et. al, 2001; Coffin et al., 2002; Ingle, 2002).The isotopic analysis and

radiometric dating of these clasts further suggests that they are sourced from Neoproterozoic and

Archean rocks of Eastern Ghat mobile belts in East India (Weis et. al, 2001). Hence, a possible correlation

between the East India and the Elan Bank can be drawn based on these evidences (Li and Powell, 2001;

Nicolaysen et al., 2001). This indicates that the Elan Bank was once attached to East India until its

breakup and separation that took place during Early Cretaceous.

Figure 1.4: Confirmation of the continental nature of the Elan Bank from ODP site -1137 drilling result (Ingle et al., 2002). Clasts within fluvial conglomerate to pebble have variable lithology including alkali basalt, rhyolite, granite-gneiss and other metamorphic rocks. The gneisses, sandstones and conglomerate matrix contain zircon and monazites with Neoproterozoic and Archean ages. This suggests the proximity of continental crust during formation of conglomerates.


7

The interpreted crustal structure of Elan Bank through seismic imaging by Borissova et al., (2003) shows

possible rift architecture (Figure 1.5). It, suggests a possible presence of rift structure on the Elan Bank

side conjugate to the East Indian margin. These new results still have to be explained and integrated

with the exiting plate tectonic reconstructions and hyper-extended passive margin architecture.

Figure 1.5: The seismic interpretation of line S179-05 (Borissova et. al, 2003). The line is located within the Enderby Basin, which is located on the side of Elan Bank facing Antarctica. The interpretation shows possible rift architecture along line S179-05. Therefore, it suggests a possible presence of rift structures on the Elan Bank side facing East Indian margin.

1.3. Plate reconstruction between India-Antarctica-Australia

A series of plate tectonic reconstruction model exist to describe pre-breakup Gondwana continental

landmass. They are based on most of the established methods of plate reconstructions. Typical

examples include geometric reconstructions (Reeves, et al., 2002; Reeves, 2009), integrated geological


8

reconstructions (Harrowfield et al., 2005), geological and paleo-magnetic methods (Collins and

Pisarevsky, 2005) magnetic anomaly and hotspot track based reconstruction (Gaina et al., 2007) and

palinspastic reconstruction (Veevers, 2009).

The geometric reconstruction (Reeves, et al., 2002; Reeves, et al., 2004; Reeves, 2009) were carried out

based on continental geology and by reversing seafloor spreading, based on detailed ocean-floor

topography through satellite altimetry to achieve a tight assembly. This model provides a relatively good

process to retrace ocean-floor creation. The series of tectonic events that describe Gondwana breakup

has been worked out with substantial geometrical details. One of the primary notions of this

reconstruction assumes that head of the Napier peninsula grossly matches the bight of India, located

within Krishna Basin with a fit of the 1000-fathom isobaths (Figure 1.6). However, according to Veevers

(2009) this pre-rift fit is overtight by 135 km along trends that parallel the flow-lines of the early phase

of seafloor spreading.

Figure 1.6: Tight assembly of Eastern Gondwana at 145 Ma, incorporating Precambrian crustal fragments (modified after Reeves, 2008). The reconstruction indicates that head of the Napier peninsula grossly matching the bight of India, located within Krishna Basin with a fit of the 1000-fathom isobaths. Note that the reconstruction has no space to fit Elan Bank microcontinent between India and Antarctica.


9

The geological reconstruction supported by revised transcontinental drainage system and coal bearing

sequences of Mahanadi Basin, Pranhita-Godavari Basin and Lambert graben, Harrowfield et al, (2005)

suggests Permian separation of the Indian and Antarctic platforms and across-rift alignment of the

Lambert graben and Godavari Valley. Thus providing a geological fit between Pranhita-Godavari Basin

and Lambert graben, the entire reconstruction has been carried out (Harrowfield et al, 2005; Figure 1.7).

Instead of creating a gap between India and Antarctica, where one could fit Elan Bank, this

reconstruction shows a distinct overlap. Furthermore, the fit between continental crust and oceanic

crust boundaries has not been tested in this model.

Figure 1.7: Permian-Triassic reconstruction of India-Antarctica (Harrowfield et al, 2005). This reconstruction shows alignment of Lambert graben and Godavari Valley and compartmentalization of intra-cratonic rift. This reconstruction involves a serious overlap to fit Elan Bank is between India and Antarctica. The fit between continent ocean boundaries has not been tested in this model.


10

Collins and Pisarevsky (2005) created a non-unique, but permissive model for the reconstruction of

Gondwana Landmass and a fit between India-Antarctica and Australia by integrating the petrological,

geochrological data for India-Antarctica-Australia landmass. Furthermore, the orientations of

Neoproterozoic shear zones were combined with the available palaeomagnetic data in this model

(Figure 1.8). Although it is a modification of original tight Gondwana assembly of Reeves and DeWit

(2000), it does not incorporate Elan Bank.

Figure 1.8: Geological outcrop map of south, central and northeast Africa, India, East Antarctica and Australia (Collins and Pisarevsky, 2005) rotated into the Gondwana tight assembly of Reeves and De Wit (2000). Precambrian outcrops older than 1000 Ma shown as dark grey and those younger than 1000 Ma are shown in light grey. Outlines of the Neoproterozoic continental blocks are marked in purple. Block abbreviations: A-A=Afif–Abas Terrane; Az=Azania; Congo=Congo/Tanzania/Bangweulu Block; L-V=Lurio–Vijayan Peninsula; R Plata=Rio de la Plata Block; Ruker=Ruker Terrane, Southern Prince Charles Mountains; S Fran=Saô Francisco Block; WA=West African Block. Orogenic belt and location abbreviations: Alb-Fr=Albany–Fraser Orogen; ANS=Arabian–Nubian Shield; D Feliciano=Dom Feliciano Belt; Cam=Cameroon; L=Leeuwin Complex; M=Mulingarra Complex; MB=Mozambique Belt; N=Northampton Complex; Of=Officer Basin, Sey=Seychelles.


11

Most notably, none of these reconstruction models include Elan Bank, which is a sizeable

microcontinent and whose continental nature was established early through for those reconstructions

(Wise et al., 2000; Coffin et al. 2000; Coffin et al. 2002; Borissova, et al., 2003).

Gaina et.al (2007) provides a comprehensive tectonic reconstruction of the breakup and early seafloor

spreading between India and Antarctica based on a robust coverage of gravity, magnetic and seismic

data off the east Antarctic margin (Figure 1.9). This reconstruction includes Elan Bank as a

microcontinent as a part once attached to East India passive margin, which was fitted against northern

part of East India passive margin, e.g. close to Mahanadi basin. Veevers (2009) used a palinspastic

reconstruction with quantitative elimination of the intervening ocean floor. This reconstruction model

uses, seismic and gravity magnetic data to determine the continent ocean boundary (COB) and thickness

of the extended continental crust. This model also provides geochronological connections across the

India-Antarctica breakup loci. Veevers (2009) suggests an area approximately 550 km * 50 km gap could

be the location of Elan Bank microcontinent (Figure 1.10). However, the proposed location of Elan Bank

again suggests proximity to Mahanadi basin. Moreover, it does not match with the reported geometry

of Elan Bank microcontinent (Wise et al., 2000; Coffin et al. 2000; Coffin et al. 2002; Borissova, et al.,

2003).

The above discussion clearly suggests that the size and shape of Elan Bank is a matter of debate and its

placement against East India is not yet deterministic. One of the major reasons behind this is linked to

unavailability of data and interpretation, which is required to have a good handle on crustal architecture

and basin evolution history and nature of continental breakup of East India. Published plate motion

models do not show any ridge jumps because the oceanic crust between Antarctic and Kerguelen

Plateau formed during the long Cretaceous Normal Superchron. Therefore it cannot be dated using the

usual technique of marine magnetic anomaly identification. However, it looks like a careful investigation

integrating potential field data, seismic and petrological data, being the core of this thesis should help to

place Elan Bank using detail tectonic criteria and such study should allow seeing its development

mechanism.


12

Figure 1.9: Plate reconstruction models for early breakup between India (IND) and Antarctica (ANT) based on new seafloor spreading isochrons between the Gunnerus Ridge and Bruce Rise, including the subsequent ridge jump to north of Elan Bank (EB) and growth of the Kerguelen Plume (KP) (Gaina et al., 2007). Stars indicate hotspots. Note that this reconstruction includes Elan Bank but fits closely towards Mahanadi Basin in India. (BB: Bunbury Basalts; BR: Bruce Rise; NP: Naturaliste Plateau; Md: Madagascar; RJ: Rajmahal Traps and Sri: Sri Lanka).


13

Figure 1.10: Antarctica and conjugate India fitted together such that the palinspastically restored continent–ocean boundaries (Solid line: COB–Antarctica, Dotted Line: COB-India) (Veevers, 2009). The match is divided into four sectors from A to D. This reconstruction does not restore Elan bank. However, it indicates that Elan Bank should be placed against India between latitude 15°N to 20°N, which covers an area approximately 550 km * 50 km. (AU=Achankovil Unit; BGB=Bhopalpatnam Granulite Belt; CGGC=Chotanagpur Granite Gneiss Complex; CITZ=Central Indian Tectonic Zone; COB'=restored continent–ocean boundary; HC=Highland Complex; KGB=Karimnagar Granulite Belt; N-EGMB=northern Eastern Ghats Mobile Belt; PB=Prydz Bay; PCM=Prince Charles Mountains; RKG=Ramakona–Katangi Granulite Belt; S-EGMB=southern Eastern Ghats Mobile Belt; SGT=Southern Granulite Terrain; VC=Vijayan Complex; WC=Wanni Complex; YB=Yamato–Belgica Complex.

The objective of the study is to decide at its end, which of the two candidate hypotheses explains the

existence of Elan Bank, its role in India-Antarctica breakup and its reconstructed position with respect to

East India passive margin. These candidate hypotheses are:


14

1. Plume re-focusing hypothesis: Elan Bank separation due to Cauvery rift zone initially capturing

the continental breakup, then Krishna-Godavari rift zone made a better breakup location choice

because of the Kerguelen plume activity influence in a close proximity and a associated ridge

jump.

2. Competing rift zone hypothesis: Elan Bank separation due to Cauvery rift zone making it into the

breakup but Krishna-Godavari rift zone never stopping its propagation. Younger propagation

end of the Krishna-Godavari zone eventually linking via Coromondal transfer zone with inner

portion of the Cauvery rift zone instead of initial linkage between end-Cauvery zone and inner

Krishna-Godavari zone.

The study incorporates reflection seismic data, potential field data (gravity and magnetic), petrological

data from the public domain, few key borehole data (ODP and exploratory wells). Regarding methods,

the study includes seismic interpretation, potential field data interpretation and modeling, gathering

petrological data from literature on Elan Bank-Antarctica-East India matches and reconstructs the

position of Elan Bank to its conjugate East India. It has to be noted that it has not exactly used the global

plate reconstruction models but a synthetic reconstruction based on detailed understanding of crustal

architecture and nature of ocean continent transition integrated with global plate reconstruction

models based on geometric method and magnetic anomalies. Furthermore, it makes seismic marriages

to test the precision of the reconstructed position and a detailed analysis on tectonic timings.

Using the above mentioned datasets and methods, the thesis decides at its end on which one of the

proposed hypotheses is correct, subsequently rejecting the remaining one.

15

Chapter 2

Plate Tectonic Evolution and Regional Geological Setting

2.1 Study Area

The study area encompasses the following areas located in and around Indian Ocean (Figure 2.1);

1. East Indian passive margin

2. Elan Bank microcontinent within Kerguelen plateau

3. East Antarctica passive margin

2.2 Breakup of Gondwana and Plate tectonic evolution

Pre-breakup Gondwana comprised of several Precambrian fragments in form of a rigid assembly that

lasted for almost 400 Ma since the widespread Pan-African orogeny (Reeves and de Wit, 2000, Figure

2.2). Prior to Gondwana breakup, Permo-Triassic intra-cratonic rift basins of central Gondwana, the

Karoo basins, located in present day southern and central Africa, evolved during the first-order cycle of

supercontinent breakup of Pangea (Catuneanu et al.,2005, Figure 2.3). The Karoo rift basins in the NW-

trending Ubende Belt in Tanzania and East Congo indicate the early stage of breakup in central

Gondwana (Delvaux, 2001). However, the Permo-Triassic tectonic regime was unable to end up in

successful continental breakup. Several other failed Permo-Triassic rift basins exist throughout the

erstwhile Gondwanaland, e.g. Pranhita-Godavari (Lakshminarayana, 2002) and Damodar valley basins

(Chakraborty and Ghosh, 2005) in India, Lambert Basin in East Antarctica (Harrowfield et al., 2005;

Chapter-2 Regional Geology

16

Veevers, 2009), Bunbury and Dandaragan Trough in Perth basin, western Australia (Song and Cawood,

2000).

Figure 2.1: Map of the study area. The area includes East India passive margin, significant part of the Indian Ocean, Kerguelen Plateau, which is a Large Igneous Province (LIP) within Indian Ocean, Elan Bank microcontinent, Enderby Basin and adjacent portion East Antarctica passive margin. The map also shows some other important physiographic features of Indian Ocean like Mid Oceanic ridges (SE and SW Indian Ocean Ridge, Carlsberg Ridge), Broken Ridge, a volcanic plateau, 90°E Ridge, a linear hotspot trail formed the ridge.


17

Figure 2.2: Entire Gondwana reconstruction at 250 Ma is represented as an aggregate of about 50 rigid Precambrian fragments (Reeves, 2009).

Figure 2.3: The distribution of Permo-Triassic intra-cratonic Karoo rift basins, the first order the first-order cycle of supercontinent breakup of Pangaea in south-central Africa (Catuneanu et al., 2005).


18

The breakup of Gondwana was initiated during Late Triassic to Early Jurassic time (Reeves and de Wit,

2000). Most of the extensional events were localized within narrow zones of Pre-existing crustal

discontinuities. The entire Gondwana breakup was achieved through four tectonic phases and

associated continental movements (Reeves et al., 2004). These phases are:

a. Rifting and breakup between Eastern and Western Gondwana (200-167 Ma)

b. Rifting between Africa and Antarctica (190-136 Ma)

c. Rifting and breakup between India-Australia-Antarctica (180-130 Ma)

d. Rifting between India and Madagascar (90 Ma)

Initially, the breakup between Eastern Gondwana (Africa and South America) and Western Gondwana

(India-Australia-Antarctica-Madagascar) was largely controlled by the dextral strike-slip movement along

the Davie and the Lebombo-Explora fracture zones, which define the eastern and western limits of

Africa-Antarctica corridor (Reeves et al., 2004; Figure 2.4a). The earliest seaway around the margin of

Eastern Gondwana developed during the Early Jurassic in Proto-Mozambique and Proto-Somali Ocean.

The separation between greater India and Australia was also preceded by Permo-Triassic intra-cratonic

failed rifting. Rifting between India and Antarctica began sometimes in Middle Jurassic. The Indian plate

rotated away from Antarctica-Australia in the northwest direction around the Early Cretaceous and

subsequent continental breakup took place in Hauterivian (132 Ma) (Choudhuri et al., 2010). The new

ocean between India and Antarctica propagated westwards from the Kerguelen hotspot after about 130

Ma (Reeves et al., 2004; Figure 2.4b). The continental fragment of Elan Bank also separated from India

and migrated to Antarctic plate as a consequence of ridge jump event at around 120 Ma (Gaina et al.,

2007).

The first outbreak of Kerguelen hotspot took place around 130-125 Ma, closely around the breakup of

India and Antarctica (Storey et al., 1992; Coffin et al., 2002; Duncan, 2002). However, the major phases

of eruption occurred during an interval of 120-95 Ma (Coffin et al., 2002). Several volcanic ridges

associated with hotspot tracks are identified. Among them, 90°E Ridge is most notable. It is a linear,

aseismic, age-progressive seamount chain in the Indian Ocean (Müller et al., 1993; Ramana et al., 2000;

Verzhbitsky, 2003) (Figure 2.1). Another important seamount chain is 85° E Ridge, which is also a linear,

aseismic volcanic ridge, now buried under sediments (Curray and Munasinghe, 1991; Ramana et al.,

1997; Bastia et al., 2010) (Figure 2.5).


19

Figure 2.4: The four phases of Gondwana breakup and drift history (modified after Reeves and DeWit, 2004). (a) Breakup between Eastern Gondwana and Western Gondwana shows largely controlled by the dextral strike-slip movement along the Davie and the Lebombo-Explora fracture zones. (b) The breakup between India and Antarctica took place around 130 Ma with evidence early ocean opening between two continents. (c) The breakup between India and Madagascar, which was largely controlled by the dextral transtensional movement. (d) The separation of India and Madagascar was followed by rapid northward drift of India. The drift was culminated at around 55 Ma when Indian plate finally collided with Eurasian plate, resulted in Alpine-Himalayan orogeny. AFC=Africa, SAM= South America, MAD=Madagascar, IND=India, AUS=Australia, ANT=Antarctica, DFZ=Davie Fracture Zone, LEFZ= Lebombo-Explora Fracture Zone


20

The breakup between India and Madagascar took place at around 93-90 Ma (Storey et al., 1995; Reeves

et al., 2004; Figure 2.4c). The Madagascar-India break-up was characterized by dextral transtension

(Reeves and de Wit, 2000; Reeves, 2003; Reeves, 2009). After the Madagascar breakup, a rapid

northward drift of India was initiated. This Indian movement was punctuated at about 65 Ma by a ridge

reorganization, which left a fossil ocean basin between Madagascar and Mascarene fragments (Reeves,

2009). This period is also marked by the major outbreak of Deccan volcanism in central and southern

India, while India was moving over Réunion hotspot (Reeves and de Wit, 2000; Singh, 2002; Seth et.al,

2004). This event is also synchronous with breakup between India and Seychelles microcontinent (Collier

et al., 2008). The drift of Indian plate was culminated at around 55 Ma when Indian plate finally collided

with Eurasian plate, which resulted in Alpine-Himalayan orogeny (Aitchison, 2007) (Figure 2.4d).

2.3 East Indian passive margin

The East India passive margin extends from shelf to an ultra-deep basin in Bay of Bengal between the

latitudes of 10°N to 21°N and longitudes of 80°E to 90°E (Figure 2.5). The East India passive margin

fringing Bay of Bengal towards east encompasses almost 2880 km length along coastline from the

southern tip of Indian peninsula to Ganga-Brahmaputra delta towards north. The island of Sri-Lanka is

also considered a part of extended East Indian passive margin with similar evolution history. Significant

part of the passive margin known as Greater India is already subducted beneath Eurasian plate during

Alpine-Himalayan orogeny. The passive margin has been evolved from Mid-Jurassic to recent following

the Gondwana breakup.

2.3.1 Physiographic Setting

The East India passive margin is characterized by narrow shelf (50-80 km in width) and a gentle slope

area (Figure 2.5). The Coastal region is characterized by well-defined deltas, beaches, lagoonal lakes with

backwater swamps, which are strongly affected by long shore drift currents. The topographic slope of

peninsular India is towards east. Therefore, a large number of rivers with large drainage basin flow

towards the Bay of Bengal. Major deltas are formed by important peninsular rivers like Cauvery, Krishna,

Godavari, Mahanadi and most significant Ganges-Brahmaputra, which drains huge sedimentary load


21

into Bay of Bengal. Ganges-Brahamaputra deltaic plain is characterized by abundant channels, small

lakes and tidal flats.

The passive margin contains five major sedimentary basins in the onshore areas. Based on detailed

geophysical data, these basins extend offshore. The underlying basement is typically dominated by

alternating set of NE–SW trending ridges and depressions (Sastri et al., 1981; Fuloria, et al., 1992;

Prabhakar and Zutshi, 1993; Bastia, 2006).

Figure 2.5: Physiographic setting of East India passive margin. The map shows the extent of study area, onshore topography, drainage system and offshore bathymetry. Both bathymetry and topography are overlaid on shaded relief map.


22

2.3.2 Geological and Tectonic Setting

The onshore geology and tectonic setting of East India passive margin varies considerably along its

length. The tectonic framework of peninsular India in the study area can be divided into tectono-

stratigraphic (Valdiya, 2010) and lithological (Ramakrishanan and Vaidyanadhan 2010) units (Figure 2.6);

1. Precambrian Indian shield region including Archean cratons, Proterozoic mobile belts and

Proterozoic sedimentary basins;

2. The Permian-Triassic intra-cratonic rift basins;

3. Jurassic-Early Cretaceous rift basins;

4. Late Cretaceous flood basalt provinces including Rajmahal and Shyllet traps and

5. Late-Cretaceous Tertiary passive margin basins

2.3.1.1 Precambrian Shield Regions

The onshore geology of East India is dominated by the Proterozoic mobile belts juxtaposed to Achaean

cratons (e.g., Ray, 1963; Radhakrishna and Naqvi, 1986; Naqvi and Rogers, 1987; Geological Survey of

India, 1993) (Figure 2.7). The peninsular India is a shield area composed of four composite, consolidated

rigid cratonic blocks evolved mostly during Archean period. These include Dharwar craton in southern

India, Bastar craton in central India, Singhbhum craton in eastern India and Bundelkhand craton in

north-western India (Valdiya, 2010) (Figure 2.8). The cratonic ages vary between 3.0 to 3.5 Ga

(Ramakrishanan and Vaidyanadhan, 2010). The cratonic stabilization took place around 3.0 Ga with

major granitic emplacements as a result of crustal scale migmatization (Naqvi and Rogers, 1987; Valdiya,

2010).

Dharwar craton is a classical greenstone-granite terrain composed of gneisses and granites and volcano-

sedimentary succession formed in shallow water platforms to deep water basins (Ramakrishanan and

Vaidyanadhan, 2010). The rock succession records regional deformation due to widespread Dharwar

and Pan-African orogenies during Neoproterozoic, with major intra-cratonic shear zones. The regional

metamorphism varies from green schist to granulite grade with significant Banded Iron Formation (BIF)

formations (Naqvi and Rogers, 1987; Ramakrishanan and Vaidyanadhan, 2010; Valdiya, 2010). A large

part of the Dharwar craton is composed of Peninsular Gneissic Complex (PGC). Peninsular Gneissic


23

Complex refers to large complex assemblages of metamorphosed granitic rocks possibly formed by

migmatization during cratonic stabilization period. It is to be noted that PGC is not unique to Dharwar

craton but it also forms a significant portion of other above mentioned cratons.

Figure 2.6: The onshore geology of East India passive margin showing the major tectonostratigraphic and lithological units (modified after the Geological Survey of India, 1993).


24

The southern part of Dharwar craton comprises highly deformed mobile belt commonly known as the

Southern Granulite Terrain (SGT) (Naqvi and Rogers, 1987; Ajaykumar et al., 2006) (Figure 2.9). This is an

intensely deformed rock suite composed of high-grade metamorphic rocks of granulite facies with

number of internal shear zones (Ramakrishanan and Vaidyanadhan, 2010). The north-eastern part of the

Dharwar craton is overlain by Proterozoic sedimentary successions and Mesoproterozoic Eastern Ghat

Mobile Belt (EGMB).

Bastar craton has a similar cratonic composition, including gneissic complexes and volcano-sedimentary

supra-crustal succession. The eastern and south eastern part of Bastar craton is bounded by Proterozoic

Eastern Ghat Mobile belt.

The Singhbhum craton is separated from the Bastar craton by the Palaeozoic Mahanadi rift zone. The

Singhbhum craton is characterized by the oldest rocks of Indian subcontinent. It comprises of

widespread Archean tonalitic to granodioritic granites and gneisses and thick volcano-sedimentary

supra-crustal succession with abundant Banded Iron Formation (BIF) and large ultrabasic intrusives. All

rocks have undergone regional deformation as well as amphibolite facies metamorphism (Ramakrishnan

and Vaidyanadhan, 2010; Valdiya, 2010). The Singhbhum craton is demarcated by two prominent mobile

belts, Satpura Mobile belt in the east, north and north-west and Eastern Ghat Mobile Belt in the south.

The Eastern Ghat Mobile Belt (EGMB) encompasses a significant portion of basement rocks of East India

passive margin. It was formed as the result of widespread Mesoproterozoic orogeny that was controlled

by oblique collision between continental plates (Valdiya, 2010). The bulk of EGMB is composed of

granulite facies metamorphic rock of charnokite-khondalite succession representing exhumed lower

crust. It has its counterpart in the Napier Complex of Enderby Land, East Antarctica (Bhattacharya,

2001). Both terrains share distinct petrologic, mineralogical similarities such as, for example, the

occurrence of saphirine- cordierite- spinel-quartz assemblages in granulites (Biswal and Sinha, 2004),

which are a product of high grade metamorphism (Harley and Hensen, 1990; Harley and Motoyoshi,

2000).


25

Figure 2.7: Precambrian geology of the onshore East India showing the major tectono-stratigraphic and lithological

units (modified after the Geological Survey of India, 1993)


26

Figure 2.8: Distribution of stable composite, consolidated rigid Archean cratonic blocks (modified after Valdiya, 2009). These include Dharwar craton in southern India, Bastar craton in central India, Singhbhum craton in Eastern India and Bundelkhand craton in north-western India. The cratonic ages vary between 3.0 and 3.5 Ga.


27

2.9: Simplified geological map of the Southern Granulite Terrain (Ajaykumar, et al, 2006).

Mobile belts of both margins indicate their thrust belt character by:

1. the presence of nappes/thrust sheets in various parts of the Eastern Ghat Mobile Belt (e.g.,

Mahalik, 1994; David et al., 1998; Neogi and Das, 1998; Biswal and Sinha, 2004);

2. the occurrence of stacked terrains of various age of metamorphism in the Eastern Ghat Mobile

Belt and Southern Granulitic terrain of India, and Napier and Rayner complexes of Antarctica

(e.g., Black et al., 1987; Biswal and Sinha, 2004);

3. the coincidence of deformation histories (e.g., Clarke, 1988; Yoshida, 1995; Kovach et al., 1997);


28

4. the occurrence of deepest rocks exhumed in the orogenic hinterland (Biswal and Sinha, 2004);

and

5. the nice correlation of the NW-verging thrustbelt detachment (Biswal and Sinha, 2004).

The contact between the Eastern Ghat Mobile Belt and its cratonic foreland on Indian side is

represented by the Terrane Boundary Shear Zone (Figure 2.8; Ray, 1963; Biswal and Sinha, 2004). It is

about 2 km thick ductile shear zone, having a thrust-dominated character along the NE-SW trending

segments, sinistral transpressional character along the N-S trending segments and dextral strike-slip

dominated character along the WNW-ESE segments. The shear zone is formed by mylonites with

predominating quartz-feldspatic composition due to the cratonic footwall influence. It contains

amphibolites on the hanging wall side due to retrograde metamorphism of the orogenic rocks.

A number of partly deformed and almost unmetamorphosed Proterozoic intra-cratonic to peri-cratonic

basins forms a significant part of Precambrian Indian shield. As a result of Eastern Ghat orogenic

movements, a number of shallow subsiding basins were formed in the active parts in front of the older

orogenic belts (Valdiya, 2010). These are mostly sub-horizontal Mesoproterozoic to Neoproterozoic

sedimentary sequences deposited over crystalline Archean gneissic basement. Their contact is

represented by a sharp angular unconformity or non-conformity in places known as Eparchean

unconformity. These basins predominantly contain quartzite-shale-carbonate suites, which were

deposited episodically and mostly in fluvial to shallow marine environments in platform setting

(Ramakrishnan and Vaidyanadhan, 2010).

The sutures between stable Indian cratons remained the zone of weakness, which were affected by

sagging and also reactivated during rifting and breakup of Gondwanaland (Figure 2.8). The Son-Narmada

rift valley is a plaeosuture, which have been reactivated many times, demarcates the boundary between

Bundelkhand craton against Bastar, Singhbhum and Dharwar cratons. The Palaeozoic Pranhita-Godavari

rift zone divides the Dharwar and Bastar cratons. Similarly Bastar and Singhbhum cratons are separated

by the Paleozoic Mahanadi rift zone (Figure 2.8).

The basement of East Indian passive margin is represented by the Southern Granulite Terrain (SGT) (e.g.,

Drury et al., 1984; Naqvi and Rogers, 1987), peninsular gneisses (PGC) (Naqvi and Rogers, 1987; Valdiya,


29

2010) and the Eastern Ghat Mobile Belt (e.g., Krishnan, 1982; Radhakrishna and Naqvi, 1986; Naqvi and

Rogers, 1987; Bhattacharya, 1996) (Figure 2.7). The orogenic units are represented by:

I. The high-grade metamorphic rocks (charnokites) belonging to Southern Granulite

terrain (SGT) in the region between the individual Cauvery basins (Ramakrishnan et al.,

1998);

II. The extensively remobilized basement with granitic gneisses belonging to Peninsular

Gneissic Complex (PGC) of Dharwar craton in the region adjacent to the Cauvery basin

and Krishna-Godavari basins (Naqvi and Rogers, 1987; Ramakrishnan, 2003);

III. The Eastern Ghat mobile belt with high grade metamorphic rocks in the region adjacent

to the western portion of the Krishna-Godavari basin (Naqvi and Rogers, 1987;

Gopalakrishnan, 1998; Ramakrishnan et al., 1998);

IV. The extensively remobilized basement with granitic gneisses belonging to gneissic

Complex of Bastar craton in the region adjacent to the eastern portion of the Krishna-

Godavari basin (Naqvi and Rogers, 1987; Ramakrishnan et al., 1998); and

V. The Singhbhum batholitic granites, granodiorites and older metamorphic group rocks

(OMG) in the region adjacent to the Mahanadi basin (Geological Survey of India, 1993;

Gopalakrishnan, 1998).

2.3.1.2 Permian-Triassic Intra-Cratonic Rift

There was a profound hiatus of about 250 Ma observed in peninsular Indian shield region after

formation of Proterozoic sedimentary basins. During Early Permian time, the entire Gondwanaland

including India experienced a regional extension. The shear zones, faults and Precambrian rift valleys

were reactivated under this new tectonic regime and a number of elongated rifted half grabens and

grabens were formed as a result of cratonic block faulting and associated subsidence. However, this

intra-cratonic rift cycle never reached up to continental breakup. This tectonic regime was active

through Early Jurassic until another tectonic regime, which caused the breakup of Gondwanaland, took

over.

The Permo-Triassic sequences of peninsular India contain mainly a thick pile of clastic sediments

preserved in grabens and half-grabens (Figure 2.10). The essential features of Permo-Triassic intra-

cratonic rift basins can be characterized as follows (Lakshminarayana, 2002; Chakaraborty et al, 2003):


30

1. Primary deposition within narrow elongated intra-cratonic rift valleys;

2. A glacial boulder bed on top of crystalline basement, followed by monotonous thick sequence of

sand, shale and coal;

3. The glacial sequence followed by fluvio-lacustrine depositional system;

4. Presence of Glossopteris or Ptilophyllum flora

5. Almost consistent west to north-west directed paleoslope

The East India passive margin contains two prominent NW-SE trending rift zones; the Pranhita-Godavari

rift zone and Mahanadi rift zone, both Gondwana basins roughly perpendicular to the coastline,

containing Early Permian-Early-Jurassic sedimentary fill.

The Pranhita-Godavari (PG) basin consists of a series of NNW-SSE grabens and half-grabens. The

Bouguer anomaly contours also conform to the shape of the basin and are closely spaced at the eastern

margin. Magnetic anomaly map of the offshore Krishna-Godavari basin indicates a deep structural grain

characterized by NW-SE trend, parallel to the trend of the Permo-Triassic onshore rift (Erram et al.,

2005). Pranhita–Godavari basin shows the characteristics of a typical rift basin with a master basin

bounding fault to the east. It underwent nearly orthogonal extensional, controlling displacement

towards WSW (Biswas, 2003). The basin is slightly asymmetric, showing the maximum sediment

thickness near its western margin.

The rift faults are steeply dipping. The easterly stratae dip indicates that downward displacement of the

hanging wall was preferentially located along rift bounding faults leading to half-graben geometry and

associated tectonically-controlled sedimentation. The tectono-sequences of PG basin are primarily

controlled by three prominent fault systems (Chakraborty et al., 2003)

i. NW-SE en-echelon fault system parallel to Dharwar trend that defined basin-bounding fault

system

ii. NNW-SSE trending faults controlling the half-graben structures and sedimentation pattern

iii. NW-SE trending faults parallel to Eastern-Ghat trend

Phased uplift of the Eastern-Ghat mobile belt also contributed significantly to the intercontinental

drainage system. Sedimentological analysis of the Permian and Triassic coal-shale-sand sequences of the

Pranhita-Godavari rift zone indicates the sediment transport from south-east, implying that the

provenance was located somewhere in Antarctica (Laksminarayana, 2002).


31

Figure 2.10: Distribution of Permo-Triassic rocks within intra-cratonic rift systems (modified after Chakraborty et al, 2003). Major faults and lineaments associated with different Permo-Triassic rift basins are given as follows. SNsF: Son-Narmada south fault; SNnF: Son-Narmada north fault; GCF: Gavilgarh fault; GVnF: Godavari valley north fault; GVsF: Godavari valley south fault; GVcF: Godavari valley cross fault; TANs: Tan Shear; BCF: Bahmani Chilpa fault; MF: Mahendragarh fault; TnF: Tapti north fault; TF: Tatapani fault; SBF: Sainthia Bahamani fault; RMF: Rajmahal fault.

The Mahanadi rift zone is an elongated NW–SE oriented rift zone, bounded on the southwest by major

easterly-dipping fault and a master bounding fault to the west. The Bouguer gravity anomaly contours

are closely spaced near the western margin, indicating the greater thickness of the basin fill located

close to the boundary fault. The basinfill strata generally strike along the length of the basin defining a

broad synclinal structure (Chakraborty et al., 2003).


32

Fission track study of zircon and apatite grains from the Mahanadi rift zone indicates a two-stage

opening history of its rift units (Lisker and Fachmann, 2001). The timing of the older event is further

constrained by the igneous intrusions associated with Indo-Australian rift zone, having an age of 300 Ma

(Veevers and Tewari, 1995; Veevers 2009). Both rift events can be correlated with Permian and Triassic

depositional systems, which are also recorded on the conjugate margin, in the Lambert graben of the

Eastern Antarctica, which has been mapped as the continuation of the Mahanadi rift zone (Fedorov et

al., 1982; Hofmann, 1996; Lisker and Fachman, 2001; Läufer and Phillips, 2007).

2.3.1.3 Jurassic-Lower Cretaceous passive margin

The studied portion of the East Indian margin can be divided into the six tectonic segments (Sinha et al.,

2010; Nemčok et al., 2013b; Figure 2.11):

I. The NE-SW trending Cauvery rift zone. Its structural grain is dominated by the Jurassic-Lower

Cretaceous pattern of NE-SW normal faults controlling the system of horsts and grabens (Rao,

1993).

II. NNW-SSE trending Coromandal transfer zone. This zone is characterized by significant dextral

strike-slip displacement, encompassing the Pennar- Palar basin. The accommodation zone forms

a hard linkage between Cauvery and Krishna-Godavari rift zones (Sinha et al., 2010; Nemčok et

al., 2013b).

III. NE-SW to ENE-WSW trending Krishna-Godavari rift zone, located in front of the pre-existing NW-

SE trending Permo-Triassic Pranhita-Godavari rift zone (Rao, 2001). The Krishna-Godavari Basin

is characterized by a set of NE-SW to ENE-WSW trending horsts, grabens and half-grabens (Rao,

2001; Choudhuri et al., 2010).

IV. NNW-SSE trending North Vizag transfer zone. This accommodation zone links Krishna-Godavari

and Mahanadi rift zones. Although there are certain strike slip character can be seen, this

margin segment forms dominantly represents a soft linkage between Godavari and Mahanadi

rift zone.

V. NE-SW trending Mahanadi rift zone. The Mahanadi rift zone lies to the SE of the outcropping

NW-SE trending Gondwana Permo-Triassic Mahanadi rift zone. Its structural grain contains the

NE-SW striking Jurassic-Lower Cretaceous fault pattern that controls a system of depressions


33

and highs. Significant volcanic activity is observed in both Permo-Triassic and Jurassic-Lower

Cretaceous rift zones (Fuloria et. al., 1992).

VI. NNW-SSE trending Konark transfer zone having a small dextral strike-slip displacement possibly

forming a hard linkage.

During middle Jurassic, another extensional tectonic regime became active with a rifting vector

orientation of NW-SE (Shah et al., 2007). This extensional regime caused the ultimate breakup of

Gondwanaland into several continental landmasses. The reactivation of older crustal fabrics, faults and

shear zones influenced the orientation of new NE-SW trending rift fault system. The East India passive

margin evolved as a result of rifting of eastern Gondwanaland when India separated from Antarctica

during the Early Cretaceous. The onshore region, in close proximity to the shoreline contains syn-rift and

post-rift Upper Jurassic-Lower Cretaceous sediments in the passive margin basins, which are formed due

to rifting of continental margin under the new tectonic regime.

The continental break-up that followed Jurassic-Lower Cretaceous rifting resulted in the oceanward tilt

of the East India passive margin, controlling the SE-ward flowing drainage system in Cauvery, Pennar-

Palar, Krishna-Godavari and Mahanadi passive margin basins (Rao, 1993). Early Cretaceous marine

transgression observed in these basins is another event associated with this SE-ward tilt.

The Cauvery Basin is located in the south-eastern part of the Indian peninsula. The initial rifting caused

the formation of NE-SW horst-graben features (Narasimha Chari, et al, 1995) . The structural grain of the

Cauvery basin is dominated by the Jurassic-Early Cretaceous pattern of NE-SW normal faults controlling

the system of horsts and grabens (Rao, 1993; Figure 2.12). As Figure 2.12 illustrates, the basin can be

roughly divided into five depressions located around four structural highs, which are relatively large and

have steep slopes. In the exploratory wells in onland and offshore, the sediments range in age from Late

Jurassic to Recent. The peneplanation of horsts was completed by the end of Santonian and the present

shelf-slope system appears to have existed since the Late Santonian (Prabhakar and Zutshi, 1993).


34

Figure 2.11: Map of the major tectonic segments of the study area and structural features of the adjacent onshore (Nemčok et al, 2013b). There are six major segments: (a) The NE-SW trending Cauvery rift zone; (b) The NNW-SSE trending dextral Coromandal transfer zone; (c) The NE-SW to ENE-WSW trending rift units of the Krishna-Godavari rift zone; (d) The NNW-SSE trending North Vizag transfer zone between the Krishna-Godavari and Mahanadi rift zones; (e) The NE-SW trending Mahanadi rift zone; (f) The NNW-SSE trending dextral Konark transfer zone.


35

The Krishna-Godavari basin is located in the central part of the East India passive continental margin.

The basin runs along the coast, abutting against the outcropping NW-SE trending Pranhita-Godavari rift

zone (Rao, 1993; Choudhuri et al, 2010; Figure 2.13). The onshore portion of the Krishna-Godavari basin

contains the three main NE-SW trending depressions. They include Krishna, West Godavari and East

Godavari depressions and two main NE-SW trending structural highs; Bapatla and Tanaku-Kaza horsts.

Their geometries coincide with NE-SW linear and planar anisotropies of the Eastern Ghat Mobile Belt.

Figure 2.12: Tectonic element map of the Cauvery Basin. It shows the onshore (modified after Narasimha Chari et al, 1995) as well as offshore rift fault trends. The structural grain of the Cauvery Basin is dominated by the Jurassic-Early Cretaceous pattern of NE-SW normal faults controlling the system of horsts and grabens, whereas the strike-slip segment faults shows mostly N-S trend along northern part of Cauvery Basin as well as the western margin of Sri Lanka.


36

Depressions and structural highs are prominent enough to have an impact on gravity and magmatic

imaging (Murty and Ramakrishna, 1980; Rao, 1980). Similar structural grain exists in the offshore, while

the onshore portion of the former syn-rift architecture is characterized by highs frequently eroded down

to the basement level (Choudhuri et al, 2010).

The Mahanadi basin is situated in the northern most region of the East India passive margin. Analogically

to the Krishna-Godavari basin, this basin is also located to the SE of the outcropping NW-SE trending

Permo-Triassic Mahanadi rift zone. Its structural grain contains the NE-SW striking Jurassic-Lower

Cretaceous fault pattern that controls a system of depressions and highs (Fuloria et al, 1992, Figure

2.14). The depressions and highs are significant enough to be identified in gravity and magnetic images

(Rao, 1980).

Majority of the passive margin basins recorded magmatic activity in some extent (Figure 2.15). The

major magmatic pluses occurred in the Early Cretaceous (Albian-Santonian), Late Cretaceous

(Maastrichtian) to Eocene time interval. These volcanic events have been interpreted in conjunction

with the Kerguelen Hotspot. The oldest magmatic activity resulted in magmatic products in the

Rajmahal lava flows while the youngest event can be dated from outcropping Razole (Paleocene-

Eocene) basalts of the Krishna-Godavari basin. Rajmahal traps, lamprophyre, dolerite dykes and sills of

the Damodar, Deogarh and subsurface Bengal basins and the offshore Mahanadi basin range in age from

120 to 110 Ma, within Early Cretaceous (Veevers and Tewari, 1995).

Unlike other passive margin basins, Mahanadi basin suffered maximum magmatic effect, possibly due to

outbreak of Kerguelen plume magmatism immediately after continental breakup. Mafic dyke intrusions,

hydrothermal overprint of faults and thrust zones at about 120 Ma in this basin are considered

synchronous with both Rajmahal magmatism in Northern India and Eastern Antarctica (Lisker and

Fachmann, 2001). 40Ar/39Ar dating of mafic dykes from Rajmahal volcanics yields an age of about 117 Ma

(Baksi, 1995; Tiwari and Tripathi, 1995).

Late Jurassic – Early Cretaceous is also the time of the igneous activity reported from Meghalaya-

Nagaland in NE India (Sarkar et al., 1996), carbonatite magmatism in the Shillong Plateau (Kumar et al.,

1996) and basaltic effusions in the Bengal basin (Sengupta, 1966).


37

Figure 2.13: Onshore (modified after Choudhuri et al, 2010) and offshore basement trends in the Krishna-Godavari Basin showing the older Pranhita-Godavari northwest–southeast trend overlain by major basement trends of the younger rift episodes. The rift structure has been delineated based on gravity-magnetic and seismic data.


38

2.3.1.4 Cretaceous flood basalt provinces

As mentioned in the earlier section, at the beginning and the end of Cretaceous period, India suffered

voluminous outpouring of lavas and volcanic explosions. The lower Cretaceous Rajmahal volcanic

province encompasses nearly 20000 km2 area of Eastern India (Figure 2.15). The Sylhat trap in

Meghalaya is another example of volcanic provinces. The type area of the Rajmahal province shows

nearly 244 m thick lava pile with individual flows varying in thickness from less than 1 m to more than 70

m (Baksi, 1995).

Figure 2.14: Tectonic map of Mahanadi basins showing several horst and graben structures due to orthogonal rifting between India and Antarctica (Fuloria et al., 1992).


39

Figure 2.15: Mafic magmatic activity in the Mahanadi Basin took place possibly due to outbreak of Kerguelen plume magmatism immediately after continental breakup (The background is formed 10 km high-pass Bouguer gravity anomaly map). Major magmatic pluses occurred during the Middle to Late Cretaceous (Albian-Santonian). The large magmatic emplacement in the offshore is represented by the 85°E Ridge, which can be roughly correlated with analogues onshore via several mafic rocks (mostly basalt), encountered in the wells and outcrops at Rajmahal trap.

The Rajmahal and Sylhat lavas are predominantly formed tholeiitic basalt, quartz-tholeiite and alkali

basalt with minor amounts of dacite, agglomerates and tuffs (Baksi, 1995; Kent et al., 2002; Ghatak and

Basu, 2011). Lava flows are associated with intertrapean sequences of siltstone, claystone and shale

(Kent et al., 2002). Petrochemical and isotope age data from Rajmahal volcanics suggest a link with

Kerguelen Plateau in Indian Ocean (Baksi, 1995; Kent et al., 2002; Ghatak and Basu, 2011). It can be also


40

inferred that there are very little crustal contamination of the mantle derived magma representing less

than 20% (Ghatak and Basu, 2011).These lavas are results of partial melting of a relatively primitive

Kerguelen plume source with components of the Indian Ocean E-MORB and contaminants derived from

the lower continental granulites. Sylhet Traps are also considered as the remnants of once larger

Rajmahal flood basalt province.

The most remarkable phase of volcanism occurred on Indian subcontinent during Masstrichtian-

Paleocene (KT boundary), while Indian plate was moving over the Reunion hotspot. This is commonly

known as phase resulting in Deccan Volcanic Province (DVP) (Valdiya, 2010) (Figure 2.16). The bulk of

Deccan volcanics can be typically characterized by continental flood basalt type, mostly of tholeitic

basalt with minor association of alkali basalts like picrites and acidic basalt like rhyolite (Sen, 2001).

There are several fossiliferous intertrapean sedimentary sequences associated with such massive lava

flows. The lava flows forming the Deccan Province are not flat lying; the western flank is steeper than

the eastern flank. The equivalent of Deccan trap is also present in in onshore Krishna-Godavari basin in

East India is known as Razole volcanics of Paleocene age. Its outcrops are well exposed in onshore

Krishna-Godavari basin (Rao, 2001).

2.3.1.5 Late Cretaceous to present day passive margin

Between the Late Cretaceous and present, the Indian peninsula has moved to the north and underwent

about 20°counterclockwise rotation in the same time (Srivastava and Chowhan, 1987; Gordon et al.,

1990; Kumar et al., 2007). This changed the initially E-W striking eastern coastline to the currently NE-

SW striking coastline (Rao, 2001).

The rate of India’s northward movement towards Asia was around 6.6 cmyr-1 during the 120-73 Ma time

interval (Aitchison et al., 2007). It increased to nearly 21.1 cmyr-1 during the 73-57 Ma time interval. At

57 ± 3 Ma, in Late Paleocene, there was a significant slowdown to nearly 9.5 cmyr-1 until 20-30 Ma

(Acton, 1999). This slowdown is most recently interpreted as associated with the Late Paleocene

collision of an intra-oceanic subduction system with the northern Indian margin (Aitchison et al., 2007;

Figure 2.17). The 20-30 Ma time interval is characterized by further major slowdown, interpreted as

associated with India-Asia intercontinental collision (Aitchison et al., 2007).


41

The response to the India-Asia collision was immediate and involved cessation of calc-alkaline arc

magmatism along the southern margin of Asia, uplift of the Tibetan Plateau, collisional orogenesis,

molasse deposition and readjustment of plate boundaries throughout eastern Asia (Aitchison et al.,

2007). During the same time, the sedimentary load of Krishna and Godavari rivers increased, initiating a

major development of the Krishna-Godavari deltas.

Figure 2.16: The distribution map of Deccan Volcanic Province and equivalent rocks in Peninsular India (A Misra, pers comm.)


42

Figure 2.17: Plot of significant geological phenomena recorded in rocks of southern Tibet and elsewhere in the region (Aitchison, 2007 and references therein). These are potentially related to collision of an island arc system with India at around 55 Ma (event 1), followed by India-Asia collision at around 35 Ma (event 2). Note the hiatus in collision-related events between these two time periods.


43

2.4 Kerguelen Plateau and Elan Bank microcontinent

The Kerguelen Plateau is one of the largest Large Igneous Provinces (LIP) and submarine plateau in the

world. It is located in southern Indian Ocean between 45°S to 65°S and 60° to 90°E.

2.4.1. Physiography and plate boundary setting

Kerguelen Plateau in present day is located in the deep waters (1000 to 4000 meters) in Southern Indian

Ocean, 3,000 km south west of Australia and 2500 km north of Antarctica. It extends for more than 2200

km in a northwest-southeast direction. The plateau was produced by the Kerguelen hotspot, starting

with or following the breakup of Gondwana about 130 million years ago. The Kerguelen Plateau

currently is part of the Antarctic Plate (Figure 2.1; Figure 2.18). It is bounded to the northeast by the

oceanic Australian-Antarctic Basin, to the southwest by the oceanic Enderby Basin and to the northwest

by the oceanic Crozet Basin. The Kerguelen Plateau is separated from the East Antarctic continental

margin by the Princess Elizabeth Trough (Bénard et al., 2010). A large sedimentary basin, the Raggatt

basin, lies on the top of the southern Kerguelen Plateau in water depths of 1500 m (Operto and Charvis,

1995). There is a small portion of the plateau is rises above sea level present day, forming the Kerguelen

Islands, Heard Island and McDonald Islands. Intermittent volcanism continues on Heard and McDonald

islands.

The southeast Indian Ridge separates the Australian-Antarctic Basin from the Indian Ocean. The ocean

floor in Australian-Antarctic Basin to the NE of the Kerguelen Plateau contains well-defined magnetic

anomalies C19 (41 Ma) (Coffin et al., 1997), which are being the oldest, caused the separation between

Kerguelen Plateau and Broken Ridge (Figure 2.19). No magnetic anomalies have been identified in the

area surrounding Kerguelen Plateau in the south towards Antarctic continental margin. The early

evolution of the south Indian Ocean is poorly constrained, as the sea-floor-spreading magnetic

anomalies have not been identified on the plateau or in much of the surrounding ocean basins. One of

the reasons may be that the most of the evolution of this ocean basin occurred during Aptian-Albian

after India-Anarctica breakup. This particular time window again encompasses the Cretaceous Quite

Zone (125-83 Ma) during which no magnetic anomalies have been produced. Similarly Ramana et al.

(2001) also described magnetic anomalies becoming progressively younger northward from the

Antarctic margin towards the central part of the basin (M10 to M0, 136.5–124 Ma).


44

Figure 2.18: The present day plate settings in the southern Indian Ocean showing Kerguelen Plateau, Elan Bank, Enderby Basin and Prince Elizabeth Trough (Bénard et al., 2010 and references there in)


45

Figure 2.19: Interpretation of gravimetric and magnetic maps associated with bathymetry (Bénard et al., 2010 and references there in). The placement of oceanic fracture zones indicated in the map. 1: Distribution of magnetic anomalies C0 to C11; 2: Distribution of magnetic anomalies C0 to C19; 3: Distribution of magnetic anomalies C0 to C21, 4: Distribution of magnetic anomalies C23 to C33; 5: Gravimetric anomalies interpreted as inferred seamount; 6: Inferred rectilinear feature.


46

2.4.2 Kerguelen volcanic province

Most of the Kerguelen Plateau was formed by emplacement of ocean-island type basalts (OIB) produced

by the Kerguelen Hotspot in the period from about 117 Ma to present. It has been recorded that

hotspot produced approximately 25 million km3 of mafic crust since 130 Ma. However, the rates of

magma production are highly variable through geological time depending upon different geodynamic

condition (Coffin et al., 2002) (Figure 2.20). The 40Ar/39Ar radiometric dating suggests that the Kerguelen

hotspot magmatism has a strong affinity with Broken ridge, Bunbury basalt of Australia and Rajmahal

traps of India and it is obtained from a relatively less depleted plume source (Ghatak and Basu, 2011).

Figure 2.20: Estimated Kerguelen magma output since 130 Ma (Coffin et al., 2002). Analytical uncertainties are variable but values are generally < 5Ma). It clearly indicates the maximum magma flux due to Kerguelen plume at various places. Note, it is quite variable in space and time. The maximum flux (0.1 km3/yr) in Elan Bank around 112 Ma.


47

2.4.3 Crustal structure of Kerguelen Plateau

The crustal architecture of Kerguelen Plateau is geologically very complex and difficult to interpret.

Several ODP and DSDP wells have been drilled at Kerguelen Plateau (Figure 2.21). The crustal structure

and nature of the crust of Kerguelen plateau has been studied mostly using potential field and refraction

seismic studies, which suggest a thicker crust in the range of 18-25 km (Operto and Charvis, 1995;

Charvis and Operto, 1999; Galdczenko and Coffin, 2001; Borissova et al, 2002) (Figure 1.3; 1.4; 1.5;

Figure 2.22). The Basement rocks recovered from drilled wells are mostly represented by magmatic

rocks. However, refraction and geochemical data also indicate unusually thick crust and possible

continental contamination in some parts of the plateau. ODP drilling on the Elan Bank recovered

continental rocks. This confirms that Kerguelen Plateau may be partly underlain by continental crust.

A typical refraction profile shows that the uppermost part of the Kerguelen Plateau is covered by an

upper sedimentary layer with velocities of 1.7 to 2.0 kms-1 representing unconsolidated Cenozoic

sediments and a lower sedimentary layer with velocities 2.6 to 3.1 kms-1 representing consolidated

Mesozoic sediments. The Kerguelen-Heard Basin within CKP is characterized by a deep crustal root,

which suggests that the crust has been probably emplaced in the Late Cretaceous (Charvis et al., 1995).

The wide angle seismic data show that the crustal thickness of the Northern Kerguelen Plateau (NKP) is

in the order of 21–23 km. The lower crust velocity ranging from 6.8 to 7.3 kms-1 is more consistent with

an oceanic origin and typically represents volcanically thickened oceanic crust (Charvis and Operto,

1999) (Figure 2.22a). The Kerguelen archipelago within Northern Kerguelen Plateau (NKP) indicates

typical oceanic crust, with volcanic roots below (Operto and Charvis, 1996)

Igneous basement of the Central Kerguelen Plateau (CKP) is 19 to 21 km thick, and consists of three

layers (Charvis et al., 1995). The upper layer is 1.2 to 2.3 km thick, and its velocities range from 3.8 to 4.9

kms-1. The middle layer is 2.3 to 3.3 km thick, and its velocities increase downward from 4.7 to 6.7 kms-1.

These two layers mostly correspond to volcanic accretion on top of the oceanic crust. CKP also

represents an anomalously thickened thick lower crust of about 17km thick, with velocities increasing

from 6.6 kms-1at about 8.0 km depth to 7.4 km/s at the base of the crust with no internal discontinuity

(Charvis et al., 1995). This high-velocity lower crust is usually typical for underplated oceanic crust

(Coffin et al., 2000).


48

Figure 2.21: Locations of geological sampling sites, ODP wells and sonobuoy stations used in the study by Borissova et al, (2002).


49

The thickest crust in Southern Kerguelen Plateau (SKP) is about 22km thick. The crustal velocity in the

Southern Kerguelen Plateau (SKP) is ranging in the order of s 3.9-6.0 kms-1 (Charvis et al., 1997) (Figure

2.22b). The igneous basement can be further divided into extrusive volcanics with velocities of 4.0 to 5.5

kms-1, which typically represent oceanic layer 2 and middle crust with velocities ranging 5.6 to 6.6 kms-1

with massive intrusives, which is typical of oceanic layer 3. It has been also observed that volcanics are

about 3-4 km thick on the eastern margin of the Southern Kerguelen Plateau.

Figure 2.22: a. Best fit velocity model across the northern Kerguelen Plateau from 1D inversion of travel times (Charvis and Operto, 1999). Dashed lines are iso-velocity contours with annotations (plain lines delineate main crustal layers). b. Final model across southern Kerguelen Plateau for profile 4 with iso-velocity lines at 0.1 km/s intervals (Operto and Charvis, 1995). Heavy lines indicate interfaces between main units. BB is basaltic basement; LC is lower crust and RZ is reflective zone.


50

However, this thickness is highly variable and thinned in many instances like in the Raggatt Basin located

within SKP (Operto and Charvis, 1995). Operto and Charvis, (1995) interpreted the crustal structure of

Raggatt Basin is as similar to volcanic passive margins. Similarly, the Labuan basin flanking the SKP also

shows thinned crust with a thickness of about 5km (Rotstein et al., 2001; Borissova et al., 2002). These

observations led to the suggestion that the southern Kerguelen Plateau may be underlain by extended

continental crust covered by basaltic flows (Operto and Charvis, 1995; Galdczenko and Coffin, 2001;

Bénard et al., 2010). Results of Leg 183 drilling on the Elan Bank (Coffin, et al., 2000; Ingle et al., 2002;

Borissova et al., 2003) indicate at least partly continental origin of its crust. These results are consistent

with low velocity (~6.8 kms-1) lower crust (Charvis et al., 1997).

2.4.4 Tectonic Provinces

There are six major physiographic and geological provinces identified within Kerguelen plateau (Coffin et

al., 1986; Borissova et al, 2002; Figure 2.23). These are:

I. Northern Kerguelen Province;

II. Central Kerguelen Province;

III. Southern Kerguelen Province;

IV. William’s Ridge;

V. Elan Bank; and

VI. Labuan Basin.

Northern Kerguelen Province: The Northern Kerguelen Province or NKP (45°-50°S) is a distinctive

bathymetric and gravity high, which also includes the Kerguelen Archipelago, Northern Kerguelen

Plateau and Skiff Bank (Figure 2.23).

ODP drilling (leg 183) and dredging results and detailed analysis confirm that the major part of Northern

Kerguelen Plateau basement is composed of Oligocene-Miocene basalts (Coffin et al., 2000). ODP Site

1140 suggests that the pre Oligocene sedimentary cover has been intersected in the northern margin

NKP. However, this cover is represented by chalk beds interbedded with basalts and it might be of latest

Eocene age. The micro-paleontological data from oozes indicate that the deposits are pelagic. They are

deposited at bathyal depth and coeval with extrusion of lava flows (Coffin et al, 2000).


51

Figure 2.23: Tectonic provinces of the Kerguelen Plateau and adjacent areas (Borissova et al, 2002)

This information further suggests that most of the Northern Kerguelen Plateau was formed in Oligocene-

Miocene, following the breakup of the Kerguelen Plateau and Broken Ridge in Late Eocene–Oligocene

(Borissova et al, 2002). However, analytical results from basal basalts from ODP site 1140 in Skiff bank,

which is situated to the west of NKP allowing estimate a much older age for basement being 66-68 Ma

(Duncun, 2002). Furthermore, it is interesting to note that underneath the Kerguelen Archipelago, some

peridotite xenoliths have been discovered and analyzed. The results indicate presence of ancient

Gondwana continental lower lithospheric elements (Hassler and Shimizu, 1998). However, Kieffer et al.,

(2002) indicated that there is neither chemical nor isotopic evidence for a continental crustal

component. Nevertheless, the presence of Cretaceous volcanics at Skiff Bank has serious implications

for understanding the plate tectonic history of the region.

Central Kerguelen Province: The Central Kerguelen Province, or CKP (50°-55°S), is located between the

bathymetric and gravity high of the NKP, and the bathymetric saddle on the northern side of the


52

Southern Kerguelen Plateau (Figure 2.23). The volcanically active Heard and McDonald Islands and a

major sedimentary basin known as the Kerguelen-Heard Basin are also important geological features

within the Central Kerguelen Province.

Several ODP wells have been drilled in the Central Kerguelen Province. The seismic, gravity-magnetic

and geological samples from ODP wells suggest that the CKP was formed in Early Cretaceous, sometimes

in Albian. Basement rocks recovered from Site 747 (Figure 2.24), located in the southern part of the CKP

are represented by basaltic basement erupted at about 85-88 Ma during Cenomanian (Munschy et al.,

1992; Borissova et al, 2002). ODP Site 1138 further reveals the felsic composition of basalt, which is

dated using 40Ar/39Ar dating yeilding 108 Ma during Albian (Coffin et al., 2002). ODP Site 1138 also

documents the occurance of glauconitic sandstones and claystones of Cenomanian to Turonian age,

which overlie the volcanoclastics and multiple basalt flows (Coffin, et al., 2000).

Figure 2.24: Composite logs for ODP Sites 1139, 1140, 736, 737, 1138, 1137 within Kerguelen Plateau (Borissova et al., 2002).


53

The Kerguelen-Heard Basin is a ‘sag’ basin located in the northern part of Central Kerguelen Province. It

covers an area of more than 40,000 km2 and contains in excess of 2000m of Cenozoic sediments. During

the later part of Oligocene and Miocene, the CKP gradually subsided and the plateau was covered by

calcareous chalks and oozes with some clastic component derived mostly from Kerguelen Island.

Southern Kerguelen Province: The Southern Kerguelen Province, or SKP (south of 55°S), is a bathymetric

and gravity high (Figure 1.3). This province is tectonically most complex area within the Kerguelen

Plateau. Several large NW-trending basement ridges have been identified through seismic as well as

gravity and magnetic studies (Rotstein et al., 1992; Borissova et al., 2002). These ridges are affected by

several stages of normal faulting resulting in a complex of NW and NE-SW trending rift units (Figure

2.23).

The 77° graben extends from north to south for more than 400 km. Similarly, the 59° graben trends

approximately east to west, to the south of the 77° graben. The Southern Kerguelen Plateau rift zone is

trending NW-SE, to the south of the 59° graben. The age and origin of these rift systems remains

unclear. It is interpreted that 77° graben formed around 75 Ma, during Campanian. It has been observed

that faults adjacent to the rift have been reactivated at different times with the latest movements taking

place in Miocene (Borissova et al, 2002). A large sedimentary basin, the Raggatt Basin, is located in the

northern part of the province, bounded by faults in the west and south and by an eroded basement

ridge in the east (Coffin et al., 2000; Borissova et al., 2002; Bénard et al., 2010; Figure 2.23).

ODP well results suggest that the volcanic basement of SKP was formed around Albian with the oldest

age being 117 Ma (Duncun, 2002). The petrographic characteristics of volcanics are not typical of

submarine volcanism, compared to the north and central provinces. On the contrary, they show more

affinity to the sub-aerial volcanism with significant amount of erosion. SKP basalts can be best described

as ocean island basalts due to their isotopic characteristics (Mahoney et al., 1995). However, they also

show a geochemical evidence of contamination by continental lithosphere. This evidence possibly

indicates the presence of continental crust beneath the Southern Kerguelen Province. It is also possible

that the thickened crust in SKP may be attributed to continental fragments along with large volume of

sub-aerial Late Cretaceous basalts (Coffin et al., 2002)(Figure 1.4). This may further increase the crustal

heterogeneity of SKP. The complex pattern of extensional structures of the SKP, therefore, can be

described as a result pre-existing heterogeneity (Borissova et al., 2002).


54

The NW trending Raggatt Basin is a major sedimentary basin in the Southern Kerguelen Province.

Seismic data show that its eastern flank is underlain by buried volcanic basement high. ODP Site 750 that

was drilled to target this high recovered Albian claystone with coal overlying tholeiitic basalt. The ODP

data also suggests presence of Turonian to Maastrichtian limestone overlying basaltic basement,

followed by oozes and chalks with foraminifera, nannofossils and cherts in the southern part of SKP.

Williams’ Ridge: William's Ridge extends in a NNW-SSE direction as a continuation of a prominent

basement ridge mapped in the eastern part of the Central province (Figure 2.23). Seismic refraction

studies indicate that William’s Ridge is underlain by 12-15 km thick crust (Gladczenko and Coffin, 2001),

which is significantly thicker than the adjacent Labuan Basin crust, but similar to that of the Kerguelen

Plateau. It rises to only 500 m below sea level and appears to consist of two blocks separated by a

narrow fault-bounded valley.

Elan Bank: Elan Bank extends westward from the boundary between the Central and Southern

Kerguelen Province (Figure 2.23). Bathymetric, gravity and seismic data indicate that there are two large

basement highs are present within Elan Bank (Figure 1.3). These highs are displaced along a NW-SE

trending feature in the central part of the Elan Bank. The western high is relatively shallow, rising up to

500 m below sea level, whereas the eastern part is about 1000 m deep (Borissova et.al; 2002, 2003). The

predominant east-west structural trend of Elan Bank is completely different from the NW-SE trends

observed at central and southern provinces.

The seismic character of the Elan bank shows non-reflective buildups, and faulted intra-basement

reflector sequences similar to those mapped on volcanic passive margins (Borissova et al., 2003, Figure

1.5). Gneissic metamorphic and felsic igneous clasts recovered from the basaltic basement complex at

ODP Site 1137 on the bank reveal its continental origins based on the fact that they underwent distance

sediment transport mechanism (Nicolaysen et al, 2001)(Figure 1.4). The conglomerate was sourced

locally in a sub-aerial environment and deposited in a braided river environment (Borissova et al., 2002).

The basement is overlain by Campanian glauconitic sandy packstone and nannofossil pelagic oozes of

Late Eocene to Pleistocene age (Ingle et al., 2002) (Figure 2.25). The margins of the bank are

characterized by massive lava flows and highly reflective layered crust at its base. Geophysical studies

indicate that the custal structure of Elan Bank is composed of mostly continental crust (Charvis et al.,

1997; Coffin et al., 2000; Frey et al., 2000; Borissova et al., 2002; 2003)(Figure 1.4; 1.5).


55

Figure 2.25: Detail of seismic data at ODP Site 1137 (Borissova et al., 2002). Seismic log and down hole interpretation show that basement complex is overlain by Campanian sandy packstone.

Elan Bank has two distinctive upper and lower basalt flows separated by the aforementioned fluvial unit

(Coffin et al., 2002;Ingle et al., 2002) (Figure 2.26). The youngest basaltic volcanism contributing to the

upper layer of Elan Bank took place around 109 Ma (Duncun, 2002). The volcanic sequence at Elan Bank

was therefore, possibly formed either during the Early Cretaceous breakup of East India-Elan Bank and

Antarctica, or during later Albian breakup of East India and Elan Bank, when the Elan bank was

transferred from the Indian plate to the Antarctic plate via a ridge jump (Borissova et al., 2003, Gaina et

al., 2007). In either case, massive Albian volcanism overprinted and radically altered the nature of the

microcontinent forming the core of the Elan Bank.

Labuan Basin: The Labuan Basin flanking the eastern margin of the Kerguelen Plateau is the largest

sedimentary basin in the region (Figure 2.23). It is about 1000 km long, 250 km wide and contains 2.5 to


56

4 km of sediment. Structural style and sedimentary fill of the basin are highly variable. The nature of the

crust underlying the Labuan Basin and the age of its sedimentary fill are very poorly constrained.

Figure 2.26: Summary of drill sites from the Kerguelen Plateau and Broken Ridge that recovered volcanic rocks (Coffin et al., 2000). Data are shown for Sites 738, 747, 748, 749, 750, 1136, 1137, 1138, 1139, 1140, 1141, and 1142. Multichannel seismic reflection profiles indicate that the volcanic rocks at all sites except Site 748, were recovered from the uppermost igneous basement of the plateau, which lies beneath the younger sedimentary cover. Basalt at Site 748 was recovered at 200 m above the seismically defined basement. It overlies a poorly recovered zone that contains smectitic clay and highly altered basalt. Radiometric ages for basalts are shown in italics. Biostratigraphic ages of sediments overlying basement are also indicated.

The Labuan Basin is floored by a highly faulted and heterogeneous crust, which contains 4 to 8 km thick

oceanic crust in the north (Gladczenko and Coffin, 2001). Based on seismic and gravity-magnetic data

the Labuan Basin has been subdivided into three different structural domains (Figure 2.27). These are;

i. Western Labuan Basin dominated by SW facing extensional faults;


57

ii. Eastern Labuan Basin characterized by large dome-shaped basement highs lacking

magnetic signature and a very deep linear trough in the east; and

iii. Southern Labuan Basin with relatively flat basement with uninterrupted sedimentary

sequences, which is largely unaffected by faulting.

Figure 2.27: The interpreted seismo-geological cross section across the Labuan Basin and South Kerguelen Plateau (Borissova et al., 2002). Note the abrupt change in depth to basement that occurs across the basin.

The origin and nature of the crust in the Labuan Basin is likely to be heterogeneous. The Western

domain is likely to represent extended and down-faulted volcanic basement of the Kerguelen Plateau. It

is possible that large non-volcanic or non-sedimentary features, which dominate the eastern Labuan

Basin, may have formed as peridotite intrusions due to mantle exhumation, during the initial breakup.

Rotstein et al. (1991) noted that the tilted-block morphology of the Labuan Basin extends onto the

Kerguelen Plateau and that there is a reasonably good correlation between sequences on the plateau

and in the basin. They suggested that the Labuan Basin was formed by crustal extension at about the

same time as the Kerguelen Plateau itself (130-95 Ma). Metamorphic and granitic rocks dredged from a


58

basement high in the northern part of the Labuan Basin and on one of the Southern Kerguelen outcrops

may indicate the possible presence of continental fragments within the Labuan Basin crust.

A smooth basement of the Southern Labuan Basin may represent an oceanic crust of Valanginian age

trapped on the Antarctic plate.

2.5 East Antarctica Passive Margin

The Antarctic continent is divided into two large geologic areas, indicating East and West Antarctica.

East Antarctica constitutes almost two third of the Antarctic continent, lying on the Indian Ocean side of

the Transantarctic Mountains. West Antarctica is situated in western hemisphere (Figure 2.28).

The plate tectonic reconstruction indicates that the passive continental margins of Enderby and Mac

Robertson Lands in the central part of East Antarctica (between 80° to 40°E latitude) passive margin was

conjugate to Greater India (Reeves and de Wit, 2000; Veevers, 2009)(Figure 1.10). The reconstruction of

Antarctic geology has been also depicted in Gondwana context in Figure 2.29 (Harley and Kelly, 2007).

Therefore, this portion of East Antarctic passive margin and the adjacent deep ocean basins are the key

to understand the Early Cretaceous breakup between Greater India and East Antarctica.

2.5.1 Physiographic setting

The continental margin of Enderby and Mac Robertson Land is characterized by a narrow continental

shelf. The slope is dominated by spur and canyon topography. These topographic features are result of

interplay between erosional and depositional processes in cold climatic conditions over a long geological

period (Stagg et al., 2005) (Figure2.28).

The eastern reach of the Mac Robertson Land is bounded by Prydz Bay, a deep embayment of East

Antarctica. The interior part of the Prydz Bay intersects with Lambert graben, the most prominent

Permian-Triassic graben in East Antarctica (Harrowfield et al., 2005). The topography of the Prydz bay is

gentler than that of the adjacent Mac Robertson land due to erosional and depositional effects of

Lambert Glacier, which flows from Lambert Graben into the Amery Ice Shelf on the southwest side of

Prydz Bay (O’Brien et al., 2001). The Enderby Basin is bounded by Princess Elizabeth Trough to the east,

South Kerguelen Plateau and Elan Bank to the northeast and north, Kerguelen fracture zone and Conrad


59

rise to the northwest, and Gunnerus Ridge to the west. The basin floor of Enderby Land reaches more

than 5000 m, whereas the depth reaches around 3000 m in Prydz bay (Stagg et al., 2005).

Figure 2.28: Physiographic division of Antarctic continent (Source: Wikipedia).


60

Figure 2.29: Map of East Antarctica in its reconstructed Gondwana context (Harley and Kelly, 2007). AF (Albany-Fraser); BH (Bunger Hills); cDML (central Dronning Maud Land); DG (Denman Glacier); Dhar + Md (Dhawar and Madras); EG (Eastern Ghats); Gr (Grunehogna); KKB (Kerala Khondalite Belt); La (Lambert); LHB (Lützow-Holm Bay); M (Mawson); Madag (Madagascar); Nap (Napier Complex); nPCM (northern Prince Charles Mountains); Oy (Oygarden Islands); PZ (Prydz Bay); RI (Rauer Islands); Ru (Ruker); SL (Sri Lanka); sPCM (southern Prince Charles Mountains); Shack (Shackleton Range); SR (Sør Rondane) TA (Terre Adélie); VH (Vestfold Hills); wDML (western Dronning Maud Land); WI (Windmill Islands); YB (Yamato-Belgica).

2.5.2 Geological and Tectonic Setting

Geology of the East Antarctica is highly variable primarily due to size of the continent and the changing

tectonic processes, environments, and climates that it has experienced over geological time. Because of

the thick ice sheet, precise geologic details are not possible to obtain for entire Antarctica.

The major tectonic provinces of the Antarctica include (Grikurov and Mikhalskii, 2002) (Figure 2.30);


61

1) Precambrian Shield Region

2) Permian-Triassic Intra-cratonic rift setting

3) Passive margin rift setting

2.5.2.1 Precambrian Shield Region

The majority of East Antarctica is a large Precambrian shield. It is similar to shield areas in Brazil, Africa,

India, and Australia. The oldest rocks found in this area are over 3 billion years old. The shield consists

of variety of Archean and Proterozoic high grade terranes with distinct crustal evolution, which

amalgamated at various times in Precambrian and Cambrian as well (Harley and Kelly, 2007)(Figure

2.31).

The rocks of the older crystalline basement, outcropping in the shield, can be described as representing

those two major types of geologic provinces. First one is known as Napier complex, which is represented

by high-grade metamorphic rocks of Archean age (Sheraton et al., 1987; Stagg et al., 2005) (Figure 2.29).

This is best known are gneiss–granulite craton of Enderby Land, dominated by an enderbite–charnockite

association, similar to the rocks of the Peninsular India (Grikurov and Mikhalskii, 2002). A highly

deformed Archean Napier Complex presents evidences for some highest grade granulite facies

metamorphism (Harley and Motoyoshi, 2000). The granitic and tonlitic gneisses are dated as upto 3.8 Ga

and hence represent the oldest rocks from Antarctica. This Precambrian complex is also petrologically

equivalent to Dharwar Super group of India.

The second type of Precambrian geologic province includes series of Proterozoic mobile belts encircling

and separating the Early Precambrian cratons known as the Rayner complex (Black et al., 1987; Sheraton

et al., 1987; Stagg et al., 2005) (Figure 2.29). Actually, it is possible that a number of inter cratonic

mobile belts and zones has originated there during various periods of the Proterozoic orogenic events

including Mesoproterozoic Rodinia and Neoproterozoic Pan African orogeny (Black et al., 1987; Grikurov

and Mikhalskii, 2002). The Proterozoic and Early Paleozoic metamorphic and igneous processes that

operated within the mobile belt are usually interpreted as the imprints of collision-related tectonic and

thermal environments that arose at the margins. The Proterozoic fold belts are characterized by well-

layered gneisses, which include garnet-biotite gneiss, clinopyroxene gneiss, hornblende gneiss and

quartzo-feldspathic gneiss (Halpin et al., 2005).


62

Figure 2.30: Schematic map of different tectono-stratographic and geological provinces of Antarctica (Grikurov and Mikhalskii, 2002). 1 – Shirmakher Oasis, 2 – Sör-Ronnane Mountains, 3 – Lützow-Holm Gulf, 4 – Kemp Coast, 5 – Mawson Coast, 6 – Larsemann Oasis, 7 – Rëuer Ils., 8 – Pryuds Gulf, 9 – Grove nunataks, 10 – Denman Glacier, 11 – Bunger Oasis, 12 – Windmill Ils., 13 – Shackleton Range, 14 – Haag nunataks

Figure 2.31: Map of East Antarctica showing the regions with Precambrian crust of different ages based on the relict U–Pb dating of zircons in some areas and other literature data (Grikurov and Mikhalskii, 2002). Note that the blank region denotes un-interpreted regions located under thick ice sheet.


63

2.5.2.2 Permian-Triassic Intra-cratonic rift setting

The Permian-Triassic sequences in East Antarctica are localized within Intra-cratonic rifts as it is

observed in other Gondwana basins in India, Africa and Australia. The Lambert graben of East Antarctica

is a prominent, wide intra-cratonic failed rift that extended from Australia’s North West Shelf, between

India and Antarctica, to southern Africa. This rift contains thick Permian-Triassic sedimentary sequences,

which are rich in coal measures (Stagg, 1985; Cooper et al., 1991; Turner, 1991; Harrowfield et al., 2005;

Stagg et al., 2005) (Figure 2.32)

The Lambert graben intersects the modern coastline at the Prydz Bay, strikes oblique to shelf

architecture, and has a geophysical signature that can be traced into 1000 km inboard of the continent

(Harrowfield et al., 2005). The stratigraphic succession of Lambert graben includes upper Palaeozoic to

Mesozoic siliciclastic sediments. They are exposed, in part, as the Amery Group in the northern Prince

Charles Mountains, which only outcrops are found at Beaver Lake region. In that region a 270-m-thick

section of non-marine sediment is reported. This sequence consists of basal conglomerate, followed by

fluvial coal-bearing sand-shale sequence of late Permian to Triassic age (McLoughlan and Drinnan, 1997)

(Figure 2.33). In Prydz Bay, the sedimentary record shows two major sequences that includes a lower

red bed unit, which may be correlated with equivalent red bed sequences of Triassic age in India. The

upper unit with sandstone and coal is dated as Aptian unit (Turner, 1991).

It is to be noted that, although Prydz Bay is an embayment in the East Antarctic margin caused by crustal

depression due to Lambert Graben, however, interpretation of seismic data indicates that the outer part

of the bay is underlain by a basin that is separate from the Lambert Graben. In inner Prydz Bay,

basement normal faults, constrained by prominent unconformities, trend sub parallel to the Lambert

graben. However, the potential field studies shows the outer bay basin contains different structural

elements , which is at high angle to Lambert graben trend (Stagg et al., 2005) This basin is known as

Prydz Bay Basin (Stagg, 1985; Stagg et al., 2005 ). Early to Late Cretaceous sedimentary succession in this

part indicates this basin is most likely formed due to Early Cretaceous breakup of Greater India and

Antarctica.

Although Prydz Bay is an embayment in the East Antarctic margin caused by crustal depression

represented by lambert graben, it needs to be noted that the interpretation of seismic data indicates

that the outer part of Lambert graben is underlain by a basin that is separate from the Lambert graben.


64

Figure 2.32: Generalized geologic map of the Prydz Bay area showing the location of the Lambert Graben filled with Permian-Triassic sediments and drill sites on the East Antarctic continental shelf (Turner, 1991).


65

Figure 2.33: Summary of the stratigraphy, measured sections and lithology of the Amery Group in the Beaver Lake area together with locations of samples (Holdgate et al., 2005).


66

In inner Prydz bay normal faults associated with prominent unconformities, trend sub parallel to the

Lambert graben. However, the potential filed studies shows the outer bay basin contains different

structural elements, which are at high angle to the Lambert graben trend (Stagg et al., 2005). This basin

is known as the Prydz bay Basin (Stagg, 1985; Stagg et al., 2005). Lower to upper Cretaceous

sedimentary succession here indicates that this basin is most likely formed due to Early cretaceous

breakup of Greater India and Antarctica.

The thermal history modeling using Apatite Fission Track studies indicates that Permian sedimentation

of the Amery Group probably represents an initial rifting in the Lambert Graben (Lisker et al., 2007)

(Figure 2.34). The initiation of rifting was likely related to the formation of the ancestral Gamburtsev

Mountains due to Variscan compression and substantial crustal thickening of the East Antarctic Craton

(Ferraccioli et al., 2011). Most of the plate reconstructions align the Lambert graben with the Mahanadi

Valley in Eastern India (Misra et al., 1999; Veevers, 2009). However, based on vitrinite reflectance data

from Permian coal allow aligning the Lambert graben with Pranhita-Godavari Basin (Harrowfield et al.,

2005).

2.5.2.3 Passive Margin rift setting

The East Antarctica margin is a rifted and mostly non-volcanic passive margin, with structural

complexities due several breakup episodes. The first breakup has occurred in Middle Jurassic (180 Ma),

when Eastern and Western Gondwana were separated. This was followed by the Early Cretaceous (132

Ma) breakup between India and Antarctica. The last breakup event occurred in Late Cretaceous (83 Ma),

taking place between Antarctic and Australia (Reeves and DeWit, 2004) (Figure 2.4). However, the

seismic and potential field studies indicate that the main rift-bounding faults approximately coincide

with the outer edge of continental shelf. Sedimentary cover of varying thickness extends inboard of the

rift-bounding faults and this cover thins above the shallowing basement towards the coast.

The margin is divided into distinct western and eastern sectors by a strong, north–south crustal

boundary located at about 58°E. Structuring in the western sector appears to be strongly influenced by

the oblique slip/strike slip setting. In contrast, the eastern sector was formed in an orthogonally rifted

margin setting, albeit with complexities caused by the major N–S trending crustal-scale Lambert Graben

and the overlying Prydz Bay (O’Brien and Stagg, 2007).


67

Figure 2.34: Apatite fission track samples from the Lambert graben (Lisker et al., 2007) Plots on the right show a. the apatite fission track ages versus sample elevation and b. the relation between apatite fission track ages and mean track lengths for the sample profiles from the southern Mawson Escarpment. Errors are plotted as ±2σ. The sample symbols are: diamonds, Riminton Bluff; circles, Riminton Bluff S; squares, Philpot Bluff.

The seaward extent of continental margin is often controversial. However, several interpretations exist

for identification of the continent–ocean boundary (COB) from seismic data, supported by gravity and

magnetic interpretation. The nature of the continent-ocean boundary (COB) is also markedly different in

eastern and western sectors. In the eastern sector, offshore Enderby Land, the boundary between

oceanic crust and the outboard edge of highly extended and/or modified continental crust, which may

include fragments of proto-oceanic crust, is generally clearly defined on the basis of seismic character

(O’Brien and Stagg, 2007)(Figure 2.35). In contrast, the western sector shows a not very well defined and

sharp COB, where its location appears to coincide with the zone of deepest basement on the margin


68

Figure 2.35: Tectonic elements of the continental margin of East Antarctica, between 38 and 164°E (O’Brien and Stagg, 2007). In the eastern sector, offshore Enderby Land, the boundary between oceanic crust and the outboard edge of highly extended and/or modified continental crust, which may include fragments of proto-oceanic crust, is generally clearly defined on the basis of seismic character


69

(Stagg et al., 2005). The COB interpretation between regions of Princess Elizabeth Trough and the

western flank of Bruce Rise is not very well constrained as the data limit does not extend far enough

landward to define the COB, although it can be located well oceanward (Stagg et al., 2005; O’Brien and

Stagg, 2007)(Figure 2.35).

The seismic data in offshore Enderby Basin and Mac Robertson Shelf (O’Brien et al., 1995) show the

evidence of synrift Mesozoic and post rift sequences on the sea floor. However, the thickness of rift-

phase sediments in this basin is uncertain due to data constraint. The continental slope and the

landward flanks of the deep ocean basins are generally blanketed by large thickness of sediments. The

post-rift sedimentary section is particularly thick, as much as 9–10 km (O’Brien and Stagg, 2007) (Figure

2.36). One of the most conspicuous structures found in offshore East Antarctica is Terre Adélie rift block

along the eastern most margin of East Antarctica. This area is located to the southwest of the major

fracture zones which separate the Otway Basin margin of Australia from Antarctica (Colwell et al., 2006)

(Figure 2.35).

Figure 2.36: Seismic profile GA-229/35 through the offshore western Enderby Land (Stagg et al., 2005). Arrow shows an inboard edge of oceanic crust.


70

The interpretation of line GA-229/06 shows that the strike-slip movement on these fractures zones has

induced structural complexities that are not seen on lines further to the western part of East Antarctica,

where extension was essentially normal to the margin (Stagg et al., 2005, Colwell et al., 2006 (Figure

2.37). It is to be noted that the crustal density distribution is also done through potential field modeling

(Figure 2.38). The most prominent features of this line are the presence of non-uniform thinning in the

lower crust and the associated major deformation of thick possibly Cretaceous and Jurassic (?)

sediments. The deformed crustal block is known as “Adélie Rift Block”. It is bounded on its landward side

towards Antarctica mainland by a major landward- dipping fault system that sole out at deep crustal

level. It is also associated with a sediment trough within the block. The oceanward side of this feature is

characterized by igneous rocks present in the continent-ocean transition zone (Colwell et al., 2006).

Figure 2.37: Seismic detail, line GA-229/06 showing the thick, faulted possibly Cretaceous and Jurassic (?) sedimentary section of the Adélie Rift Block, which is located immediately inboard of the continent-ocean transition (Colwell et al., 2006).


71

The sedimentary succession (O’Brien et al., 1995; Truswell et al., 1999) and the crustal architecture

(O’Brien and Stagg, 2007) of the conjugate East Antarctica margin of India can be summarized as follows

(Stagg et al., 2005; O’Brien and Stagg, 2007):

1. Deposition of sediments in the early rift basin during the Toarcian– Bajocian;

2. Extension and normal faulting active from Callovian to at least the Aptian;

3. Continental shelf sedimentation from the Palaeocene to the late Eocene, possibly continuing

until the Middle Miocene; and

4. Erosion of the shelf by grounded ice since the Middle Miocene.

It is interesting to note that, until the mid-Cretaceous the passive margin evolution is consistent with the

geology of the basins of the east coast of India, Elan Bank, with fragments of the Kerguelen Plateau,

which are conjugate to the Enderby Land margin. Even the nature of COB and difference between

margin setting in the eastern and western sector of East Antarctica closely resembles orthogonally rifted

and strike slip dominated segments of conjugate East India margin.

Figure 2.38: Potential field model for line GA-229/06 through the offshore Terre Adélie rift block. COB: continent-ocean boundary (Colwell et al., 2006).


72

73

Chapter 3

Data

Diverse and voluminous data sets have been used to complete this study, including;

a. Reflection seismic data;

b. Gravity and magnetic data;

c. Borehole data (mostly for biostratigraphy and paleoenvironment); and

d. Petrological data

Some of these data are public and some confidential.

3.1 Reflection seismic data

A series of profiles along East India and Elan Bank is selected to understand the crustal architecture,

structural and tectonostratigraphic elements along the margins. The data set includes reflection:

a. Time and few depth-migrated 2D seismic profiles from East India (owned by Reliance Industries

Ltd.)

b. Time and depth-migrated 2D seismic profiles from ION-GXT Indiaspan

c. Time-migrated 2D seismic profiles from Elan Bank (owned by Geoscience Australia)

All profiles are displayed with seismic amplitude with variable density display. The data have normal

polarity convention, showing peak to though with the first hard surface (Seabed) being a peak. The

Chapter-3 Data

74

adopted colour scheme is a black – white- red scale. The black represents the peak. The red is the

trough.

3.1.1 East Indian margin data

Four key seismic profiles (Figure 3.1; 3.2; 3.3; 3.4) through the Cauvery rift zone have been selected to

show the nature of the southern horsetail of the dextral Coromandal strike-slip fault system (Nemčok et

al., 2013b; Chapter 2; Figure 2.11). The three profiles (Figure 3.1; 3.2; 3.3) are roughly NW-SE oriented,

being roughly parallel tectonic transport. All profiles image thinning of continental crust towards the

ocean-continent transition. Most of the thinning occurs over a short distance to the east of major high-

angle breakaway fault. These profiles are time-migrated reflection seismic data with the vertical axis

being in two way travel time. The record length is 10 Sec, which is sufficient to record the major crustal

boundaries in hyper-extended continental margin. The overall resolution of these profiles is good. The

main crustal are reliably imaged.

The first profile (Figure 3.1) is located in the southernmost part of the Cauvery rift zone. The profile is a

west-east trending section, oriented roughly parallel to tectonic transport. The water depth of the

profile varies from 0.5 to 3 sec two-way travel time (TWT). It has to be noted that the crustal thinning

occurs over a distance of 100 km. The section clearly demarcates the contact between crystalline

basement and its sedimentary cover as a result of the high impedance contrast between the two. In the

image, the character of basement reflectivity is relatively transparent due to low amplitude and

frequency content. It has to be noted that the lower part of the basement is quite reflective due to

relatively higher amplitude content. However, its frequency content is still low. Apart from that, there

can be certain prominent reflective events observed within the basement image. The most prominent

are two such events, which can be observed in the western end of the profile. The lower of these two

events appears to be continuous while the upper one is losing its character in the middle of the profile.

Here, the basement reflection drops from almost 2 sec to 6 sec two way travel time against the

breakaway fault. The seismic reflection is quite chaotic, with relatively higher amplitude content near

the ocean-continent transition in the south-east. The sedimentary sections are very clearly imaged with

visible layered reflection patterns. The amplitude and frequency content varies from layer to layer. It has

Chapter-3 Data

75

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Chapter-3 Data

76

to be noted that the sedimentary sections west of the breakaway fault is thinner in comparison to the

thicker section towards the deeper offshore in the eastern part of the profile. It is possible to divide the

sedimentary sections based on their own internal characters including onlap, downlap, and truncations.

One of the clear truncations can be seen on the top of the wedge-shaped sedimentary section above the

basement. This unconformity is the most prominent one in this region.

The second profile (Figure 3.2) is located to the north of the previous profile (Figure 3.1) also in the

Cauvery rift zone. This is also a west-east trending section, oriented roughly parallel to tectonic

transport. The water depth of the section varies from 0 to 3 sec two-way travel time (TWT). The crustal

thinning here is more gradual in comparison to the first profile covering a distance of 120 km. the high

impedance contrast between the sedimentary and basement sections is apparent as well. Similarly, the

lower part the basement is reflective due to high amplitude content while the frequency remains low.

The prominent discrete reflection within the basement is common. However, there are two discrete

reflections here. They can be observed at a higher angle to the continuous patchy reflective section at

the bottom of the profile. The basement reflectivity in the eastern part appears to be quite chaotic. The

basement reflection typically drops from 3 sec to 6 sec TWT against the main breakaway fault. The

sedimentary sections are quite comparable to the first profile, where the thicker sediment column is

present in the deeper water part. The internal reflection characters including onlap, downlap and

truncations are visible.

The third profile (Figure 3.3) is located further north of Profile-2 (Figure 3.2) also in the Cauvery rift

zone. This is also a west-east trending profile, oriented roughly parallel to tectonic transport. The water

depth of the section varies from 0 to 3.5 sec two-way travel time (TWT). It has to be noted that the

seabed is rugged in this profile. The crustal thinning is more gradual in comparison to the second profile,

covering a distance of 140 km. The high impedance contrast between the sedimentary and basement

sections is apparent as well. Similarly the lower part the basement is reflective due to high-amplitude

content while the frequency remains low. Here, in the western part, two small wedge-shaped basins are

imaged close to the basement bounded by faults. The highly reflective lower part of the basement

shows some degree of layer-parallel reflections. The patchy reflective boundaries are almost continuous

from west to east at a depth represented by 6 sec TWT. Two discrete reflective boundaries can be

observed at a higher angle to the patchy continuous boundary. The basement reflection drops from

initial 2 sec to 4 sec TWT against the breakaway fault and from 4 sec to 6.5 sec TWT from west to east.

Chapter-3 Data

77

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sh

allo

w t

o d

eep

er b

asi

n a

rea

is

gra

du

al

wit

h a

ma

jor

bre

aka

wa

y fa

ult

dem

arc

ati

ng

th

e b

ou

nd

ary

bet

wee

n p

roxi

ma

l a

nd

dis

tal

ma

rgin

s. S

B=S

eab

ed,

UC

C=t

op

up

per

co

nti

nen

tal

cru

st,

LCC

= t

op

lo

wer

con

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

- o

cea

nic

cru

st, C

M =

to

p c

on

tin

enta

l ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

ho

rizo

ns

(T9

0

to T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

78

Fig

ure

3.3

: Th

e re

flec

tio

n s

eism

ic P

rofi

le-3

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts W

-E

tren

din

g d

ip-o

rien

ted

sec

tio

n T

he

loca

tio

n o

f th

is p

rofi

le

is t

he

sou

ther

n C

au

very

rif

t zo

ne.

Th

e p

rofi

le r

epre

sen

ts W

- E

tre

nd

ing

dip

ori

ente

d s

ecti

on

. N

ote

hig

hly

ref

lect

ive

up

per

co

nti

nen

tal

cru

st.

The

cru

sta

l

tra

nsi

tio

n f

rom

sh

allo

w t

o d

eep

er b

asi

n a

rea

is

gra

du

al

wit

h a

ma

jor

bre

aka

wa

y fa

ult

dem

arc

ati

ng

th

e b

ou

nd

ary

bet

wee

n p

roxi

ma

l a

nd

dis

tal

ma

rgin

s.

SB=S

eab

ed, U

CC

=to

p u

pp

er c

on

tin

enta

l cru

st, L

CC

= t

op

low

er c

on

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

-oce

an

ic c

rust

, CM

= t

op

co

nti

nen

tal

ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

the

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

79

Fig

ure

3.4

: Th

e re

flec

tio

n s

eism

ic P

rofi

le-4

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts W

- E

tre

nd

ing

dip

-ori

ente

d s

ecti

on

. Th

e lo

cati

on

of

this

pro

file

is t

he

sou

ther

n C

au

very

rif

t zo

ne.

Th

e cr

ust

al t

ran

siti

on

fro

m s

ha

llow

to

dee

per

ba

sin

are

a is

gra

du

al.

No

te t

he

pro

min

ent

cen

tra

l hig

h in

th

is s

ecti

on

.

SB=S

eab

ed, U

CC

=to

p u

pp

er c

on

tin

enta

l cru

st, L

CC

= t

op

low

er c

on

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

-oce

an

ic c

rust

, CM

= t

op

co

nti

nen

tal

ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

80

This provides a terrace like appearance in the basement reflection. The sedimentary sections are quite

comparable to the one from previous profiles, where the thicker sediment column in present in the

deeper water part. The internal reflection characters, including onlap, downlap and truncations, are

visible. The unconformity is also clearly imaged. It seals the smaller wedge shaped packages bounded by

discontinuous surfaces.

The fourth profile (Figure 3.4) is located further north. This is also a west-east trending section. The

water depth of this section varies from 0 to 3.5 sec two-way travel time (TWT). This profile shows a bit

different structural architecture. Furthermore, this profile allows characterizing the deeper part of basin.

Except the common observations including the basement-sediment interface and nature of the

basement reflectivity, the most prominent feature in this profile is a central high, below which the

reflectors are folded, to the contrary to the previous profiles. Although there is a contrast between

sediment cover and basement, the reflectivity quotient is not as that in the high as previous profiles.

This high reflectivity is not continuous for sediment –basement interface as well. Here, the sediment

thickness is quite uniform from west to east in comparison to the previous profiles.

The ION profile IE-800 (Figure 3.5) in this region also shows similar features. The water depth is ranging

from 0 to 4 sec TWT. This profile is oriented roughly NW to SE and at an angle to the profile -4 (Figure

3.4), which is more close to the tectonic transport direction. In this profile, the continuous reflection is

seen across the central high. The strong basement reflection is again discontinuous in this profile,

particularly to the west of the central high. The basement reflectors are quite inclined within the central

high itself. Towards the east of the central high, the basement reflection becomes almost horizontal

towards the south-eastern part of the profile. There is a highly reflective intra-basement continuous

reflector with very high amplitude and frequency, which can be observed in the western part of the

central high. In this profile, rotated fault blocks can be observed in the north-west. The sedimentary

section appears to be quite reflective in this profile. The sedimentary section can be subdivided into

different stratigraphic sections, based on the internal character and boundaries associated with onlap,

downlap and truncations. There is a clear truncation surface on the top of the central high. It can be

mapped across it from north-west to south east. However, the sedimentary column between this

unconformity and the basement is quite variable. There is a thicker sedimentary section between the

unconformity and basement. It can be observed within the central high, which is thinner in the north-

western part and thinnest in the south-eastern part. The sediments display a thickening against the

Chapter-3 Data

81

central high on both eastern and western sides. Such thickening is more prominent in the western part,

which contains higher amplitudes.

Fig

ure

3.5

: IO

N r

efle

ctio

n s

eism

ic p

rofi

le 8

00

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sect

ion

. Th

e lo

cati

on

of

this

pro

file

is

the

sou

ther

n C

au

very

rif

t zo

ne.

Th

e cr

ust

al

tra

nsi

tio

n f

rom

sh

allo

w t

o d

eep

er b

asi

n a

rea

is

gra

du

al.

No

te t

he

pro

min

ent

cen

tra

l hig

h in

th

is s

ecti

on

. Th

e re

flec

tio

ns

wit

hin

th

e ce

ntr

al h

igh

are

ima

ged

ver

y w

ell.

Def

orm

ed a

nd

fold

ed g

eom

etry

is o

bse

rved

wit

hin

th

is h

igh

. SB

= S

eab

ed, U

CC

= t

op

up

per

co

nti

nen

tal c

rust

, LC

C =

to

p lo

wer

co

nti

nen

tal c

rust

, O

C

= to

p o

cea

nic

cru

st,

PO

C=

top

pro

to-o

cea

nic

cru

st,

CM

= t

op

co

nti

nen

tal

ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

the

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

82

There are three profiles (Figure 3.6; 3.7; 3.8), which have been selected to characterize the central part

of the Coromandal strike-slip fault system. All of them are composites. Their proximal parts are

represented by Reliance profiles. Their distal parts are represented by ION Indiaspan IE- 880 (Figure 3.6);

IE-970 (Figure 3.7) and IE-930 (Figure 3.8) profiles respectively.

The first profile through the the Coromandal strike slip system (Figure 3.6) is located in the

southernmost part of the Coromandal system to the north of Profile-4 (Figure 3.4) near the southern

horsetail structure. The profile is oriented roughly north-west to south-east and parallel to the tectonic

transport. The water depth in this profile varies from 0 to 4 sec TWT. Notably, the slope is quite steep.

The basement reflector is easy to identify, as it is imaged as a highly reflective surface due to sharp

impedance contrast. The basement reflector appears to be rugged, due to upper crustal faulting. In

contrast to the previous profiles (Figure 3.1, 3.2, 3.3, 3.4 and 3.5) in the Cauvery Basin, the crustal

thinning here occurs within a very short distance of 50 km. The main breakaway fault dips at a higher

angle in comparison to the previous profiles. The basement reflection significantly drops from initial 0.5

to 7 sec TWT against the main breakaway fault, and then from 7 to 8 sec TWT from north-west to south-

east, emphasizing a terrace like geometry. The intra basement reflection is quite transparent and

chaotic, due to low amplitude and low frequency content towards the north-western part of the profile.

However, in the south-western part, the basement character is quite different. In this part, the

basement reflector appears to be almost horizontal. The frequency content is still low, but some layer

parallelism can be observed below the basement reflector. One of such reflector at 9.5 TWT is highly

reflective and almost parallel to the basement reflector. In between these two characteristically

different basement reflection pattern, there are some more chaotic, high amplitude-low frequency

basement reflections are observed. Here, there are some small but discrete high-amplitude reflectors,

which are oriented at random angles.

The sedimentary sections are very clearly imaged with visible layer-parallel reflective surfaces.

Interestingly, there is almost no sedimentary thickness observed in the shallow water portion and thin

layer of sediments on the slope, which indicates heavy slope bypass. The main sedimentary basin occurs

only in the deep water region beyond 3.5 sec TWT bathymetry. The sedimentary sections can be

characterized and sub-divided according to their internal characteristics including onlap, downlap and

truncations. A clear fault bounded basin can be observed in the north-western part immediately to the

east of the main breakaway fault. The triangular wedge shaped sedimentary basin is almost 2 sec TWT

Chapter-3 Data

83

Fig

ure

3.6

: Th

e re

flec

tio

n s

eism

ic P

rofi

le-5

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sec

tio

n.

This

is

a c

om

po

site

pro

file

, w

her

e th

e SE

sid

e is

pa

rt o

f IO

N p

rofi

le 8

80

. Th

e p

rofi

le c

uts

th

rou

gh

th

e ce

ntr

al C

oro

ma

nd

al s

trik

e sl

ip f

au

lt z

on

e.

No

te a

sh

arp

tra

nsi

tio

n b

etw

een

pro

xim

al a

nd

dis

tal m

arg

ins

alo

ng

a m

ajo

r b

rea

kaw

ay

fau

lt.

A d

eep

ba

sin

is a

lso

fo

rmed

ad

jace

nt

to m

ain

bre

aka

wa

y fa

ult

. SB

= S

eab

ed, U

CC

= t

op

up

per

con

tin

enta

l cru

st,

LCC

= t

op

low

er c

on

tin

enta

l cru

st,

OC

= t

op

oce

an

ic c

rust

, P

OC

= to

p p

roto

-oce

an

ic c

rust

, C

M =

to

p c

on

tin

enta

l ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

the

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

Chapter-3 Data

84

thick and topped by an angular unconformity. Above this unconformity, a relatively transparent section

is present. On top of that, a prominent onlap surface is observed, roughly parallel to 6 sec TWT level.

The second profile in the central part of the Coromandal strike slip system (Figure 3.7) is located to the

north of the previous profile (Figure 3.6). The profile is oriented roughly north-west to south-east,

parallel to the tectonic transport. The water depth in this profile varies from 0 to 4 sec TWT. In this

profile, the observations are quite similar to those from the previous one (Figure 3.6). The crust is

thinned over a very short distance of 50 km, with a very steeply dipping breakaway fault. The basement

reflector is easy to identify due to its inherent impedance contrast. Against the breakaway fault, the

basement reflector drops from 1 to 6 sec TWT. Here, the intra-basement reflections can be categorized

into three distinct zones from north-west to south-east. The north-western part represents chaotic and

relatively transparent character, while the central part represents high amplitude-low frequency chaotic

reflection with occasional randomly oriented discrete high-amplitude reflectors. The south-eastern part

contains low amplitude with relatively higher frequency content with some layer parallelism and

presence of the highly reflective basement parallel reflector at around 9.5 sec TWT level. The sediment

thickness in deep-water region is thinner than in shelf and slope regions. This includes some amount of

slope bypassing as well. The sedimentary sequences can be sub-divided as based on their internal

character including onlap, downlap and truncations.

The third profile through the central part of Coromandal strike slip system (Figure 3.8) is located to the

north of previous profile (Figure 3.7). This profile is oriented roughly north-west to south-east, parallel

to the tectonic transport. The water depth in this profile varies from 0 to 4 sec TWT. In this profile, the

observations are again almost the same in comparison to the previous ones (Figure 3.6, 3.7). The crust is

thinned over a very short distance of 40 km, with a very steeply dipping breakaway fault. Against the

breakaway fault, the basement reflector, which can be identified using same criteria as those used in the

previous profiles, it drops from 1 to 6 sec TWT. Then it again drops from 6 to 8 sec TWT before

becoming horizontal. It again, displays a terrace like geometry of the basement reflector. Here, the

intra-basement reflections can be categorized into three distinct zones, from north-west to south-east.

The north-western part represents chaotic and relatively transparent character, while the central part

represents a high amplitude-low frequency chaotic reflection with occasional randomly oriented high

amplitude discrete reflectors. In the central part, however, a characteristic more or less continuous

reflector is observed. It rises from the north-western part to the south-east towards the horizontal

Chapter-3 Data

85

basement reflector. The south-eastern part is, again, low amplitude, with relatively higher frequency

content, with some layer parallelism. Like in previous profiles (Figure 3.6 and 3.7), the presence of the

highly reflective basement-parallel reflector is not clear in this profile due to a high noise content in the

data. The thicker sediment can be found in the deep-water region. Less sediment characterizes in shelf

and slope regions. This indicates some amount of slope bypass as well. The sedimentary sequences can

be sub-divided as based on their internal character including onlap, downlap and truncations.

Another important profile used in this study is ION profile IE-900 (Figure 3.9). This profile is located

between the profiles (Figure 3.6 and 3.7) through the central part of the Coromandal strike-slip system.

The profile is oriented roughly north-west to south-east, parallel to the tectonic transport. The water

depth in this profile varies from 0 to 4 sec TWT. Here, also the crustal thinning is also supposed to occur

within a distance of 50 km. The observations are almost similar to those in the previous profiles but the

image quality is better. General observations, in accordance with those made in previous profiles (Figure

3.6, 3.7 and 3.8) include basement reflectivity character, three distinct zones typical for the intra-

basement reflection pattern, and a high angle breakaway fault against which basement reflection drops

from initial 1.2 to 6 sec TWT. The terrace like geometry is also quite evident in this profile so is a thick (2

Sec TWT) fault bounded sedimentary wedge immediately east of the breakaway fault, topped by a clear

angular unconformity (Figure 3.6). The sediment thickness can be mostly observed in the deep-water

part. There are not many internal deformations within the sediments above the unconformity. One of

the primary observations in this profile includes the sediment thickness in shallow water, which is

comparatively higher than that in the previous profiles but there is a complete slope bypass here,

because there is almost no sediment deposited along the steep slope. The sedimentary sequences can

be, again, clearly subdivided according to their internal character including onlap, downlap and

truncations.

The northern part of the study area is represented by the roughly orthogonally extended Krishna-

Godavari rift zone (Nemčok et al., 2013b). The southern part of the Krishna Basin represents an oblique

rift zone, which is also the northern end of the Coromandal horsetail structure (Nemčok et al., 2013b;

Figure 2.12). Four key seismic profiles (Figure 3.10; 3.11; 3.12; 3.13) have been selected to showcase the

major seismic characteristics of the northern horsetail structure and orthogonal rift zone. These profiles,

like those through the Cauvery rift zone, are also oriented parallel to the tectonic transport roughly from

north-west to south-east.

Chapter-3 Data

86

Fig

ure

3.7

: Th

e re

flec

tio

n s

eism

ic P

rofi

le-6

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sec

tio

n.

This

is a

co

mp

osi

te

pro

file

, w

her

e th

e SE

sid

e is

pa

rt o

f IO

N p

rofi

le 9

30

. Th

e p

rofi

le c

uts

th

rou

gh

th

e ce

ntr

al

Co

rom

an

da

l st

rike

-slip

fa

ult

zo

ne.

No

te a

sh

arp

tra

nsi

tio

n

bet

wee

n p

roxi

ma

l an

d d

ista

l ma

rgin

s a

lon

g a

ma

jor

bre

aka

wa

y fa

ult

. A

dee

p b

asi

n is

als

o f

orm

ed a

dja

cen

t to

ma

in b

rea

kaw

ay

fau

lt.

SB =

Sea

bed

, UC

C

= to

p u

pp

er c

on

tin

enta

l cru

st, L

CC

= t

op

low

er c

on

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

-oce

an

ic c

rust

, CM

= t

op

co

nti

nen

tal m

an

tle,

OM

= to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

the

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

Chapter-3 Data

87

Fig

ure

3.8

: Th

e re

flec

tio

n s

eism

ic P

rofi

le-7

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sec

tio

n.

This

is

a c

om

po

site

pro

file

, w

her

e th

e SE

sid

e is

pa

rt o

f IO

N p

rofi

le 9

70

. Th

e p

rofi

le c

uts

th

rou

gh

th

e ce

ntr

al

Co

rom

an

da

l st

rike

-slip

fa

ult

zo

ne.

No

te t

he

sha

rp t

ran

siti

on

bet

wee

n p

roxi

ma

l an

d d

ista

l ma

rgin

s a

lon

g a

ma

jor

bre

aka

wa

y fa

ult

. A

dee

p b

asi

n is

als

o f

orm

ed a

dja

cen

t to

ma

in b

rea

kaw

ay

fau

lt.

SB =

Sea

bed

, U

CC

=

top

up

per

co

nti

nen

tal

cru

st,

LCC

= t

op

lo

wer

co

nti

nen

tal

cru

st,

OC

= t

op

oce

an

ic c

rust

, P

OC

= to

p p

roto

-oce

an

ic c

rust

, C

M =

to

p c

on

tin

enta

l m

an

tle,

OM

=

top

oce

an

ic m

an

tle.

Th

e d

eta

il o

f th

e h

ori

zon

s (T

90

to

T1

0)

an

d (

K1

00

to

K2

0)

is g

iven

in T

ab

le-3

.1. S

ee t

ext

for

det

aile

d in

terp

reta

tio

n

Chapter-3 Data

88

Fig

ure

3.9

: IO

N r

efle

ctio

n s

eism

ic p

rofi

le 9

00

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e p

rofi

le r

epre

sen

ts N

W t

o S

E tr

end

ing

dip

-ori

ente

d s

ecti

on

. Th

e p

rofi

le i

s

loca

ted

wit

h c

entr

al

Co

rom

an

da

l st

rike

-slip

fa

ult

zo

ne.

No

te t

he

sha

rp t

ran

siti

on

bet

wee

n p

roxi

ma

l a

nd

dis

tal

ma

rgin

s a

lon

g a

ma

jor

bre

aka

wa

y fa

ult

. A

dee

p b

asi

n i

s a

lso

fo

rmed

ad

jace

nt

to m

ain

bre

aka

wa

y fa

ult

. SB

= S

eab

ed,

UC

C =

to

p u

pp

er c

on

tin

enta

l cr

ust

, LC

C =

to

p l

ow

er c

on

tin

enta

l cr

ust

, O

C =

to

p

oce

an

ic c

rust

, P

OC

= to

p p

roto

-oce

an

ic c

rust

, C

M =

to

p c

on

tin

enta

l m

an

tle,

OM

= t

op

oce

an

ic m

an

tle.

Th

e d

eta

il o

f th

e h

ori

zon

s (T

90

to

T1

0)

an

d (

K1

00

to

K2

0)

is g

iven

in T

ab

le-3

.1. S

ee t

ext

for

det

aile

d in

terp

reta

tio

n

Chapter-3 Data

89

The first profile from this group is the southernmost profile in the Krishna-Godavari rift zone (Figure

3.10). This profile is a composite one from shallow to deep-water region. The water depth ranges from

less than 1 to 3 Sec TWT. Here, the crustal thinning takes place over a distance of 130-140 km. The

basement-sediment interface is highly reflective and continuous throughout the profile, due to high

impedance contrast between proximal north-western to distal south-eastern part. The proximal and

distal parts are separated by a breakaway fault. The basement reflector drops down from initial 1 to 3.5

sec TWT and then from 3.5 to 5 sec TWT, against the main breakaway fault. It further drops down from

6 to 7.5 Sec TWT in the distal south-eastern part of the profile. The entire basement reflector is thus

provides a terraced like geometry, like that in the previous profiles observed in the Cauvery and

Coromandal regions. The amplitude and frequency content within the basement is quite noisy in the

proximal part. However, in the southern part it is quite transparent due to low amplitude content. Only

the lower part of the section is quite reflective and shows some kind of layered geometry. The proximal

part of the profile images a large fault bounded sedimentary basin. The geometry of the basin suggests

that it is a proximal graben with a relatively thick sedimentary fill. The fill can be subdivided into several

sequences based on their internal characteristics including onlap, downlap and truncations. The

sedimentary section to the east of the main breakaway fault in deep-water region is thicker than that in

the proximal part. However, the size of the individual grabens in the distal margin, which are also

characterized by fault founded sedimentary wedges, is smaller than that in proximal margin. These

grabens are overlain with erosional unconformity.

The next profile (Figure 3.11) is located to the north of the previous one (Figure 3.10). This profile is,

again, a composite one from shallow to deep-water regions. Its character is almost similar to that of

previous profile with respect to frequency content, amplitude response and impedance contrast

demarcating sediment-basement interface (Figure 3.10). The water depth ranges from less than 1 to 3

Sec TWT. Like the previous profile, the crustal thinning is gradual and takes place over a distance of 130-

140 km. The basement-sediment interface is highly reflective and continuous throughout the profile

from proximal north-west to distal south-east parts. The proximal and distal parts are separated by a

breakaway fault. The basement reflector drops down from initial 1 to 5 sec and then from 5 to 7.5 sec

TWT against the main breakaway fault. Therefore, the terrace like geometry of the basement reflector is

quite obvious. The basin-bounding fault is also evident in this profile. The intra-basement reflection

characteristics are common in terms of amplitude –frequency content, where the upper part remains

transparent and the lower part is more reflective. It is interesting to note that the dip of the breakaway

Chapter-3 Data

90

fault is relatively steeper than the dip of other normal faults in both profiles (Figure 3.10 and 3.11). A

proximal margin graben with thicker sedimentary fill and a series of distal margin grabens with thinner

sedimentary fill are also present in this section. The sedimentary sequences can be, again, subdivided as

based on their internal character including onlap, downlap and truncations.

The next profile (Figure 3.12) represents the northern part of the orthogonally extended Krishna-

Godavari rift zone. The profile is oriented roughly from northwest to southeast, parallel to the tectonic

transport. It extends from shallow-water to deep-water. This profile is also a composite one, where the

proximal north-western part is a Reliance profile with the south-eastern distal one being the ION profile

IE-1240. The water depth ranges from less than 1 sec TWT in shallow water to 3 sec TWT in the deep-

water regions. The total crustal thinning takes place over a distance of 120-130 km. Although the

proximal part of this profile is quite noisy, being composed of legacy seismic lines in East India, the

typical sediment-basement interface can be identified. This basement reflector is prominent due to its

high reflectivity and amplitude content towards the distal margin section. The proximal and distal

margin parts are separated by a breakaway fault. The basement reflection gradually drops from initial 1

to 5 sec TWT though gradual faulting, and then it drops suddenly from 5 to 7 sec TWT against the main

breakaway fault. It has to be noted that the throw against the breakaway fault is less (2 sec TWT) in

comparison to that in the other previous profiles, where the throw was in order of 4-5 sec TWT. The

basement reflection is also discontinuous in the proximal margin, where several fault bounded

sedimentary wedges are present. In the distal margin, the basement is transparent and chaotic due to

its low amplitude, as well as frequency content. However, the lower part of the basement is highly

reflective with some layer parallelism. The sedimentary wedges in the distal margin to the east of the

breakaway fault are quite small in size. They contain some high amplitude –low frequency reflective

bodies within. The basement reflection becomes almost horizontal in the south-eastern deep-water part

of the profile. In this part, the intra-basement section is quite transparent due to lack of amplitude.

However, some layer parallel dim reflectivity can be observed in this zone. The sediment thickness

increases in the deep-water region. The sedimentary sequences can be, again, subdivided according to

their internal character including onlap, downlap and truncations. A prominent truncation surface,

which clearly envelops all fault affected sedimentary sections, can be recognized as a major angular

unconformity. There are certain bulges in the younger sedimentary section. However, these bulges do

not react to the basement architecture.

Chapter-3 Data

91

Fig

ure

3.1

0:

The

refl

ecti

on

sei

smic

Pro

file

-8 t

hro

ug

h E

ast

Ind

ian

ma

rgin

. Th

e co

mp

osi

te p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sec

tio

n. T

he

pro

file

is l

oca

ted

in

th

e n

ort

her

n h

ors

eta

il st

ruct

ure

of

the

Co

rom

an

da

l fa

ult

sys

tem

. Th

e p

roxi

ma

l a

nd

dis

tal

ma

rgin

s a

re s

epa

rate

d b

y a

ma

jor

bre

aka

wa

y fa

ult

.

No

te t

ha

t th

e g

rab

ens

in t

he

dis

tal

ma

rgin

are

mu

ch s

ma

ller

is s

ize

tha

n t

ho

se i

n p

roxi

ma

l m

arg

in.

SB

= S

eab

ed,

UC

C =

to

p u

pp

er c

on

tin

enta

l cru

st,

LCC

=

top

low

er c

on

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

-oce

an

ic c

rust

, CM

= t

op

co

nti

nen

tal m

an

tle

, OM

= t

op

oce

an

ic m

an

tle.

Th

e d

eta

il o

f th

e

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

92

Fig

ure

3.1

1:

The

refl

ecti

on

sei

smic

Pro

file

-9 t

hro

ug

h E

ast

Ind

ian

ma

rgin

. Th

e co

mp

osi

te p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip-o

rien

ted

sec

tio

n. T

he

pro

file

is lo

cate

d in

th

e n

ort

her

n h

ors

eta

il st

ruct

ure

of

the

Co

rom

an

da

l fa

ult

sys

tem

. Th

e p

roxi

ma

l an

d d

ista

l ma

rgin

s a

re s

epa

rate

d b

y a

ma

jor

bre

aka

wa

y fa

ult

.

No

te t

ha

t th

e g

rab

ens

in t

he

dis

tal m

arg

in a

re m

uch

sm

alle

r is

siz

e th

an

th

ose

in p

roxi

ma

l ma

rgin

. S

B =

Sea

bed

, U

CC

= t

op

up

per

co

nti

nen

tal c

rust

, LC

C =

top

low

er c

on

tin

enta

l cru

st, O

C =

to

p o

cea

nic

cru

st, P

OC

= to

p p

roto

-oce

an

ic c

rust

, CM

= t

op

co

nti

nen

tal m

an

tle,

OM

= t

op

oce

an

ic m

an

tle.

Th

e d

eta

il o

f th

e

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

93

Fig

ure

3.1

2:

The

refl

ecti

on

sei

smic

Pro

file

-10

th

rou

gh

Ea

st In

dia

n m

arg

in. T

he

com

po

site

pro

file

rep

rese

nts

NW

- S

E tr

end

ing

dip

ori

ente

d s

ecti

on

wh

ere

the

SE s

ide

is p

art

of

ION

pro

file

12

40

. Th

e p

rofi

le is

loca

ted

in t

he

no

rth

ern

rif

t zo

ne.

Th

e cr

ust

al t

hin

nin

g f

rom

sh

allo

w t

o d

eep

er b

asi

n a

rea

is g

rad

ua

l wit

h a

ma

jor

bre

aka

wa

y fa

ult

dem

arc

ati

ng

pro

xim

al

an

d d

ista

l m

arg

ins.

No

te t

ha

t th

e g

rab

ens

in t

he

dis

tal

ma

rgin

are

mu

ch s

ma

ller

is s

ize

tha

n t

ho

se i

n

pro

xim

al m

arg

in.

SB =

Sea

bed

, U

CC

= t

op

up

per

co

nti

nen

tal c

rust

, LC

C =

to

p lo

wer

co

nti

nen

tal c

rust

, O

C =

to

p o

cea

nic

cru

st,

PO

C=

top

pro

to-o

cea

nic

cru

st,

CM

= t

op

co

nti

nen

tal m

an

tle,

OM

= t

op

oce

an

ic m

an

tle.

Th

e d

eta

il o

f th

e h

ori

zon

s (T

90

to

T1

0)

an

d (

K1

00

to

K2

0)

is g

iven

in T

ab

le-3

.1. S

ee t

ext

for

det

aile

d

inte

rpre

tati

on

.

Chapter-3 Data

94

The next profile (Figure 3.13) is located to the north the previous profile (Figure 3.12). Roughly oriented

from northwest to southeast and parallel to the tectonic transport, this profile also extends from

shallow water to deep water. Again, it is a composite one where the proximal north-western part is a

Reliance profile with the south-eastern distal one being the ION profile IE-1290. This profile also shows

characteristics common to the previous profile in terms of amplitude-frequency content of the data,

continuity and geometry of the basement reflector, intra-basement reflection characteristics and

sedimentary characters (Figure 3.12). The water depth ranges from less than 1 sec TWT in shallow water

to 3.5 sec TWT in the deep-water regions. The total crustal thinning takes place over a distance of 110-

120 km. towards the proximal margin, the basement reflection is quite discontinuous due to low

frequency and high noise content in the data. The fault bounded sedimentary wedges can be observed.

They are sealed by a top truncation surface. This is a large basin-bounding fault in the north-western

end of the profile. The geometry of basement configuration is clear in the distal part of the section.

Here, the basement reflection drops from initial 1 to 4.5 sec TWT through a series of breakaway fault

splaying off the main breakaway fault. Then it finally drops down from 4.5 to 6 sec TWT, before

becoming almost horizontal in the distal south-eastern end of the profile. The intra-basement reflectors

in proximal margin are quite chaotic and low frequency, due to homogeneous noise content. The intra-

basement reflectors in the south eastern region below the near horizontal basement are quite

transparent, showing some degree of parallelism to the basement. There are few low-frequency layers

with high reflectivity above the prominent basement reflector. The sedimentary section is relatively

thicker in the deep water region. The internal geometries of the sedimentary cover including the onlap,

downlap and truncations are used to subdivide the sedimentary cover into different sequences. It has to

be noticed that here, the slope bypass occurs at a late stage, as indicated by presence of relatively thick

sediments in the slope region. There are certain bulges in the younger sedimentary section. They do not

react to the basement architecture.

The seismic data described above helped to interpret the major tectono-stratigraphic surfaces and

tectono-sequences, based on internal reflection characteristics and equivalent tectonic and major

climatic events associated with this margin. This approach was applied to in all seismic profiles in this

study. The same has been summarized in Table 3.1.

Chapter-3 Data

95

Fig

ure

3.1

3:

The

Ref

ecti

on

sei

smic

Pro

file

-11

th

rou

gh

Ea

st I

nd

ian

ma

rgin

. Th

e co

mp

osi

te p

rofi

le r

epre

sen

ts N

W -

SE

tren

din

g d

ip o

rien

ted

sec

tio

n

wh

ere

the

SE s

ide

is p

art

of

ION

pro

file

12

90

. Th

e p

rofi

le is

loca

ted

in t

he

no

rth

ern

rif

t zo

ne.

Th

e cr

ust

al t

hin

nin

g f

rom

sh

allo

w t

o d

eep

er b

asi

n a

rea

is

gra

du

al w

ith

a m

ajo

r b

rea

kaw

ay

fau

lt d

ema

rca

tin

g p

roxi

ma

l an

d d

ista

l ma

rgin

s. N

ote

th

e g

rab

ens

in t

he

dis

tal m

arg

in a

re m

uch

sm

alle

r is

siz

e th

an

tho

se in

pro

xim

al m

arg

in.

SB =

Sea

bed

, UC

C =

to

p u

pp

er c

on

tin

enta

l cru

st,

LCC

= t

op

low

er c

on

tin

enta

l cru

st,

OC

= t

op

oce

an

ic c

rust

, P

OC

= to

p p

roto

-

oce

an

ic c

rust

, C

M =

to

p c

on

tin

enta

l ma

ntl

e, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ail

of

the

ho

rizo

ns

(T9

0 t

o T

10

) a

nd

(K

10

0 t

o K

20

) is

giv

en in

Ta

ble

-3.1

.

See

text

fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

96

Table 3.1: Tectono-stratigraphic chart for East coast India used for identification of seismic sequence boundaries.

Chapter-3 Data

97

The entire ION-Geophysical IndiaSpan East dataset has been used in this study to understand the crustal

architecture (Figure 3.14). The IndiaSpan was a multi-client seismic API program for regional geological

study in east and west coast of India acquired by ION GXT. Phase I of IndiaSpan covers entire India and

comprises about 13,574 lkm of 2D images. It is a high quality dataset with a long offset (10 km) and a

deep record length (18 sec), with source strength approximately 170b-m peak-peak. The data is

available in both Pre Stack Time Migration (PSTM) and Pre Stack Depth Migration (PSDM) format. In the

depth domain the data is available up to 24-30km. For the present study in East coast of India, the total

data quantum is 11587 lkm (8087 lkm + 3500 lkm) including both Phase-1 and Phase-1 infill data.

Although the data was acquired on a very broad regional scale, the data quality is much better for

regional interpretation and crustal architecture studies, because of its continuity, coverage and deep

penetration.

Usually, the ideal reflection seismic data for the interpretation of the crustal type distribution preferably

should be deep enough to image the Mohorovicic discontinuity (Moho) to provide the image of the

entire crust. Although, the crustal architecture is documented by various profiles in the study area, the

profile ION IE-1000 is the best one for deep crust documentation (Figure 3.15). Profile ION IE-1000

represents the example of the margin formed by the roughly orthogonal extension controlling normal

fault patterns. Profile ION IE-1000, like many other GXT regional profiles, is long enough to show not

only the continental margin in the north-west but also a deep ocean basin towards the south-east. The

crustal thinning in this profile occurs within 130-140 km distance.

In this profile, there are several rotated fault blocks in the proximal margin, with relatively high-angle

normal faults. The upper part of the basement in the NW is relatively chaotic and noisy, with high

frequency content. However, the deeper basement is quite reflective and layered in an anastomosing

pattern. One can observe the lowermost strong reflector in NW, which is rising towards SE. The crust is

thinned oceanward and the top of the crystalline basement is gradually becoming flat towards the deep-

water region. The same is imaged by a strong continuous near horizontal reflector. Interestingly, there

are two continuous reflectors below the flat deeper basin basement, which do not reach the acoustic

impedance contrast typical for the boundary between the top basement and the overlying sedimentary

cover. These are not continuous throughout the profile but lose their reflectivity towards west. The

section between these two major reflectors is quite transparent due to low amplitude content, but

distinctly and weakly layered. The section between the inclined basement and the near-horizontal

basement reflector is again quite distinct. While the top part of this region is imaged by higher

Chapter-3 Data

98

reflectivity zone, the base in most cases does not have any distinct signature, having mostly chaotic

reflection with some discrete high reflectivity surfaces at random angles. There is no continuous

basement reflection in this part. Similar observations can be made in ION-GXT profile IE-1200 and IE-

1600 (Figure 3.16, Figure 3.17).

The ION profile iE-1200 is located in the Krishna Godavari Basin (Figure 3.14) oriented in the north-west

to south-east direction. On the proximal margin side, the basement reflector is highly reflective but

quite discontinuous, as it is affected by faulting. The proximal margin faults are usually dipping westerly

against the easterly dipping main basin-bounding fault. The distal margin is quite distinctive and

separated by a major breakaway fault. The throw against the breakaway fault is low (2 Sec TWT) like

that in the previous profile (Figure 3.12). It has to be noted that the distal margin grabens contain

thinner sediment thickness in comparison to the proximal margins. The grabens are sealed by a

truncation surface. The basement reflectivity is quite transparent in the upper part due to low

reflectivity and frequency content. However, the lower part of the basement is highly reflective. Some

clear and distinct reflective surfaces can be observed, which are rising from below the basement

towards the deep-water part, towards the ocean-continent transition. It can be noted that a highly

reflective fault is truncating against this uprising bright reflector. The basement reflector is near

horizontal and continuous in the south-eastern region. The intra-basement reflectivity below this

section is transparent but some degree of layer parallelism to the basement can be observed. The major

difference of this profile from the previous one is an absence of a continuous reflector below the deep-

water basement. Another difference is the layers on top of the deep-water basement, which are highly

reflective and contains lower frequency.

The ION profile IE-1600 is a northwest-southeast oriented (Figure 3.17). The proximal margin can be

seen in this image. The proximal margin is separated from distal margin through a high-angle breakaway

fault. This profile is located in the close proximity of the post-rift hotspot volcanic province in the

Mahanadi Basin in the north (Bastia et al., 2010). Here, the reflectivity in the top basement may be

related to acoustic impedance contrast between the basalt and sedimentary systems. Here, also a series

of grabens can be observed in the distal margin. However, the amplitude content within the graben is

quite high with frequency being quite low. The grabens are topped by a clear truncation surface.

Towards the deep-water south-eastern part of the profile, again, basement is becoming near-horizontal

and there is some layer parallel reflectivity. However, the intra-basement reflectivity gradually becomes

chaotic towards the north-west.

Chapter-3 Data

99

Figure 3.14: Map of the study area, offshore East India, showing the location of the IndiaSpan reflection seismic

profiles (ION, IndiaSpan report, 2007), onshore topography, drainage system and offshore bathymetry.

Chapter-3 Data

100

Fig

ure

3.1

5:

Ref

lect

ion

sei

smic

pro

file

IO

N/G

XT

-10

00

th

rou

gh

th

e K

rish

na

rif

t zo

ne

sho

win

g a

dis

trib

uti

on

of

cru

sta

l ty

pes

in

clu

din

g

thin

ned

up

per

co

nti

nen

tal

cru

st (

UC

C),

th

inn

ed l

ow

er c

on

tin

enta

l cr

ust

(LC

C),

pro

to-o

cea

nic

cru

st (

un

roo

fed

co

nti

nen

tal

ma

ntl

e),

an

d

oce

an

ic c

rust

(O

C),

ma

pp

ed u

sin

g t

he

seis

mic

sig

na

ture

s. T

he

con

tin

enta

l M

oh

o u

pri

sin

g t

ow

ard

s th

e SE

is

very

wel

l im

ag

ed.

The

top

of

the

con

tin

enta

l ba

sem

ent

(TC

B)

is f

au

lted

, in

dic

ati

ng

th

e b

ritt

le n

atu

re o

f th

e u

pp

er c

rust

. Th

e to

p o

f th

e lo

wer

co

nti

nen

tal c

rust

(TL

CC

) is

ma

pp

ed a

s b

ase

d o

n t

he

ob

serv

ati

on

s d

escr

ibed

in

th

e te

xt.

Tow

ard

s th

e SE

, th

e m

ore

org

an

ized

la

yere

d n

orm

al

oce

an

ic c

rust

wit

h a

pro

min

ent

Mo

ho

res

olu

tio

n (

oce

an

ic M

oh

o)

can

be

ob

serv

ed.

The

top

oce

an

ic b

ase

men

t (T

OB

) is

re

lati

vely

fla

t a

nd

tec

ton

ica

lly

un

dis

turb

ed.

Bet

wee

n t

he

exte

nd

ed c

on

tin

enta

l cr

ust

an

d t

he

oce

an

ic c

rust

lie

s th

e p

roto

-oce

an

ic c

rust

, re

pre

sen

ted

by

the

un

roo

fed

con

tin

enta

l lit

ho

sph

eric

ma

ntl

e. I

t is

ch

ara

cter

ized

by

the

cha

oti

c se

ism

ic r

efle

cto

r p

att

ern

an

d p

rese

nce

of

thro

ug

h-g

oin

g f

ract

ure

zo

nes

.

See

text

fo

r fu

rth

er d

eta

ils o

n s

eism

ic c

ha

ract

er o

f in

volv

ed c

rust

al t

ypes

.

Chapter-3 Data

101

Fig

ure

3.1

6:

Det

ail

of

the

refl

ecti

on

sei

smic

pro

file

GX

T-1

20

0 t

hro

ug

h t

he

Go

da

vari

rif

t zo

ne

sho

win

g t

he

con

tin

enta

l an

d p

roto

-oce

an

ic c

rust

s. N

ote

th

at

the

top

of

the

pro

to-o

cea

nic

cru

st (

un

roo

fed

ma

ntl

e) is

co

vere

d b

y vo

lca

nic

ro

cks.

UC

C -

up

per

co

nti

nen

tal c

rust

; LC

C -

lo

wer

co

nti

nen

tal c

rust

; TL

CC

- t

op

of

the

low

er c

on

tin

enta

l cru

st.

Chapter-3 Data

102

Fig

ure

3.1

7:

Ref

lect

ion

sei

smic

pro

file

GX

T-1

60

0 t

hro

ug

h t

he

Ma

ha

na

di r

ift

zon

e. T

he

up

risi

ng

co

nti

nen

tal M

oh

o is

no

t im

ag

ed w

ell i

n t

his

pro

file

an

d t

he

oce

an

ic M

oh

o i

s n

ot

very

pro

min

ent.

Tw

o g

rab

ens

can

be

seen

on

th

e co

nti

nen

tal

ma

rgin

. Th

e o

cea

nic

cru

st i

s fl

at

wit

h p

ara

llel-

laye

red

ref

lect

ors

in

its

up

per

pa

rts.

Th

e se

ism

ic im

ag

e o

f th

e u

nro

ofe

d m

an

tle

is d

istu

rbed

an

d c

ha

ract

eriz

ed b

y w

avy

ref

lect

ors

. Th

e to

p o

f th

e co

nti

nen

tal b

ase

men

t (T

CB

) a

nd

top

of

the

oce

an

ic b

ase

men

t (T

OB

) a

re in

terp

rete

d o

n t

he

ba

sis

of

the

last

pro

min

ent

con

tin

uo

us

refl

ecto

r in

co

nti

nen

tal a

nd

oce

an

ic c

rust

s, r

esp

ecti

vely

.

Chapter-3 Data

103

Based on afore mentioned observations the following crustal horizons have been interpreted for East

India. These are;

I. Top upper continental crust (UCC);

II. Top lower continental crust (LCC);

III. Continental Moho (CM);

IV. Top Proto-Oceanic Crust (POC);

V. Top Oceanic Crust (OC); and

VI. Oceanic Moho (OM).

The same principle for interpreting crustal boundaries was applied to the Elan Bank images.

3.1.2. Elan Bank data

The time-migrated 2D seismic data from the Elan Bank (Figure 3.18) were obtained from Geoscience

Australia, an agency of the Australian federal government for geoscientific research. The seismic data

encompasses the survey S-179, which was acquired over the Elan Bank and part of the Kerguelen

Plateau. There are seven seismic profiles acquired (both strike and dip profiles) in survey S-179, which

mostly cover the Elan Bank. The data quality is good, having a sufficient record length of 12 sec TWT,

required for the crustal interpretation, although geophysical artifacts including multiples, noise and

sideswipes are found in some profiles.

All seven profiles through the Elan Bank have some common distinct characteristics. In all profiles, the

first strong reflector below seabed showing equivalent impendence contrast for crystalline basement

may be equivalent to top basalt. The strong contrast between physical properties of basalt and overlying

sediments in terms of both density and velocity provides a strong impedance contrast in seismic imaging

between the two rock sequences. There are significant reflections observed also below the basalt layer.

The sedimentary cover on top of the basalt is quite thin and condensed. The reflections in the

sedimentary cover are continuous, showing the internal geometries and stratal termination patterns

including onlap, downlap and truncation, and small scale faults.

Chapter-3 Data

104

Figure 3.18: The 2D seismic grid in Elan Bank. The 2D seismic was acquired and processed by Geoscience Australia for scientific research. The data are part of survey 179, which was focussed on Elan Bank.

Profile 1 (Figure 3.19) is located in the central part of the Kerguelen Plateau (Figure 3.18). This profile is

a north-east to south-west trending profile oriented roughly parallel to dip direction. It has to be noted

that this profile does not pass through the Elan Bank microcontinent. The bathymetry is 1.8 sec TWT

only. The bathymetry is quite flat and undisturbed in this profile. There is an ODP well 1138 situated on

this profile. The first prominent reflective surface in the image is assumed to be basalt top, which can be

characterized by high amplitude and low frequency content. Additionally, this can be further correlated

with stratigraphic picks of the well, which penetrates through the top of the basalt layer. The basalt

layer is also near-horizontal at around 2.5 sec TWT level. The sedimentary section above the basalt is

quite thin (1 sec TWT thick). A series of alternate low amplitude transparent and high amplitude bright

layers can be observed. There are not many internal stratal termination patterns in the sedimentary

cover. There is a low amplitude transparent package immediately below the top basalt reflector.

Although the layer is almost transparent, still some occasional reflectors with higher reflectivity can be

mapped in it. They are truncating against the top basalt layer at an angle. The base of the package is also

almost horizontal at a depth level represented by 3 sec TWT. The section below this package is quite

Chapter-3 Data

105

noisy due to high amplitude and very low frequency content, which increase further down. However,

occasional discontinuous high amplitude reflective surfaces are present. These surfaces appear to be

near horizontal as well. Therefore, no significant crustal thinning can be recognized in this profile.

Figure 3.19: The Reflection seismic Profile-1 through the central Kerguelen Plateau. The profile represents a NE-SW trending dip oriented section. The ODP site 1138 is located at this profile. S1 = Sediment Unit 1, B1 = top Basalt, IB1 = top intra-basalt unit, OC = top oceanic crust, OM = top oceanic mantle. The details of the horizons are given in Table-3.2. See text for the detailed interpretation.

Chapter-3 Data

106

Profile 2 (Figure 3.20) is located in the north-western part of the Elan Bank microcontinent (Figure 3.18).

This profile is an almost east-west trending profile, oriented as roughly parallel to the tectonic transport.

The water depth in this profile varies from 3.5 to 6 sec TWT from east to west. The assumed top basalt

layer is the most prominent high amplitude reflective surface, which runs roughly parallel to the seabed.

Thus, the thickness of the overlying sediment cover is quite thin, whose thickness varies from 1 sec TWT

in the east to 0.5 sec TWT in the west. The sedimentary layer is quite transparent due to its low

amplitude content, except for the very top part near the basalt top. Therefore, the internal stratal

termination patterns are difficult to recognize under such circumstances. However, a prominent highly

reflective sediment layer can be identified as based on its amplitude patterns just above the basalt layer.

It has to be noted that the overall amplitude and noise content also increases with depth and data

gradually loses its frequency content. In spite of that the sub-basalt seismic domain can be easily

subdivided into two distinct domains, which are separated in the middle of this profile. These domains

are the eastern and western domains. In the eastern domain, the reflections below basalt are clearly

imaged. One major intra-basalt layer can be identified as this surface is quite reflective due to high

amplitude content and more or less present seismic continuity. Another important traceable reflector

below appears to be discontinuous and affected by faulting. This surface provides a better impedance

contrast between the layers below. Notably, this surface is sealed by another truncation surface above.

The seismic packages between these two surfaces are quite thin, transparent and weakly layered. The

reflection pattern of this interval package is quite different from the basaltic reflection pattern.

Interestingly, the faults that affect the lower surface do not affect the upper truncation surface. They

are essentially detaching at a level below. It is possible to map the detachment surface, which is clearly

rising towards the west. It is possible to map another couple of more or less continuous reflective

amplitudes, which run near parallel to the detachment and rise towards the west. Compared to the

eastern domain, the western domain is full of noisy and chaotic reflections and devoid of any distinct

high amplitude reflectors. Therefore, an overall gradual crustal thinning is observed from east to west

over a distance of 60-80 km.

Chapter-3 Data

107

Fig

ure

3.2

0:

The

Ref

lect

ion

sei

smic

Pro

file

-2 t

hro

ug

h t

he

no

rth

-wes

tern

ma

rgin

of

Ela

n B

an

k. T

he

pro

file

re

pre

sen

ts a

n E

-W t

ren

din

g

dip

-ori

ente

d s

ecti

on

. Th

e d

ata

set

con

tain

a r

ela

tive

ly s

ma

ll a

mo

un

t o

f m

ult

iple

s. N

ote

th

at

the

cru

sta

l th

inn

ing

is

gra

du

al o

cea

nw

ard

.

S2 =

Sed

imen

t U

nit

2,

B1

= to

p B

asa

lt,

IB1

= to

p i

ntr

a-b

asa

lt u

nit

, SR

= t

op

syn

rift

, O

C =

to

p o

cea

nic

cru

st,

UC

C =

to

p u

pp

er c

on

tin

enta

l

cru

st,

MC

C/L

CC

= m

idd

le/l

ow

er c

on

tin

enta

l cr

ust

, LC

C =

lo

wer

co

nti

nen

tal c

rust

, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ails

of

the

ho

rizo

ns

are

giv

en in

Ta

ble

-3.2

. See

tex

t fo

r d

eta

iled

inte

rpre

tati

on

Chapter-3 Data

108

Profile 3 (Figure 3.21) is also located in the north-western part of the Elan Bank microcontinent (Figure

3.18). This profile is, again, an east-west trending profile, oriented roughly parallel to the tectonic

transport. The water depth in this profile varies from 3 to 6.5 sec TWT from east to west. Here, also the

most prominent highly reflective boundary is supposedly being the basalt top. The top basalt layer is

visibly faulted in this profile. It has to be noted that seismic multiples are very much evident in this

profile. The bottom most part is, thus, mimicking the seabed reflections. This creates a difficulty to

identify the continuity of reflectors in the bottom part of the profile. Otherwise, this profile is very

similar to profile 2 (Figure 3.20), in terms of reflection continuity and reflection content in each layer

below the top basalt reflector. The overlying sedimentary layer is quite thin being 1 sec TWT thick in the

eastern part. It thins towards the deeper water region in the west. The bottom most part of the

sedimentary layer again can be recognized as a separate sequence due to its higher amplitude content.

The internal geometry of the sedimentary sequence is quite uncharacteristic and homoclinal in nature.

The upper part of the basalt layer immediately below the high amplitude top basalt reflector is quite

transparent. Its base can be recognized due to its continuity up to the middle of the section from east to

west and change in amplitude, which increases vertically downward. Like in previous profile (Figure

3.20) also a rather discontinuous and faulted reflective surface can be identified in this profile. The

larger number of faults creates multiple small-size rotated fault blocks. The faulted surface is capped by

a thinly layered reflective package sealed by a truncation surface. The faults, again, appear to be

detaching at a traceable common base. There is one a highly discontinuous and discrete reflector below

can also be traced below, which is apparently rising towards the west. Therefore, this profile also shows

very gradual crustal thinning from east to west over a distance of 90-100 km. There is a mound like

feature in the eastern part, before the initiation of a relatively flat-topped basalt and equivalent layer.

The reflections below the mound are highly chaotic.

Profile 4 (Figure 3.22) is located in the south-western part of the Elan Bank microcontinent (Figure 3.18).

This profile is, again, an NE-SW trending profile, oriented at an oblique angle to the tectonic transport.

The water depth in this profile varies from 3 to 6 sec TWT from north-east to south-west. However, this

profile is full of noise, multiples, migration smiles and sideswipes. Certain mound like features can be

also observed here. These mounds are also affecting the current seabed, making the seabed quite

rugged, particularly in the south-western part. The base of the mound appears to be deflected

downwards. Like in other profiles (Figure 3.19, 3.20 and 3.21, the most prominent reflective surface is

assumed to be the basalt top. However, the sediment cover in this profile is almost negligible. The

Chapter-3 Data

109

reflection package immediately below the assumed basalt reflector is quite transparent and free of

amplitudes. However, the package below is full of high-amplitude reflectors. The internal geometry of

this package clearly shows homoclinal south-westerly dipping reflectors. The most prominent high

amplitude reflector is quite continuous up to the middle of the profile. The package is faulted and the

faults presumably detach at a common base, which is not very continuous. The presence of a

discontinuous bottom most reflector, which appears to be rising towards the south-west, can be traced,

but only very locally. Its continuity disappears below the mound feature, located almost in the middle of

the profile. Overall, this profile shows a possible crustal thinning towards the south-west over a distance

of roughly 60-70 km distance.

Chapter-3 Data

110

Figu

re 3

.21

: Th

e R

efl

ecti

on

se

ism

ic P

rofi

le-3

th

rou

gh t

he

no

rth

-we

ster

n m

argi

n o

f th

e El

an B

ank.

Th

e p

rofi

le r

epre

sen

ts a

n E

-W t

ren

din

g

dip

-ori

ente

d s

ecti

on

. Th

e d

atas

et c

on

tain

s a

set

of

mu

ltip

les.

No

te t

hat

th

e cr

ust

al t

hin

nin

g is

gra

du

al o

cean

war

d.

S2

= S

edim

en

t U

nit

2,

B1

= to

p B

asal

t, I

B1

= to

p i

ntr

a-b

asal

t u

nit

, SR

= t

op

syn

rift

, O

C =

to

p o

cean

ic c

rust

, U

CC

= t

op

up

per

co

nti

nen

tal

cru

st,

MC

C/L

CC

=

mid

dle

/lo

wer

co

nti

nen

tal c

rust

, LC

C =

low

er

con

tin

enta

l cru

st, O

M =

to

p o

cean

ic m

antl

e. T

he

det

ails

of

the

ho

rizo

ns

are

give

n in

Tab

le-3

.2.

See

text

fo

r d

etai

led

inte

rpre

tati

on

.

Chapter-3 Data

111

Fig

ure

3.2

2:

The

Ref

lect

ion

sei

smic

Pro

file

-4 t

hro

ug

h t

he

sou

th-w

este

rn m

arg

in o

f El

an

Ba

nk.

Th

e p

rofi

le r

epre

sen

ts a

NW

-SE

tren

din

g d

ip-

ori

ente

d s

ecti

on

. Th

e d

ata

set

con

tain

s a

set

of

mu

ltip

les.

No

te t

ha

t th

e cr

ust

al

thin

nin

g i

s g

rad

ua

l o

cea

nw

ard

.

S2 =

Sed

imen

t U

nit

2,

B1

= to

p

Ba

salt

, IB

1=

top

in

tra

-ba

salt

un

it,

SR =

to

p s

ynri

ft,

OC

= t

op

oce

an

ic c

rust

, U

CC

= t

op

up

per

co

nti

nen

tal

cru

st,

MC

C/L

CC

= m

idd

le/l

ow

er

con

tin

enta

l cru

st,

LCC

= lo

wer

co

nti

nen

tal c

rust

, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ails

of

the

ho

rizo

ns

are

giv

en in

Ta

ble

-2.

See

text

fo

r d

eta

iled

inte

rpre

tati

on

.

Chapter-3 Data

112

Profile 5 (Figure 3.23) is located in the southern part of the Elan Bank (Figure 3.18). This profile is, again,

an NE-SW trending profile, oriented roughly parallel to the tectonic transport. The water depth in this

profile varies from 1.5 to 6 sec TWT from north-east to south-west. The profile also shows a clear

demarcation of shelf, slope, and deep-water parts of the Elan Bank margin. This section is also relatively

free of multiples and other noise. So far, this is the only profile in the Elan Bank, which shows an

exceptionally well imaged crustal thinning towards the south-west, within a deep ocean basin and over a

distance of 80-90 km. This profile also images the entire passive margin from its proximal to distal parts.

Here, again, the most prominent high amplitude reflector is presumed to be the basalt top. The

thickness of sedimentary cover overlying the basalt is very thin in the proximal part over the shelf,

relatively thicker on slope and thickest in deep-water basin. The proximal and distal margins can be

differentiated due to presence of the high-angle breakaway fault. The top basalt reflector drops down

initial 1.5 sec to 6 sec TWT. The base of the basalt can be also recognized as a continuous high amplitude

reflective surface in the proximal shelf region, where it is 1 sec TWT below the top basalt reflector, and

runs almost sub parallel to the base-basalt reflector. This indicates a quite uniform thickness of basalt

layer. The basalt package in the proximal part in the north-east appears to be quite transparent due to

low amplitude content. However, in the distal part in the south west, the basalt package is layered with

a different degree of reflectivity. The basalt layer is also typically faulted by high angle normal faults. The

stratal pattern below basalt in the proximal margin is not clear but, in the distal margin some smaller

fault bounded sedimentary wedges can be recognized in the north-eastern part. A typical bottom most

discontinuous reflector can be recognized only in the north-eastern part. It appears to be rising towards

the south-west. The top basement in the deep-water region can be recognized due to its high

impedance contrast. The basement appears to be quite horizontal. However, there is a mound like

feature in the extreme south-western part, which is faulted on its flank as well.

Profile 6 (Figure 3.24) is located in the central part of the Elan Bank (Figure 3.18). This is a roughly east-

west trending profile, roughly parallel to the tectonic transport. The water depth in this profile varies

from 1.25 to 2.5 sec TWT from west to east. This profile is free of multiples but the amplitude and noise

content increases rapidly along the vertical axis. Here, also the most prominent highly reflective

boundary is supposedly being the basalt top. The sedimentary section above the basalt layer is quite

thin (less than 1 sec TWT) and it thickens from west to east. The intra-layer stratal amplitude pattern

within the sedimentary package is difficult to distinguish. The homoclinal layering of the alternate high

and low amplitude pattern in observed from west to east in this profile. The upper part of the basalt

Chapter-3 Data

113

layer immediately below the high amplitude top basalt reflector is relatively transparent. Its base can be

recognized due to its continuity throughout the section from west to east. The seismic reflection below

the transparent package is quite reflective and continuously layered. A discrete and discontinuous

reflector can be observed below the transparent basalt package, which appears to be highly faulted. This

reflector gradually steps down along faults. There are some rotated seismic wedges in the profile, which

may be sedimentary in nature and may represent rifted grabens. However, this is indistinguishable in

seismic image due to high noise content in the data. However, the top part of this package is capped by

a clear truncation surface, which can be traced throughout the profile from west to east. The

sedimentary nature of sub-basalt reflections can be only confirmed by presence of IODP well 1137,

which is present on this profile (Figure 3.24). Typical continental fragments were recovered from the

clasts of the fluvial facies. The details of the findings from this well will be described later.

Profile 7 (Figure 3.25) is a north-east to south-west trending profile, at an angle to the major tectonic

transport. This is the longest and continuous profile of the entire survey (S-179), connecting the Elan

Bank, the Kerguelen Plateau and the Labuan Basin (Borissova et al., 2002) (Figure 3.18). The water depth

in this profile varies from 3 to 6 sec TWT. The profile shows a very rugged and uneven sea bottom along

its entire length. It contains some multiples. The shallowest part of the profile is located in the central

part, which represents the southern Kerguelen Plateau. The south-western part of the profile is partly

passing through the Elan Bank, while the Enderby Basin is located towards the very end of the profile.

On the north-eastern side, a clear image of the sedimentary basin can be found. It is known as the

Labuan basin (Rotstein et al., 2001). It appears that the sedimentary section in the Labuan Basin is

tectonically deformed. The thickening of sediment is observed in this basin toward the faults. The most

prominent high amplitude reflector is, again, presumably the basalt top, which can be continuously

traced throughout the profile across the southern Kerguelen Plateau. Apart from the Labuan Basin, the

sedimentary section overlying the basalt layer uniformly thick over Kerguelen Plateau. Its thickness is

approximately 1 sec TWT. The internal geometry of the sediments is uncharacteristic, except the

alternate layering of high and low amplitude reflectors. The data below the top basalt are highly chaotic

and noisy. Therefore, it is almost impossible to recognize continuous reflection patterns and internal

stratal amplitude geometries below the top basalt layer due to high noise content and presence of

multiples. There are certain high-angle discontinuities in the north-eastern part of the profile. The top

basalt surface is faulted, particularly in the north-western part of the profile. The crust appears to be

more deformed in the north-eastern part in comparison to the south-western part.

Chapter-3 Data

114

Fig

ure

3.2

3:

The

Ref

lect

ion

sei

smic

Pro

file

-5 t

hro

ug

h t

he

sou

ther

n m

arg

in o

f El

an

Ba

nk.

Th

e p

rofi

le r

epre

sen

ts a

NE

- SW

tre

nd

ing

dip

-

ori

ente

d s

ecti

on

. Th

e d

ata

set

con

tain

s a

rel

ati

vely

sm

all

am

ou

nt

of

mu

ltip

les.

No

te t

ha

t a

ver

y sh

arp

tra

nsi

tio

n b

etw

een

pro

xim

al a

nd

dis

tal

ma

rgin

wit

h m

ajo

r se

t o

f b

rea

kaw

ay

fau

lts.

S

2 =

Sed

imen

t U

nit

2,

B1

= t

op

Ba

salt

, IB

1 =

to

p i

ntr

a-b

asa

lt u

nit

, SR

= t

op

syn

rift

, O

C =

to

p

oce

an

ic c

rust

, UC

C =

to

p u

pp

er c

on

tin

enta

l cru

st, M

CC

/LC

C =

mid

dle

/lo

wer

co

nti

nen

tal c

rust

, LC

C =

low

er c

on

tin

enta

l cru

st, O

M =

to

p o

cea

nic

ma

ntl

e. T

he

det

ails

of

the

ho

rizo

ns

are

giv

en in

Ta

ble

-2. S

ee t

ext

for

the

det

aile

d in

terp

reta

tio

n

Chapter-3 Data

115

Figure 3.24: The Reflection seismic Profile-6 through the central part of Elan Bank. The profile represents an E-W trending dip oriented section. ODP site 1137 is located on this profile. S2 =Sediment Unit 2, B1 = top Basalt, IB1 = top intra-basalt unit, SR = top synrift, OC = top oceanic crust, UCC = top upper continental crust, MCC/LCC = middle/lower continental crust, LCC = lower continental crust, OM = top oceanic mantle. The details of the horizons are given in Table-2. See text for detailed interpretation.

Chapter-3 Data

116

Fig

ure

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The seismic sequence boundaries and have been given in Table 3.2. The crustal boundaries interpreted

in the Elan Bank:

I. Top upper continental crust (UCC);

II. Top middle/lower continental crust (MCC/LCC);

III. Continental Moho (CM);

IV. Top oceanic crust (OC); and

VI. Oceanic Moho (OM).

Table 3.2: Tectono-stratigraphic chart for Elan Bank used for identification of seismic sequence boundaries.

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3.2. Potential field data

Generally, marine gravity anomaly data can be estimated from satellite radar altimetry. They are

sometimes referred to as satellite gravity. In this study, the satellite gravity data come from Sandwell

and Smith, (1997). This gravity dataset was compiled from sadG data, version 11.2, a 1-arc-minute grid

compilation, and satellite altimetry mission data sets. These data are accurate down to wavelengths of

20-25 km. They are available at http://topex.ucsd.edu/WWW_html/mar_grav.html. Additionally, KMS01

and KMS02 data are also available for the study. The KMS01 mean sea surface dataset was derived from

a combination of T/P, ERS (ERM and GM modes) data. The resolution of the marine gravity field is 2 by 2

min (1/30 degrees) globally. The global marine gravity field and mean sea surface is available at the

following website: ftp://ftp.spacecenter.dk/pub/GRAVITY/KMS02/.

The ship-track gravity data are available from ION India span series, which were acquired simultaneously

with seismic data acquisition. Thus, the gravity data acquired are present only along the ION-GXT

seismic profiles.

The magnetic dataset was delivered as profile data along marine track lines and was unleveled. These

data were leveled and extrapolated away from the track lines for qualitative analyses of regional

magnetic trends. Unfortunately, a ship track magnetic coverage from ION is incomplete due to some

security and territorial issues. However, additional magnetic coverage was obtained from Bird

Geophysical who merged the IndiaSpan magnetic data with reprocessed GeoDas coverage. The merged

data are available from Bird Geophysical at http://www.birdgeo.com.

3.2.1 Gravity Maps

The Free-air gravity map of the East India margin shows a distribution of anomalies in the offshore

region (Figure 3.26). If the Free-air gravity anomaly map is compared to the bathymetry map (Figure

3.14), it shows that the positive anomalies are restricted to the shelf region only. The Free-air anomaly

map reveals that anomalies increase up to +112 mGal from ocean towards shelf. The positive anomaly in

the shelf region abruptly decreases to negative anomaly on the ocean side. It decreases to a minimum of

-197 mGal over a short distance across the slope region. The anomaly then again increases at a slower

rate towards the deep-water in the east, but remains negative (mGal). The map also contains several

slightly positive anomalies in areas where there are major delta and deep-water fans, with expected

thick sedimentary cover in the areas of the Bengal and Godavari fans. In the Free-air anomaly map, a

linear negative gravity anomaly trend, flanked by positive anomaly, is observed in the deeper part of the

http://topex.ucsd.edu/WWW_html/mar_grav.html

ftp://ftp.spacecenter.dk/pub/GRAVITY/KMS02/

http://www.birdgeo.com/

Chapter-3 Data

119

Bay of Bengal basin. This trend coincides almost with 85°E longitude. This particular anomaly is known as

85°E Ridge anomaly. It represents a chain of buried aseismic volcanic mounds known as the 85°E Ridge

(Curray and Munasinghe 1991; Muller et. al, 1993; Ramanna, et al., 2001; Bastia et al., 2010). The ocean

floor anomaly becomes again positive to the east, towards the eastern side near the Andaman Sea.

Figure 3.26: Free-Air gravity map of the East Indian offshore. The anomaly map along the margin shows a distribution of anomalies representing the crustal transition. Note that the linear negative gravity anomaly trend runs almost north-south in the deeper part of the Bay of Bengal Basin. It coincides with almost along 85°E longitude. This particular anomaly is known as the 85°E Ridge anomaly.

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The Elan Bank represents a relatively positive topographic feature in comparison to the surrounding

oceanic basins as shown in the Figure (3.27). In this map, the shallowest part is located in the central

and northern parts of the Kerguelen Plateau where the bathymetry increases up to 900 m below Mean

Sea Level (MSL) and the deepest part is located in the ocean basins in the west, where the bathymetry

decreases up to 6000 m below MSL . In Elan Bank, the bathymetry is ranging roughly between 1400 m to

4000 m below MSL. The typical shelf-slope break in Elan Bank can be only observed along northern and

southern margin. The western margin represents a gradual bathymetric change from the central part of

Elan Bank to the deep ocean basin. In this map only the Heard-McDonald Island in the central Kerguelen

Plateau is only a positive topographic feature. It has to be noted that there are some linear but random

topographic features surrounding the Elan Bank.

Figure 3.27: Topography of the Elan Bank and surrounding Kerguelen Plateau. Note that a positive elevation is associated with the Heard-McDonald Island (Data source: http://topex.ucsd.edu).

Free-air gravity map of Elan Bank shows a positive gravity anomaly over Elan Bank and Kerguelen

Plateau in comparison to the adjacent deep ocean basins (Figure 3.28). The positive anomaly increases

up to +96 mGal in the interior of Elan Bank. The negative anomaly on the ocean side decreases to a

minimum of -79 mGal in the Enderby Basin. However, the change is not uniform along margins. In the

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121

southern and northern margin, the maximum positive value changes to negative (-79 mGal) in the

deeper ocean basin, almost instantly from shelf to deeper basin. This transition is also observed in the

western margin also but a positive gravity anomaly values changes to less a negative gravity anomaly

value compared to the southern margin. A trend of discontinuous linear positive gravity anomaly

roughly oriented northeast to southwest is observed in the southern margin of Elan Bank. It has to be

noted that the feature also exists along the same trend in the bathymetry map (Figure 3.27). In the

central part of Elan Bank, one can clearly notice that there is a right lateral shift in the positive anomaly

feature along the roughly northwest-southeast trend. Overall, the Kerguelen Plateau is associated with a

positive anomaly. However, few linear trends of negative anomalies are visible, particularly in the

southern Kerguelen Plateau. The Labuan Basin, which is located to the south-east of the Kerguelen

Plateau, is showing a linear negative anomaly trend roughly oriented from northwest to southeast.

Figure 3.28: Free air gravity map of the Elan Bank. The anomaly distribution shows a positive gravity anomaly over the Elan Bank and the Kerguelen Plateau in comparison with anomalies in the adjacent deep ocean basins (Data source: http://topex.ucsd.edu). Note the positive anomaly associated with Elan Bank.

http://topex.ucsd.edu/

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122

Bouguer gravity map represents a gravity response of the bathymetry removed; thus leaving only the

gravity response of the underlying geology. The Bouguer gravity derivative map of 10km high pass in

East India (Figure 3.29) is enhancing the short wavelength regional anomalies. This map, if compared to

the bathymetry in East India (Figure 3.14), it shows that the shelfal part is mostly represented by

positive anomaly coupled with a long linear parallel trend of low-wavelength negative anomaly.

However, it can be observed that these positive anomalies are not continuously parallel to the coast as

they are intersected by intermittent breaks. Further outboard, this negative anomaly is again coupled

with parallel long-wavelength broad positive anomaly. The map also helps to identify the 85oE Ridge by

subdued, broad, elongated N-S trending negative anomaly coupled with two external thin, elongated,

subdued positive anomalies. The 85°E Ridge anomaly is also very prominent in the Bouguer gravity data.

There are two circular negative anomalies in offshore Krishna-Godavari and Cauvery basins. The deep-

water region east and west of 85°E Ridge is represented by broad long-wavelength anomaly. Another

interesting observation in offshore the Cauvery Basin indicates that there is a circular low-wavelength

positive anomaly sandwiched between the two short-wavelength NE-SW and N-S trending linear

anomalies.

The Bouguer gravity maps of Elan Bank shows a distribution of anomalies from central part to the deep

basin part of Elan Bank (Figure 3.30). This map shows features and trends in anomaly pattern similar to

those observed in the Free-air gravity map (Figure 3.29). However, the features and their edges have

been enhanced. The Elan Bank mostly represents a positive anomaly, so does the most of the Kerguelen

Plateau. In comparison to the Free-air gravity, it can be seen that the change in gravity gradient at

southern margin in Bouguer gravity is much steeper in comparison to the northern margin or western

margin of Elan Bank. The linear short-wavelength, discontinuous anomalies, oriented roughly in the

northeast-southwest direction, are also present in the Enderby Basin located to the south and south-

west of the Elan Bank. Although, most of the features in Bouguer gravity map are enhanced, the right

lateral shift in positive anomaly feature along the roughly northwest-southeast trend in the central part

of Elan Bank is more subdued. Nevertheless, this feature is still recognizable in the Bouguer gravity map.

The isostatic residual gravity map of the East Indian offshore (Figure 3.31) allows a bit finer distinction of

imaged features. The shelf and deep-water region have been characterized by coast-parallel positive

anomaly. The two circular anomalies observed in the Bouguer gravity map are further enhanced in this

map, which shows them as negative anomalies. The isostatic residual anomaly map also shows a long,

linear anomaly associated with 85°E Ridge. It has to be noted that while the west of the 85°E Ridge is

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123

showing an overall positive anomaly pattern in general, except two circular negative anomalies

mentioned above, the eastern region of the 85°E Ridge anomaly is characterized by mostly negative

anomalies. This image contains several data gaps, though.

Figure 3.29: 10 km High pass Bouguer gravity map of East India. The anomalies of ocean-continent boundary and 85°E Ridge anomaly are also prominent.

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124

Figure 3.30: Bouguer gravity map of Elan Bank. The anomaly of the Elan Bank and the Kerguelen Plateau is compared to the adjacent deep ocean basins (Data source: http://topex.ucsd.edu). Note positive anomalies associated with Elan Bank and Kerguelen Plateau.

3.2.2 Gravity Profiles

Apart from the maps, the gravity profiles extracted along the seismic profiles are also used in this study.

In East India, high-resolution gravity and magnetic data were collected in tandem with ION Geophysical

IndiaSpan dataset. In Elan Bank, the profiles have been extracted from the prepared gravity-grids along

the seismic profiles of Survey-179 (Figure 3.18).

The Bouguer gravity profiles used in this study for East India have been shown in Figure (3.32). The key

ION profiles are located in the Krishna-Godavari (IE-1000, 1200, and 1290) and the Mahanadi basins (IE-

1600) (Figure 3.14). All gravity and magnetic profiles along these seismic transects show several

following characteristic features. The most common observation is the gravity value decreasing steeply

over a short distance in the north-west, followed by a small bump, and then increasing towards the

south-east. If these profiles are compared with bathymetry (Figure 3.14), then the sharp decrease

coincides with the shelf break.

http://topex.ucsd.edu/

Chapter-3 Data

125

Figure 3.31: Isostatic residual anomaly map with 3*3 km resolution derived from interpolated ship-track gravity data (Fugro Robertson Inc., 2006). Lack of ship-track coverage in some areas is indicated as data gaps. The positive anomaly on continental side coupled with negative anomaly on oceanic side represents a good candidate for identifying the crustal boundary. The 85

o E Ridge is imaged as a prominent gravity low representing crustal roots

underneath a chain of volcanic loads. Two isolated circular anomalies of similar origin and located to the W of the ridge in the Krishna-Godavari and Cauvery basins.

In the profile IE-1000 (Figure 3.32a) the gravity value decreases from initial 0 mGal in shelf to -80 mGal

on the slope. The gravity value then increases to -65 mGal, and then again decreases to -80 mGal.

Therefore, the gravity value shows a characteristic bump in profile IE-1000. This bump is a short-

Chapter-3 Data

126

wavelength high-amplitude anomaly. In the south-east, the gravity values increase to -25 mGal in a

steady fashion. This rise is a part of a very large wavelength mega-regional anomaly.

The profile IE-1200 (Figure 3.32b) also shows a similar trend in gravity curve. However, in this profile,

the initial 0 mGal anomaly value in the shelf decreases to -75 mGal, then increases to -40 mGal before

coming back to -75 mGal. The bump in this profile, therefore, has a shorter wavelength but higher

amplitude in comparison to that in the previous profile. In the south-east, the gravity profile increases

again to -25 mGal but again a decreasing trend is observed, along which the anomaly value decreases to

-75 mGal. This indicates that a short wavelength regional anomaly may exist further south-east within

the long-wavelength mega-regional anomaly.

The profile IE-1290 (Figure 3.32c) shows a little bit different trend. Here the initial 40 mGal value

decreases to -75 mGal very sharply along the slope. The value then increases to -60 mGal and drops

down to -65 mGal before steadily increasing to -10 mGal. Thus, the bump in this profile is very subtle

and short-wavelength anomaly. It has to be noted that at the south-eastern end, a short wavelength

anomaly within the mega-regional long-wavelength anomaly exists across which the value decreases

near to -35 mGal from -10 mGal.

The gravity anomaly value along profile IE-1600 (Figure 3.32d) shows some kind of dissimilarity from the

previous ones. Here, the gravity anomaly decrease along the slope in north-west is relatively gradual.

The value decreases from initial 35 mGal to more than -60mGal. After that, it increases to -30 mGal,

before dropping down to 50 mGal. The anomaly curve characteristically becomes almost horizontal at -

45 mGal anomaly value in the south-east. This indicates a very large wavelength anomaly in the south-

east.

Various characteristic features can be observed in the Elan Bank gravity profiles extracted along the

seismic lines (Figure 3.33). The Free-air gravity values are quite different in each profile. The gravity

value along profile S179-01 (Figure 3.33b), located in the central Kerguelen Plateau (Figure 3.33a) shows

no major anomaly variation along the profile indicating a presence of a very large wavelength anomaly.

The profile S179-02, located in the north-western margin of the Elan Bank (Figure 3.33a), however,

shows a characteristic anomaly variation along the profile. The maximum gravity anomaly value of 45

mGal in the east decreases to near 7-8 mGal within a distance of 80 km. Now, if compared to the

bathymetry map of Elan Bank (Figure 3.27) and seismic profile (Figure 3.20), this profile does not show

any steep change in bathymetry characteristic of shelf to slope transition. Therefore, this indicates a

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127

high amplitude moderate wavelength anomaly. The value increases further east up to 12 mGal before

steadily decreasing up to 0 mGal. Therefore, a small characteristic short-wavelength low-amplitude

bump is observed in this profile.

Figure 3.32: Bouguer gravity and magnetic anomaly curves along ION IndiaSpan profiles. The curves are generated from ship-track gravity data acquired during ION-IndiaSpan acquisition in East India. a. Profile IE-1000 in the Krishna-Godavari rift zone, b. profile IE-1200 in the Krishna-Godavari rift zone, c. Profile IE-1290 in the Krishna-

Godavari rift zone, d. Profile IE-1600 in the Mahanadi rift zone.

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128

Figure 3.33: Calculated Bouguer gravity curves along S179 survey profiles in the Elan Bank and the southern Kerguelen Plateau. a. Location map of survey profiles, b. S179-01 located in the central Kerguelen plateau c. S179-02, located in the north-western margin of Elan Bank, C) S179-03 located at the north-western margin of Elan Bank d. S179-04 located at the south-western margin of Elan Bank, e. S179-05 located at the southern margin f. S179-06 located at the central part of Elan Bank and g. S179-07 located from Enderby Basin to the Labuan basin across the southern Kerguelen Plateau.

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129

The profile S179-03 is also located in the north-western margin of the Elan Bank (Figure 3.33a). It shows

a different anomaly distribution (Figure 3.33c). Here two short wavelength anomalies are observed from

east to west. The initial anomaly value of 35 mGal decreases to 20 mGal then becomes flat for a while

before dropping down to near 12 mGal. This creates a stair-like pattern in gravity value, which

represents a local anomaly within a long-wavelength regional anomaly. The anomaly value is further

increases to 17 mGal towards the west, before decreasing to 0 mGal. Thus, it again, features a small

bump representing a short-wavelength anomaly in the profile. At the western end, the anomaly value

steadily increases to 5 mGal. This increasing trend indicates very long-wavelength anomaly feature

towards west.

The fourth profile S-179-04 is located at south-western margin of the Elan Bank (Figure 3.33a). It shows,

again, a characteristic gravity anomaly (Figure 3.33e). The gravity value decreases from initial 45 mGal to

almost 10 mGal from north-east to south-west. Then the profile shows two consecutive gravity rises,

where the value increases to 38 mGal then decreases down to 20 mGal and again increases to 33 mGal.

These are very short-wavelength anomalies. The gravity anomaly value finally decreases down to 5 mGal

in the south-west.

The profile S179-05 is located at the southern margin of the Elan Bank (Figure 3.33a). If compared with

the bathymetry of Elan Bank (Figure 3.27) and seismic profile (Figure 3.23), this profile represents typical

shelf slope geometry of passive margin (Figure 3.33f). It can be observed that the gravity anomaly value

decreases from the initial 70 mGal in the shelf break region to -50 mGal along the slope in this profile in

the north-east. The gravity anomaly values from there on are on a steadily rising trend, climbing to near

20 mGal on the south-eastern end. This south-eastern part, therefore, represents a very long-

wavelength regional anomaly. It has to be noted that this curve is quite simple, with no short-

wavelength bumps like those in the previous profiles through the Elan Bank. The profile S179-06 is

located in the central part of the Elan Bank (Figure 3.33a). This profile does not show much variation in

anomaly values (Figure 3.33g), which are ranging between a minimum of 50 mGal to maximum of 70

mGal. This anomaly pattern represents part of a very long-wavelength regional anomaly.

The final profile S179-07 runs across the southern Kerguelen Plateau from the Enderby basin in the

south-west to the Labuan basin in the north-east, encompassing 600 km length (Figure 3.33a). It covers

a small part of the Elan Bank. The overall gravity trend shows a very large wavelength regional anomaly

rising towards the north-east. However, there are numerous short-wavelength high-amplitude local

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130

anomalies present within the large wavelength trend. Four such anomalies can be easily identified. The

values vary from -20 mGal to 25 mGal and then come down to -5 mGal in the first one from south-west.

In the next one, the value climbs to 25 mGal and drops down to 15 mGal. In the next one, it again climbs

to 45mGal and decreases down to 10 mGal. The last one in the north-east is quite symmetrical, where

the value climbs to 50 mGal from 10 mGal and then comes back to 10 mGal.

3.2.2 Magnetic Profiles

Magnetic profiles are only available in East India along few of the ION Geophysical IndiaSpan-East

profiles in Krishna-Godavari rift zone (IE-1000, 1200 and 1290) and the Mahanadi rift zone (IE-1600)

(Figure 3.14). All profiles show mostly long-wavelength reverse magnetic anomalies. The profile iE-1000

shows three reverse short-wavelength magnetic anomalies from north-west to south-east (Figure

3.32a). Along this profile, the reverse magnetized values go down to -80 nT, more than -160 nT and -160

nT in three such events from north-west to south-east. There are two short-wavelength reverse-

magnetized anomalies present at both ends of the profile (Figure 3.32b), where the reverse-magnetized

values are -112 nT and more than -240nT, respectively, from north-west to south-east. In the next

profile IE-1290, again, two reverse-magnetized anomalies are present (Figure 3.32c). Both the short-

wavelength anomalies go to -100 nT. The final profile in the Mahanadi Basin (IE-1600) is incomplete due

to no data acquisition in the south-eastern end. However, it also shows two reversely-magnetized short-

wavelength anomalies, where the reverse magnetization reaches to -340 nT in the north-western one

and -285 nT in the south-eastern one.

3.3 Borehole data

Data from a series of boreholes, particularly biostratigraphic information have been used to constrain

the seismic interpretation in East India. The wells in East India are proprietary data and, therefore,

confidential. As a result of that, the names and coordinates of the wells are not provided in this thesis.

The wells have been selected in a systematic matter. They are geographically distributed over the study

area and penetrate maximum stratigraphic depths. However, the geographic well distribution is also

skewed due to various degrees of exploration activities in the area. Out of the 15 selected key wells in

East India, eight wells are present in the northern Krishna-Godavari rift zone, three wells are present in

the northern Coromandal horsetail structure, one well in the central Coromandal zone and two wells in

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southern horsetail structure of the Coromandal strike-slip fault system and Cauvery rift zone and one

well in Mahanadi rift zone. The data including biostratigraphy and paleo-environment section from each

well are also proprietary. Therefore, details of the data could not be documented in this thesis. A

summary chart of the observations from these wells is given in Figure 3.34.

The paleo-environment data indicate a change in bathymetry in each well from continental through

shallow marine to deep marine environments. They also indicate the well (B1) in the Mahanadi rift zone

is terminated within the basalt only and it only exhibits a deep water environment. Out of 15 wells in

total, only four well reach the basement. There are three wells (T1, T3 and T4) located in the northern

Krishna-Godavari rift zone and one (P4) in the Cauvery rift zone. The oldest sediment are drilled in well

P-1 and P2 in the northern Krishna-Godavari rift zone, having the Early Jurassic Toarcian (182 Ma) age.

The thickest continental sediments are drilled in well P2, having penetration of 1650 m in total. This

sediment was deposited in a continental environment. The age of continental sediments shows a

deposition since Toarcian (182 Ma) in Early Jurassic to Hauterivian (129 Ma) in Early Cretaceous, which

encompasses an almost 53 Ma long time span. These deposits are present in all wells except the one in

the Mahanadi basin.

Three wells (P1, P2 and D1) in the northern Krishna-Godavari rift zone show shallow-water sediments of

200m, 300m and 50m thickness respectively. The 800 m thick shallow water deposits are also observed

in one well (P3) at the northern end of the Coromandal horsetail structure. This represents a maximum

thickness of shallow water deposits ranging from Hauterivian (129 Ma) to middle Aptian (122 Ma)

present in any well in the East India. It has to be also noted that the wells (P3 and D1) are the only wells,

which show a continuous sedimentary succession, without any detectable erosion or non-deposition.

The shallow-water deposits are also observed in the Cauvery rift zone wells (P4 and T10) but the

thickness is considerably less (50m and 140m, respectively).

The erosion or non-deposition has been documented all wells. This can be recognized by presence of

missing stratigraphic sections. Stratigraphic gaps are variable in all wells used in this study. The

maximum missing section is observed in well T3 in the northern Krishna-Godavari rift zone ranging from

Bajocian (168 Ma) to middle Aptian (123 Ma). Another interesting observation can be made in two wells

(P1 and P2) in the northern Krishna-Godavari rift zone and two wells (P4 and T10) in the Cauvery rift

zone. In these four wells missing stratigraphic sections are present between continental to shallow-

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water transition. However, in the two wells (P4 and T10) in the Cauvery rift zone, missing stratigraphic

section comes from both above and below the shallow-water deposits.

The deep-water sections are more or less continuous from middle Aptian (122 Ma) onwards in all wells

except the two wells (P1 and P2) in the northern Krishna-Godavari rift zone. In these two wells, the

deep-water deposits start from middle Hauterivian (130 Ma). The well (T9) located in the central

Coromandal horsetail structure also exhibits a late deep-water transition from Albian-Aptian onwards

(113 Ma).

There are two borehole datasets, which could be actively used for the Elan Bank and Kerguelen plateau.

These wells have been drilled during the ODP leg 183 (Coffin et al., 2000). As mentioned earlier, the site

1138 is located within the main Southern Kerguelen Plateau, on seismic profile-1 of survey S-179. This

well penetrates a depth of 850 m. The complete stratigraphic column of well is given in Figure 3.35). The

well ended within the basalt sequence after penetrating 150 m of basalts. The basalt section is overlain

by a thin Upper Cretaceous silty claystone sequence topped by a glauconite bearing sandstone. The

section from 500 m to 600 m is dominated by a chalk deposit. The overlying section between 150 m to

500 m is again, dominated by nanofossil ooze, whereas the top section is dominated by diatom and

foraminifera ooze.

The Site 1137 is located on the Elan Bank, at a water depth of 1004 m. This well penetrated 219 m of

marine sediment and 151 m of basement (Coffin et al., 2000). This well also ties to seismic profile-6 of

survey S-179. This well reaches at a depth of more than 350 m. The complete stratigraphic column of

the well is given in Figure 3.36. The top 200 m of the well contain dominant nanofossil ooze. There is

some amount of clayey sand present before the well penetrates 65 m of massive basalt. The next 30 m

long section is a volcanoclastic conglomerate followed by 20 m thick massive basalt sequence. The

bottom most section is represented by brecciated massive basalt and conglomerate with locally sourced

fluvial deposits containing clasts of garnet-biotite gneiss recovered at Site 1137 (Nicolaysen et al., 2001)

Comprehensive ODP reports for these sites are publicly available (Coffin et al., 2000).

Chapter-3 Data

133

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Figure 3.35: Composite stratigraphic section of ODP Site 1138, which is located on the Kerguelen Plateau and tied to seismic Profile-1 (Coffin et al., 2000).

Chapter-3 Data

135

Figure 3.36: Composite stratigraphic section of ODP Site 1137, which is located on the Elan Bank and tied to seismic Profile-6 (Coffin et al., 2000).

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3.4. Petrological data from literature

A comprehensive work has been carried out by previous workers detailing the petrological similarities

between Eastern Ghats basement samples and samples found in the ODP site 1137 at Elan Bank. The

correlation was established using different parameters from petrological information. These parameters

include petrography, geochemistry, trace element geochemistry, radiometric dating (Clarke, 1988;

Gopalakrishnan, 1998; Mezger and Cosca, 1999; Coffin et al., 2000; Harley and Motoyoshi, 2000; Krause

et al., 2001; Nicolaysen et al., 2001; Rickers et al., 2001; Frey et al.; 2002; Ingle et al., 2002; Chetty et al.;

2003; Ghatak and Basu, 2011). The summary is described below.

Basement units in ODP site 1137 (Figure 3.37), including units 1–4, 7, 8, and 10, consist of basalt with

brecciated and oxidized flow tops. Unit 5 (286–291 m) consists of interbedded sandstone and siltstone

layers containing volcanic lithic fragments as well as minor quartz, feldspar, hornblende and biotite.

Heavy minerals, like garnet and zircon, are abundant in these sandstone layers. Unit 6 is a conglomerate

consisting of poorly sorted clasts, ranging from granule to boulder size in a polylithologic matrix of

coarse-sand grain size. Basalt is the most abundant clast type, but there are also clasts of flow-banded

rhyolite, sanidine-phyric trachyte, garnet-biotite gneiss, actinolite gneiss, and granitoids. Unit 9 is a

sanidine-rich, crystal vitric tuff that contains rare seimi-rounded lithic clasts of the same lithologies

found in the unit 6 conglomerates (Ingle et al., 2002). Isotopic analysis data allow comparing of these

granite clasts of cratonic parts with those from the Eastern Ghat and India (Mezger and Cosca, 1999).

The quartz-feldspathic gneiss clasts from fluvial samples contain quartz, alkali feldspar, plagioclase, and

garnet, and trace amounts of zircon and monazite (Nicolaysen et al., 2001). Biotite and muscovite are

present in some clasts. Zircons from the gneiss clast fall into two groups on the basis of shape, size, and

clarity. The first group consists of a few large (200 mm long) angular, brownish fragments, whereas the

second group is characterized by clear, elongate to equal subrounded grains, which are 200 mm long.

Detail electron microscope analysis reveals distinct cores with embayed margins that are surrounded by

one or two overgrowth rims. The internal structures of the monazite grains are similarly complex. They

have irregular cores with one or two rims (Nicolaysen et al., 2001).

Individual zircons were abraded prior to dissolution. Three of the large brown zircon fragments yielded

Pb 207/ Pb 206 dates of 795.8 ± 1, 836.1 ± 2.1, and 938.5 ± 3.2 Ma, respectively (Figure 3.38). Four biotite

grains from the second garnet-biotite gneiss clast were analyzed by the Ar40/Ar39 single-crystal step-

heating techniques. The age spectra of the 500–600-mm-diameter biotites suggest that the high-

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temperature steps record a minimum peak-metamorphic age of ca 550 Ma (Nicolaysen et al., 2001). The

radiogenic isotopic ratios for selected sample from ODP-1137 are tabulated in Figure 3.39.

Figure 3.37: The fluvial samples from ODP -1137 site confirming the presence of continental crust at Elan Bank (modified after Coffin et al., 2000 and Ingle et al., 2002). Clasts from the fluvial conglomerate to pebble have variable lithologies including alkali basalt, rhyolite, granite-gneiss and other metamorphic rocks. The gneisses, sandstones and conglomerate matrix contain zircon and monazites with Neoproterozoic and Archean ages. This suggests the proximity of continental crust during deposition of fluvial conglomerates.

The Eastern Ghat is characterized by charnockitic and khondalitic group of rocks, which are extensively

invaded by coarse-grained megacrystic gneisses (Figure 3.40a). Megacrystic gneisses include both

hypersthene- and biotite-bearing types (Chetty, et al, 2003). In general, the margin of shear zones is

enriched by the hypersthene-bearing gneisses. Similarly, the biotite-bearing ones are essentially located

within central part and comprise enclaves of the country rocks (Figure 3.40b). Leucogneisses are the

other predominant rock type in the southeastern part of the area. They are mostly restricted to the axial

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parts of shear zones. They display a continuous compositional banding of biotite-rich and biotite-poor

quartz-feldspathic layers. Leucogneisses are medium- to fine-grained and display an alternation of mafic

and felsic layers. The layering is formed by the transposition and alternation of pre-existing gneissic

banding, and/or mafic dykes.

Figure 3.38: Radiometric age dating and comparison of continental basement origin of India and Elan Bank (Nicolaysen et al., 2001). a. Concordia plot shows a slightly discordant array of zircon and monazite data from unit 5 sandstone, unit 6 conglomerate matrix, and garnet-biotite gneiss (sample 183- 1137A-34R2, 118–120 cm). b. 40

Ar/39

Ar furnace step-heating spectra for four individual biotites from second garnet-biotite gneiss clast (sample 183-1137A-35R2, 42–47 cm). Spectra show a diffusive loss of argon, but high temperature steps suggest a minimum peak-metamorphic age about ca. 550 Ma.

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Under the microscope, one can see that the coarse-grained charnockite is constituted of plagioclase,

perthite, quartz, garnet, clinopyroxene, orthopyroxene, apatite and rounded zircons in decreasing order

of abundance. Foliation is defined by the alternation of layers of mafic and felsic minerals. The

hypersthene-bearing megacrystic gneisses are generally coarse-grained and consist of minerals such as

plagioclase, perthite, quartz, garnet, orthopyroxene, biotite, hornblende, iron oxide, apatite and

rounded zircons. Foliation is well defined by the alignment of biotite and hornblende, which are

abundant as retrograded products from pyroxene.

Figure 3.39: Radiogenic isotopes of selected samples from ODP site 1137 at Elan Bank showing both composition and provenance of rock clasts from the recovered fluvial conglomerate sample (Ingle et al., 2002)

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Figure 3.40: Field photographs of different lithounits of the Eastern Ghat mobile belt from the Nagavalli shear zones, supposedly adjacent to the Elan Bank (Chetty et al., 2003). a. Coarse-grained megacrystic gneisses with enclaves of calc-granulites b. Sectional view of mesoscopic shear zone at the transition between hypersthene bearing megacrystic gneiss and biotite-bearing gneiss.

141

Chapter 4

Methods

This study utilizes different kinds of scientific methods and approaches. The following methods have

been systematically approached to envisage the nature of various technical aspects of the study;

a. Seismic interpretation

b. Potential field data interpretation

c. Gathering petrological data on Elan Bank/Antarctica/E India for geological correlation

d. Tectonic reconstruction between Elan Bank and East India

e. Seismic marriages

f. Tectonic timing

g. Match of Elan Bank with hot-spot track in Bay of Bengal

4.1 Seismic Interpretation

In the current study, the seismic interpretation was carried out over a vast dataset in East India and Elan

bank, where two major types of interpretation were carried out. These are:

a) Seismic and sequence stratigraphic interpretation to recognize tectono-stratigraphic elements,

depositional environments and seismic facies changes.

b) Structural interpretation to understand crustal architecture, basin margins, rifting style and fault

geometry

Additionally, two approaches were incorporated to fine tune the sensitivity of interpretation. These are:

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Top-down approach: Top-down approach is an analytical technique that starts with the big picture. It

breaks it down subsequently into smaller segments. The interpretation is mostly based on real

observation supported by background data. For example, a sinuous channel is reported in the literature

as well as seen in seismic data and eventually penetrated by a well. This represents a true and verified

data to be integrated into the seismic interpretation. Overall, this is a data-driven technique.

Bottom-up approach: Bottom-up approach is a synthetic technique to bind scattered information with

collective observations to give rise to a logical model, thus validating the original top down

interpretation. For example, if a high amplitude body represents sand in a well nearby, then similar high

amplitude bodies may represent sand as well. Overall, this is a model-driven technique.

4.1.1 Seismic and sequence stratigraphic interpretation

Seismic stratigraphy has been developed as a technique to interpret stratigraphic information from

reflection seismic data. The resolution of the seismic reflection follows the gross sedimentary sequences

and, as such, they approximate chrono-stratigraphic boundaries. The key contrasts represented by

seismic amplitudes and frequency come from compositional variation of geological formations from

both vertical and lateral variations in terms of facies changes.

The step by step methodology for seismic sequence stratigraphic interpretation begins with dividing the

seismic data into the discrete natural stratigraphic packages that make up the section. This is primarily

done by Identification and marking geometry of reflection and reflection terminations (Mitchum and

Vail, 1977; Galloway, 1989; Catuneanu, et al., 2008). The terminations can be represented by lapouts or

truncations (Figure 4.1). The lapout includes downlap, which is commonly seen at the base of

prograding clinoforms. It usually represents progradation of the basin margin (Figure 4.1). The onlap

represents termination of low-angle reflections against a steeper seismic surface (Figure 4.1). Similarly,

the toplap is a termination of inclined reflections against an overlying lower angle surface (Figure 4.1).

The different kind of truncation includes erosional truncation, where the strata are terminated against

an overlying erosional surface (Figure 4.1). Yet another type is a fault truncation, which is simply a

termination of reflections against a syn- or post-depositional fault, slump, or intrusion surface. As a

common rule, when particular reflectors terminate in a consistent manner they define a line in the

section, which represents a seismic surface. Once the seismic data has been divided into its component

depositional packages further geological interpretation may be attempted. Together with well log data

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143

this interpretation can be made into a sequence stratigraphic interpretation, predicting identification of

various sequence boundaries, transgressive surfaces and maximum flooding surfaces along with genetic

units determined by major allogeneic controls (Wagoner, 1995; MacNeil and Jones, 2006; Catuneanu et

al., 2008) (Figure 4.2).

Critical to the interpretation process is comparing how horizons and faults tie at intersections of seismic

profiles. Significant effort is expended correcting misties of faults, horizons, and sequence boundaries at

every profile intersection. The well penetrations were very helpful to correct the misties. In this regard,

closing the interpretation in loops around the seismic grid is a particularly effective technique.

In the present study, the surfaces have been identified as tectono-stratigraphic surfaces in broad

regional scale. These surfaces generally correspond to the second order sequence stratigraphic surfaces,

which are controlled by major tectonic events and major climatic changes. After all profiles were

interpreted and tied, the results of the interpretation were summarized and documented (See Table 3.1

and 3.2).

Figure 4.1: Stratal terminations that can be observed above or below a stratigraphic surface in seismic profiles and larger scale outcrops (from Mitchum and Vail, 1977)

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Figure 4.2: Architecture of facies, genetic units (systems tracts) and sequence stratigraphic surfaces (Catuneanu et al., 2008)

4.1.2 Structural interpretation

In passive margin setting, seismic data is commonly used for interpreting different types of crust using

“rift indicators” as a proof for the presence of the continental crust. This indicator would be a wedge of

sediments or volcanic rocks thickening towards the boundary fault, assuming that it may represent fill of

the failed rift unit, sedimentary fill of grabens, half-grabens and pull-apart basins deposited during the

syn-rift phase (Rosendahl et al., 2005; Figure 4.3). However, these indicators have rather limited values

unless there is depth of imaging allows seeing the entire crust hosting the rift-like features. Furthermore

some of the continental extensional allochthons may be present on top of exhumed mantle, which could

be may interpreted wrongly as continental crust (Manatschal, 2004; Rosendahl et al., 2005; Reston,

2007).

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145

Figure 4.3: Segment of the PROBE profile 5 showing the occurrence of the “rift indicators” on the proto-oceanic crust (Rosendahl et al., 2005).

A good quality wide-angle, deep seismic reflection profile became the primary tool for crustal

architecture interpretation at passive margins. The profile should be depth-migrated and deep enough

to image the entire crust, up to its Moho depth, to constrain the interpretation of crustal architecture.

Thanks to a large imaging depth in the ION Indiaspan data, the type of the crust could be determined

from the seismic image itself.

Based on the images themselves, the crustal layers could be defined. For example, the brittle upper

continental crust in reflection seismic images is generally transparent (Mooney and Meissner, 1992). It

contains extensional deformation structures such as grabens and half-grabens. Generally, the extension

related faults sole out at middle crustal levels at proximal margin, while they sole out at lower crustal

and sometimes at the top of continental mantle in distal margin (Figure 3.3). Images of the lower

continental crust, in general, are devoid of any discrete pre-rift and syn-rift deformational features due

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146

its ductile rheologic nature (Mooney and Meissner, 1992; Brun and Beslieb, 1996; Condie, 2005). The

lower crust (Figure 4.4) is characterized by relatively high reflectivity and parallel reflectors. The

boundary between upper and lower continental crusts is not a petrological boundary but a phase

transition from brittle to ductile state. In general, there is no sharp acoustic impendence contrast

between these two. Therefore, it is mapped as a zone instead of a surface.

Seismically, Moho is usually well imaged (Figure 3.15) because it represents a petrological boundary

between crust and mantle, both having distinct physical properties. The contact of underlying high

density and high velocity mantle rocks with overlying crustal rocks, which is have a relatively lower

velocity and density, produces a high enough acoustic impedance contrast to cause a distinct Moho

image. However, it also has been seen that at volcanic margins with thicker crust due to underplating.

The reflection seismic imaging is sometimes unable to image Moho.

Based on the impedance contrast Moho, can be interpreted on either continental or oceanic side. The

Moho depth in depth migrated seismic profiles can be further calibrated by thickness from seismic

refraction experiments in the known regions. The global average of oceanic crust thickness is around

7.08 km (White et al., 1992). Although there certain variations within normal range, the thickness of

oceanic crust is considered globally uniform. Therefore, it provides a nice calibration of Moho imaging.

On the continental side of passive margins, the crustal thickness is much less uniform. This further

complicates a Moho depth calibration. The crust is attenuated, and as the depth-dependent extension

theory describes, the upper continental crust thins down to almost zero at the continent-ocean

transition zone, where mantle is unroofed (Davis and Kusznir, 2004; Karner 2008). This means that the

Moho rose up in this zone (Figure 3.15). Therefore, the Moho depth, is therefore, depends on rate and

nature of extension.

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Figure 4.4: A strike-oriented seismic profile showing the continental crust in detail (Nemčok et al., 2013b). Top of the continental basement (TCB-1) is faulted. The upper continental crust (UCC) is marked by the presence of faults, shear zones and other pre-existing deformational features. It is relatively transparent in terms of reflectivity due to its brittle nature. The lower continental crust (LCC) is highly reflective with sub-horizontal reflector pattern due to its “flowable” ductile nature. The faults sole out at the top of lower continental crust (TLCC-1). As the TLCC-1 is not a petrological boundary, it is identified based upon above mentioned deformation environment criteria.

The seismic signature of unroofed continental mantle, referred as proto oceanic crust, is quite

distinguishable from continental and oceanic crusts (e.g. Meyers, 1996; Rosendahl and Groschel-Becker,

1999; Odegard et al., 2002; Odegard, 2003; Rosendahl et al., 2005; Nemčok and Rosendahl, 2006a and

b). The proto-oceanic seismic image is chaotic with noticeable discontinuities (Figure 3.15).

Theoretically, when mantle is exhumed at the continent-ocean transition and in a direct contact with

sea water and initial post-break-up sediments, the majority of exhumed peridotite is transformed into

serpentinite due to chemical reactions caused by downward percolating sea water (Manatschal and

Bernoulli, 1998; Odegard et al., 2002; Odegard, 2003). The invasive percolation is associated with

intensive fracturing reported from areas with unroofed continental mantle in outcrop (Manatschal,

2004).

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4.2 Potential field data interpretation

A rather modest methodology of crustal architecture interpretation comes from the study of potential

field data, particularly using gravity. Free-air, Bouguer and Isostatic residual gravity anomaly data can be

each used, although the combination is the best approach. The fundament of this methodology is the

isostatic equilibrium between adjacent oceanic and continental masses. The mass deficit due to water

column and mass excess due to presence of mantle below more or less uniformly thick oceanic crust,

which is less than the adjacent continental crust produces the characteristics gravity (Lillie, 1999). The

potential field study can be carried out in two ways

Qualitative analysis: This study is basically includes independent map based interpretation without

other constraints including seismic interpretation, crustal densities etc.

Quantitative analysis: The quantitative analysis eventually leads to forward or inversion modeling of

gravity values constraining the different crustal densities and interpreted seismic section.

4.2.1 Qualitative Analysis

The qualitative analysis includes both map based analysis and profile based analysis. Both the methods

have been used for East India and Elan Bank.

The following gravity maps have been used for in qualitative analysis of the data.

a. Free-air gravity anomaly

b. Bouguer gravity anomaly

c. Residual gravity anomaly

d. Isostatic residual gravity anomaly

The gravity data obtained from the station needs to be corrected at the beginning. Initially the raw

gravity data is processed to generate several anomaly maps. In this study a standard gravity processing

workflow is followed (See Blakely, 1995). The processing outline is given in Figure 4.5.

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Figure 4.5: The gravity data processing workflow.

The most primary data is the gravity readings (gobs) observed at each gravity station after corrections

have been applied for instrument drift and earth tides. The next step is the latitude correction where

the correction subtracted from Gobs which accounts for Earth's elliptical shape and rotation. The gravity

value that would be observed if Earth was a perfect (no geologic or topographic complexities), rotating

ellipsoid is referred to as the normal gravity (gn) given by:

gn = 978031.85 (1.0 + 0.005278895 sin2(lat) + 0.000023462 sin4(lat)) (mGal) (Eqn. 4.1)

Free-air gravity map: The Free-air gravity signal contains the integrated gravity response of the entire

earth from the centre of the earth to the farthest reaches of the atmosphere, including effects from

deep-earth structure, the Moho, basement, bathymetry, and geological structures of interest.

In order to correct for variations in elevation, the vertical gradient of gravity (vertical rate of change of

the force of gravity, 0.3086 mGal/m-1) is multiplied by the elevation of the station (h) and the result is

added, producing the Free-air anomaly. Free-Air gravity anomaly (gfa) correction is then shown as:

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150

gfa = gobs - gn+ 0.3086h (mGal) (Eqn. 4.2)

where h is the elevation (in meters) at which the gravity station is above the reference datum (typically

sea level).

The Free-air gravity anomaly includes the strong gravity effect of the water layer, which makes it difficult

to judge if an observed anomaly is caused by a geological feature in the subsurface or by the seafloor

topography. For average ocean bottom sediment densities in the range 1.9 to 2.2 g/cm3 and a seawater

density of 1.03 g/cm3, there exists a density contrast of approximately 1.0 g/cm3 at the sea floor. This

large density contrast may produce a large imprint on the Free-air gravity and consequently the Free-air

gravity map shows a strong correlation with bathymetry. The Free-air gravity map has been used in this

study to identify the shelf to slope break, crustal boundaries and volcanic ridges within oceanic crust.

Bouguer gravity map: The sharp Free-air gravity anomaly at the continental shelf edge can be explained

directly in terms of changing water depth. For this reason, it is customary to apply a Bouguer reduction

to the Free-air data. This is a first-order correction to account for the excess mass underlying

observation points located at elevations higher than the elevation datum (sea level or the geoid).

Conversely, it accounts for a mass deficiency at observation points located below the elevation datum.

Bouguer gravity anomaly (gfa) correction is given by:

gb = gobs - gn + 0.3086h - 0.04193r h (mGal) (Eqn. 4.3)

where r is the average density (gm/cm3) of the rocks underlying the survey area. Bouguer slab (r) is

considered 2.2 g/cm3 as a standard background value. The correction was computed on the topographic

surface so that the correction is referenced to the topographic surface.

This correction minimizes the predictable portion of the anomaly. The Bouguer reduction for a marine

survey replaces the water column with a density equal to the average density of the sediments at the

seafloor, thus ideally eliminating this large density contrast. During the Bouguer correction, the removal

of the water bottom density contrast enhances positive rise in gravity caused by the shallower mantle.

Thus, the Bouguer gravity anomaly would be expected to have eliminated the sharp negative Free-air

gravity effect associated with the rapid change in water depth across the shelf edge, but would be

anticipated to retain the longer wavelength, positive rise caused by the shallowing of Moho depth

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seaward along a passive margin. Positive anomalies can be associated with a thickened oceanic crust in

Bouguer gravity, which is sometimes non-unique in Free-air gravity.

Residual gravity anomaly map: Typically, a high-pass filter emphasizes short wavelength regional to

local anomalies removing the effect of long wavelength mega regional anomalies. It primarily enhances

gravity anomalies originating from basement structures (horsts, grabens), local basement intrusions, and

sedimentary structures. It has to be noted that for smaller wavelength features, the Bouguer anomaly

data tend to highlight the location of sedimentary basins and flexurally compensated features such as

rift shoulders if a high-pass filter is applied. A high-pass filter applied to the Bouguer gravity anomaly is

often effective for isolating gravity anomalies originating at basement and sedimentary intervals from

anomaly interference caused by deeper Moho undulations and crustal density changes, which are

associated with rifting and magmatic processes in the lower crust. A 10 km high-pass filter was applied

to the Bouguer gravity (Figure 3.29). The 10 km cut-off wavelength is the removes more than 50% of the

contribution of longer wavelength anomalies.

Isostatic residual gravity map: It is already implicit in the removal of the water bottom density contrast

with the Bouguer reduction is that there is no similar treatment for the positive increase in gravity

caused by shallow mantle isostatically compensating for deep water and thin oceanic crust. There is

typically a large density contrast at the lower crust/upper mantle, in the order of 0.30 g/cm3 to 0.40

g/cm3. However, the calculated gravity effect of the Moho is generally produces a long wavelength

anomaly. Therefore, the isostatic correction involves deriving a predicted depth to the Moho surface,

and calculating its gravity effect. When the isostatic effect is subtracted, the resultant anomaly map

shows both the density contrasts at the seafloor (Bouguer reduction) and at the Moho. This map is

obtained from Reliance industries.

The coupling of positive anomaly and negative anomaly has been uses as a guiding tool to interpret

crustal boundaries (Figure 3.31). It also enhances short wavelength local uncompensated anomalies

including solitary volcanoes or hotspot trails.

Qualitative gravity profile examination: Observations from gravity profiles in several passive margins

indicates that the crust attenuates during rifting there is a zone of thinning between the continental and

oceanic crust where there is a major increase in the gravity signal (Stewart et al., 2000; Bird, 2001;

Watts, 2001; Wyer and Watts, 2006; Kusznir, 2009; Watts et al., 2009). This position roughly indicates

the ocean-continent transition, which can be easily identified in gravity profile curves. The careful

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gravity profile examination also indicates several anomaly pattern, which can be originated from several

or common source, depth of the source and nature of the anomaly. This method is quite useful in Elan

Bank as the gravity profiles indicate that a transitional zone can be observed on all five profiles where

the gravity anomaly values drop significantly.

4.2.2 Quantitative Analysis

Seismic interpretation can be constrained through forward gravity modeling. The seismic interpretation

allows placing Moho surface on the lowest strong reflector to differentiate the oceanic, continental

crusts and unroofed mantle according to their thickness, seismic nature. The 2D gravity inversion along

seismic profile allows assigning different densities to crustal blocks and layers interpreted from seismic

images and matching the observed and calculated gravity curves. While it is possible to end up with

several possible scenarios using only gravity modeling but the unique scenario comes from constraints

provided by interpreted seismic data. The magnetic data can be used as additional constraints.

4.2.2.1 Forward modeling along profiles

The 2D forward modeling was carried out in cross sections along interpreted seismic profiles. The

primary objective of the 2D forward modeling was to define the geological source of the gravity and

magnetic anomaly. This method is useful to create a geologic model of the subsurface with integrated

seismic interpretation, well calibration, and observed gravity and magnetic measurements. However,

the objective of forward modeling in each profile may be different, in the attempt to achieve some

primary understanding. The forward modeling has been carried out selected survey lines in East India

along with Bird Geophysical using the selected ION Indiaspan profiles and ship-track gravity and

magnetic data acquired along these profiles as mentioned in previous chapter (Bird, 2009). This attempt

could not be possible in Elan Bank profiles due to lack of depth converted seismic profiles.

Since Bouguer anomalies express only relative gravity data, density contrasts of all lateral

inhomogeneities also affects the modelled field (Christensen and Mooney, 1995). Therefore and to

avoid edge effects, a reference density model has to be considered for the use of absolute densities,

which on top provides better opportunities to compare the density model with petrologic or seismic

velocity models. The typical crustal density incorporated in this study is provided in Table-4.1

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(Christensen and Mooney, 1995). For consistency, density values were held constant for each of the

modelled layers for all models.

Mass Type Densities (gm/cc)

Water 1.03

Sediment 1 2.19

Sediment 2 2.25

Basalt 1 2.9

Basalt 2 (Intra-Basalt 1) 2.9

Sediment 3 (Synrift top) 2.6

Upper Crust 2.7

Lower Crust 3.0

Mantle 3.36

Table 4.1: Density used for gravity forward modeling

The modeling assumed specific initial subsurface density model to calculate gravity and then compare it

with the data driven synthetic interpretation. The modeled densities were adjusted in order to match

the calculated gravity curve with the observed gravity data curve within a constraint of geologically

logical interpretation (Blakely, 1995). Therefore, this method is an iterative way by modifying source

body configurations and/or densities and susceptibilities after each computation until a “best fit” was

obtained between the observed and calculated data profiles.

Cross sectional modeling of source body geometries and rock properties, like density, is usually an

effective method and it is useful when different scenarios are tested with seismic and potential field

observations. The process of modeling involves changing rock property values and modifying source

body geometries. Modifying source body geometries, in turn, involves moving-deleting-adding vertices

to source body polygons, splitting-joining source body polygons, and changing their dimensions

perpendicular to the cross section (Blakely, 1995; Bird et al., 2011). It is again critical to note that all the

modeling solutions are non-unique and they should be used for validation or ruling out alternate

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interpretation. The non-unique nature of potential field modeling means that for any given observed

anomaly, there are multiple source body models, which can yield matching, computed curves. As such,

the confidence level in the resultant geologic model is dependent not only upon the degree of forward

model fit but also upon the quality and quantity of geologic and geophysical constraints employed

during the modeling. Such constraints often include, for example, control from well logs, seismic

sections, correlation to surface geology, sampled susceptibilities and densities, isopach maps and others.

Concurrent integrated modeling of both gravity and magnetic data often eliminates some ambiguity. At

the end, a synthetic approach was considered, where seismic interpretation can be constrained through

forward gravity modeling.

4.3 Gathering petrological data on Elan Bank/Antarctica/East India for geological correlation

This technique is primarily based on matching the stratigraphic sections, crustal provinces and petro-

tectonic assemblages (Scotese et al., 1988; Coffin and Eldholm, 1992; Harley and Motoyoshi, 2000;

Reeves and de-Wit, 2000; Alvey et al., 2008; Kearey et al., 2009). This is a nice qualitative technique to

validate quantitative geometric reconstruction of separated continental masses. As the location of the

rift is often controlled by the geology of the supercontinent, which takes place along zones of pre-

existing weakness, it is possible to trace continuous geologic features from one continent to another

across the reconstructed continental boundaries, if they are associated with true ancient configurations

of continents (Cawood et al., 2006; Alvey et al., 2008). The presence of Nagavali Shear zone in East India

may show a relationship with right lateral offset feature in Elan bank as it can be seen in predicted

bathymetry map (Figure 3.27). The analytical technique including textural and mineralogical

interpretation and whole rock geochemistry for major and trace elements from samples collected from

Eastern Ghat rocks in East India (Gopalkrishnan, 1998; Chetty et al., 2003; Biswal and Sinha, 2004) and

ODP site -1138 in Elan Bank (Nicolaysen et al., 2001; Frey et al., 2002; Ingle et al., 2002; Ewart et al.,

2004) and Antarctica (Harley and Hensen, 1990; Harley and Motoyoshi, 2000) is crucial to achieve the

crustal continuity between these fragmented land masses.

The correlation of the stratigraphic age patterns across the conjugate margins is a good tool to match

both margins. In this technique, distinctive and unique igneous provinces including extrusive and

intrusive rocks can be traced between the continents. They can be geochemically and radiometrically

analyzed as a qualifier for reconstructions (Ingle et al., 2002; Ewart et al., 2004). 40Ar/39Ar dating is found

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to be very useful to determine synchronicity of volcanic emplacements during final breakup or outbreak

of any hotspot related volcanism. This particular dating method successfully separates the whole-rock

samples of rhyolite ignimbrites and basaltic lavas from the pre- and syn-rift flood volcanic units and thus

establishes a temporal link between conjugate rifted margins like East India and Elan Bank (Ingle et al.,

2002).

4.4 Tectonic reconstruction between Elan Bank and East India

The tectonic reconstruction between East India, Elan Bank and Antarctica is important to understand the

microcontinent separation during continental breakup. As mentioned in Chapper-1, most of the

available plate reconstruction model do not account for Elan Bank in their India-Antarctica fit, although,

few models do account for Elan Bank juxtaposed against the Mahanadi Basin. However, the exact

methodology for such fit is unexplained.

The initial reconstruction was followed using the geometric reconstruction (Reeves, 2008) and marine

magnetic anomalies (Müller et al., 1997). However, the final reconstruction is a synthetic reconstruction

model developed using all the relevant information to restore the hyper-extended passive margins.

Similar method has been applied to restore hyper-extended passive margins in Iberia-Newfoundland

(Corsby, et al., 2008; Sutra et al., 2013).

Geometric reconstruction: This method involves matching continental borders based on series of poles

of rotation for each pair of continents arranged in a grid of latitude and longitude positions (Bullard et

al., 1965; Schettino, 1998; Reeves and de-Wit, 2000; Figure 4.6). For each pole position the angle of

rotation is determined that brings the continental margins together with the smallest proportion of gaps

and overlaps. As the angle of rotation is finalized, the sensitivity of the fit is quantified by a degree of

mismatch. Both the crustal type and the degree of extension within the ocean-continent transition zone

are estimated on a point-by-point basis from bathymetry and basement depth (Reeves and de-Wit,

2000). The root-mean-square (RMS) misfit in the apparent total closure of both oceanic crust and

extension within the continental crust is generally computed for a range of possible poles of opening

(Dunbar and Sawyer, 1987). Whenever possible, finite rotations are calculated by matching conjugate

isochrones (Müller et al., 1997). A contour map of misfit versus pole location reveals a global minimum

RMS misfit of an Euler pole. The computed sensitivity of fit is generally known as the objective function

(Kearey et al., 2009). Values of the objective function are entered on the grid of pole positions, and

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contoured. The location of the minimum objective function provides the pole of rotation for which the

continental edges fit most exactly (Kearey et al., 2009).

Figure 4.6: Fit of the continents around the Atlantic Ocean, obtained by matching the 500 fathom (920 m) isobaths (Bullard et al., 1965)

It has to be noted that the geometric reconstruction considers a tight continental assembly using

present day coastline or 1000 m isobaths. In such geometric reconstruction, it is evident that present

day coastline fit is highly erroneous and can be used as a guiding tool only. The present day coast line is

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a geomorphic expression and does not limit the continental masses in passive margin. A major problem

with isobaths arises from the observation that the shape of these lines is strongly affected by the

sedimentary cover particularly the area where high riverine sedimentary inputs are there (Schettino,

1998). The delta progradation in these regions shifts the 1000m isobaths further seawards by hundreds

of kilometers in few million years. This phenomenon clearly indicates that even though conjugate

isobaths may fit accurately at the beginning of the ocean opening, subsequent sedimentation generally

decreases the precision of geometric fit.

Magnetic anomalies and fracture zones: This method is based on the Vine–Matthews hypothesis (Vine

and Matthews, 1963). According to this hypothesis, geometric reversal of the Earth’s magnetic field

explains the sequence of magnetic anomalies away from ocean ridges in terms of normal and reversed

magnetizations of the oceanic crust acquired during polarity reversals (Vine, 1966; Figure 4.7).

Interestingly, rates of geomagnetic reversals are highly variable in the geologic past. There has been a

series of quick rates of reversals during the Cretaceous, followed by long quite period for about 35 Ma

within, during which the field was of constant normal polarity.

Figure 4.7: Sea floor spreading and the generation of magnetic lineations by the Vine-Matthews hypothesis (Kearey et al., 2009)

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Once the geomagnetic reversal timescale has been calibrated, oceanic magnetic anomalies may be used

to date oceanic lithosphere (Vine, 1966; Torsvik et al. 2001; Müller et al., 2008; Figure 4.8).

Simultaneously, it is possible to develop very close models of all observed stripped anomaly sequences

to determine the spreading rates at oceanic ridges. Combining these two observations, it is now possible

to deduce ages back to mid-Jurassic times with an accuracy of few million years (Müller et al., 1993). In

this method, the entire ocean basin is zipped up using the magnetic anomaly data together with oceanic

fracture zones to provide relative fits between continents. This technique is also useful for

understanding ocean basin evolution (Dunbar and Sawyer, 1987). This method is globally applicable

from mid-Jurassic times to present (180 Ma onwards). It is to be noted that magnetic anomaly fits must

be considered minimum fits since they do not account for a pre-drift extension. For example, the

majorities of Pangaea reconstructions are Late post Triassic/Early Jurassic or younger reconstructions

and assume insignificant intra-plate deformation since the Permian (Torsvik et al. 2001). Additionally,

this method is a difficult one for reconstruction in the areas, where the oceanic crust developed during

Cretaceous long normal magnetic super-chrorn (124 to 83.5 Ma).

Synthetic reconstruction: A simple method for estimating the amount of extension at passive conjugate

margins requires the combined use of gravity, magnetic and seismic interpretations to get a firm control

over the basement configuration, nature of ocean-continent transition and Moho depth. A series of

synthetic cross sections, which depicts the water column, the sedimentary section, the crustal

architecture with Moho depth has been prepared by combined interpretation for both conjugate

margins. These profiles and cross sections must act as crustal constrains to achieve a closer

approximation of the shapes of rifted margins to fit together for a more precise pre-rift geometry. This

general strategy was adopted in reconstructing East India and Elan Bank and Antarctica. The structure of

the crust has been constrained by at least five modern wide-angle seismic sections. However, the

Antarctica constraints are coming only from public domain information.

Hence, in this study the fit between East India-Elan Bank-Antarctica was not carried out following any

standard plate reconstruction model. Instead, a combination of magnetic anomaly stripes and geometric

correction was applied based on interpreted architecture of thinned continental crust, exhumed mantle

and oceanic crust. The geometry of crustal boundaries like continental-proto-oceanic or proto oceanic-

oceanic boundaries has been used to constrain the geometry of the East India margin and continental

margin of Elan Bank.

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Figure 4.8: Ocean floor age map (Müller et al., 2008). (a) Distribution of age observations based on marine magnetic anomaly identifications (rotated magnetic anomalies are in red and observed ones are in black). (b) Age-area distribution of the ocean floor based upon magnetic anomaly (c) Gridded age uncertainties. All the panels are shown in the Molleweide projection. Continental margins are medium gray, continents are light grey, and plate boundaries are indicated by black lines.

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4.5 Seismic profiles marriages from conjugate margins

A precise way of testing correctness of the interpreted crustal type distribution and crustal architecture

of conjugate margins is the seismic profile marriage technique (Rosendahl et al., 2005). This technique is

used for the verification of placing Elan Bank against East India. This methodology suggests bringing

interpreted reflection seismic profiles from conjugate margins and pairing their external proto-oceanic

ends or external continental ends together (Rosendahl et al., 2005). This technique is well applicable to

ocean segments with roughly orthogonal rifting. It may be complicated in a case of complex drift history

with strike-slip or oblique-slip tectonics preceding the continental break-up like in the same segments of

East India margins. However, it is possible to achieve some precision in strike-slip margins by considering

the sensitivity of interpretation and other geological and geophysical observations.

In this study, based on the tectonic fit, the profiles have been chosen as these intersect at conjugate

positions near the merged oceanic/proto-oceanic or proto-oceanic/continental crustal boundaries. It is

to be noted that both the paired profiles is kept in same vertical and horizontal scale. The married or

paired seismic profiles are then brought together to test the continuity of the crustal architecture and

synrift geometry. It is very important to constrain the marriage using the interpreted faults related to

thinning at present day distal margins in both continents. Again, the match is only valid up to continental

breakup (e.g. true onset of organized sea floor spreading). Therefore, the stratigraphic section younger

than that is beyond consideration. The seismic marriage is constrained by profiles through the

orthogonal rifted margin in the Krishna-Godavari region to confirm the restoration of Elan Bank against

East India. It has to be noted that the geographic East India-Elan Bank fit does not provide any scope for

seismic marriage with orthogonal Cauvery rift zone. However, its proximity to Coromandal strike slip

system has been used to constrain the reconstruction.

4.6 Tectonic timing

The details of timing of various tectonic events recorded by passive margin are a powerful technique to

document the important tectonic events and to constraint the regional plate reconstruction models.

Although, this technique provides relative timing arguments, these are essentially used for fine tuning

the seismic interpretation by building tectono-stratigraphic framework. This technique is applied to

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create the time frame for continental breakup along different parts of the margin and to constrain the

fit between East India and Elan bank. The timing technique is proved to be crucial in this study to decide

the best possible one between the two hypotheses introduced in Chapter-1.

4.6.1 Break-up timing based on borehole data

A series of borehole data drilled on East India margin have been selected to define the continental

breakup timing along East India. These wells have been systematically sampled, cored and logged. The

well samples have been analyzed for biostratigraphy as well as paleo-environment based on planktons,

nannofossils and palynology data. The composite information for age markers as well as plaeo-

environment data has been used for regional integration and stratigraphic boundary constraints. The

timing argument is primarily based on the interpretation of paleo-environment derived from

biostratigraphic information.

In rift basins, the nature of extensional faulting and climatic factors controls the morphology and

topography. Hence, the drainage and related sedimentary depositional environments are influenced by

these controls. In a global rift model, it has been observed that the continental fluvial-lacustrine

environment prevails during the synrift syn-tectonic phases (Ravnas and Steel, 1998; Fraser et al., 2007).

These sections are generally identified and benchmarked by palynological information. As the

subsidence progresses with more extension, a marine incursion occurs at the end of synrift syn-tectonic

phase or later during the syn-extensional sag phases. In a completely preserved section, this event is

followed by prolonged thermal subsidence as well as rapid drowning by deeper marine sequences.

Therefore, it is possible in a gross sense to delineate the synrift and post rift sequence changes in

depositional environment and subsidence pattern. This information is further supported by

palynological data, which can help us to determine the timings of major events vis. a vis. timing of

continental breakup. Differentiations of the shallow water deposits are bit difficult as they are very

much condensed and in most cases these can be either late synrift or early post-rift ones. However, this

can be combined with seismic interpretation to resolve the issue.

The biostratigraphic information is also vital to identify missing and condensed sections. These gaps or

unconformities are usually caused by erosion instigated by the isostatic uplift occurring as a result of

continental breakup. Isostatic uplift of rift margin is a response to displacement on major fault, which is

gradually ceasing at the end of extension in the lithosphere vis. a vis. close to continental breakup.

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Identification of uplift events is crucial to determine the breakup time as well. To add precision to timing

arguments, 1D subsidence modeling was carried out to determine the timing of the events including

subsidence and uplift by integrating the interpretation with reflection seismic data.

It is to be noted that the borehole location is important. The subsidence and isostatic uplift history in

proximal and distal margin may vary in a hyper extended margin. In this study, the boreholes are

selected in way, which gives a combination of proximal and distal margin wells. The continental breakup

timing is thus determined by careful examination of borehole information by interpreting environment

of deposition, biostratigraphic information and 1D subsidence models.

4.6.2 Break up timing based on seismic interpretation

As mentioned earlier, the typical rift indicators are represented by wedge pattern of reflectors

thickening against fault. Such pattern is called half graben geometry (Figure 4.9). The top of the half-

grabens are usually sealed by an unconformity surface. This unconformity can be angular unconformity

or disconformity. The disconformity cannot be detected in reflection seismic images without

biostratigraphic help. On the contrary, the angular unconformity can be interpreted and mapped across

the region rather easily. The most typical one is the breakup unconformity. This unconformity is well

interpreted in the studied dataset in East India and Elan Bank. The regional correlation of breakup

unconformity tied with borehole information has been used as a tool to determine the timing of

continental breakup. This technique is very useful for quick regional interpretation. However, as this

unconformity surface is diachronous in spatial as well as lateral sense, more biostratigraphic information

has been used to loop-tie the information in regional scale.

It is very important to note that the typical breakup unconformity at proximal margin may not represent

the actual continental breakup (Manatschal, 2004). The syn-tectonic activity may soon cease after the

stretching phase in proximal margin and extension may continue in thinning phase in distal margin

(Manatschal, 2004; Karner, 2008). Therefore, it is important to identify some equivalent sections of syn-

extensional sag deposits at proximal margin. These deposits are shallow marine in nature and often

restricted to the central part of the basin. They are represented by some bowl shaped reflection packet,

tapering at two ends. However, it is extremely difficult to identify such sections, particularly in East

India, which represents a fast rifting scenario (Nemčok et al., 2013b). As a consequence, the syn-

extensional sag deposits are very much condensed or absent. In such a case, the biostratigraphic

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information was essential to interpret or proving the absence of such phases. Therefore, the sequence

or surface, which is represented by the first deep marine incursion that envelopes proximal and distal

margin has been used as a proxy to date the true continental breakup timing.

Figure 4.9: Progressive extension along major basin bounding fault and sedimentary deposition found in the intra-basin combine to produce a wedge-shape sedimentary unit that thickens toward the major basin bounding faults.

4.6.3 Break-up timing based on magnetic stripe anomaly data from literature

Interpreted magnetic anomaly stripes and associated geometry of fracture zones is an undisputed tool

to date genesis of oceanic crust due to organized sea-floor spreading. The magnetic anomaly

fingerprinting is a very specialized technique and is not adopted in this study. However, the interpreted

datasets (Royer et al., 1991; Ramanna et al, 1994; Krishna et al., 1995; Krishna and Rao, 2000; Ramanna

et al, 2001; Müller et al., 2008) were taken to validate the interpretation (Figure 4.10).

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The global database on the age of ocean floor was created by Müller et al. (2008), by synthesising the

most of global magnetic anomaly datasets available in public domain. Associated map is useful for global

interpretation (Figure 4.8). However, the caution was taken for the same reason as that with afore

mentioned magnetic anomaly interpretation in Bay of Bengal. The problem is that the dataset is

interpolated in many parts of Bay of Bengal due to lack of reliable magnetic anomaly data.

Figure 4.10: Different magnetic anomaly datasets in Indian Ocean combined from various published sources. The Cenozoic magnetic anomaly data is derived from Krishna et al., (1995), The Magnetic anomaly near Sri Lanka is from Desa et al., (2005) and the Mesozoic magnetic anomaly is from Ramanna (2001).

In general, the oldest magnetic anomaly represents the oldest oceanic crust adjacent to passive margin

setting. Therefore, it directly indicates a relative timing of continental breakup, which takes place prior

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to development of the oldest oceanic crust. If the anomalies are located close to the proto-oceanic-

oceanic crust boundary, the precision is relatively good. However, there are certain disadvantages to

this technique. The magnetic anomaly may form over serpentinized exhumed mantle of the proto-

oceanic crust corridor (Sibuet et al., 2007), as a result of magnetite remagnetization sometime after its

emplacement in those rock suites. Therefore, this may mislead the interpretation of the actual

continental breakup timing. Additionally, as the oceanic crust adjacent to the East India margin is

developed mostly during Cretaceous long normal magnetic polarity zone, the imprints of magnetic

anomaly are weak. The close interaction with Kerguelen plume may also modify the magnetic anomaly

signatures of the oceanic crust. Due to such complications, this technique is grossly used to determine

the timing along with borehole and seismic information.

4.6.4 Hot-spot activity timing for exact locations

This technique is used as based only on published information of radiometric dates. The Kerguelen

plume activity is well documented and dated event. The basalts from various parts of Kerguelen plateau

and associated volcanic provinces like Rajmahal Trap in East India, Bunbery basalt in Western Australia

have been preciously dated using radiometric methods (Storey et al., 1992; Coffin et al., 2002; Ghatak

and Basu, 2011). A comparative study was carried out between these actual radiometric dates from

basalts due to Kerguelen plumes and the continental breakup timing from other methods. This exercise

again provides a clue between relative time difference between Kerguelen plume emplacement and

Elan Bank-India separation.

4.7 Match of Elan Bank with hot-spot track in Bay of Bengal

Several hotspot tracks are present in Bay of Bengal. Among them, the 85° E Ridge hotspot track (Figure

3.26) has a close proximity to East India and thus Elan bank. Hotspot tracks are again not directly dated.

However, the published information about the radiometric dates has been plotted on the anomaly

pattern of hotspot track. This provides a relative idea of emplacement timing of 85°E ridge and its spatial

variation. The reconstruction and the weak magnetic anomaly pattern in Bay of Bengal have been used

to create the trail of Elan Bank microcontinent that broke up from East India. The relative timing

between hotspot activity and separation of Elan bank thus can be a tool to determine the time

differences between hotspot activity, oceanic crust formation and continental breakup.

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Chapter 5

Interpretations

The interpretation of different data sets described in this chapter follows the contents of the Methods

chapter.

5.1. Seismic Interpretation

The seismic interpretation has been carried out on available seismic data as described in section 3.1 in

East India as well as Elan bank in order to recognize the rifting style, rift fault geometry, margin

segmentation, crustal architecture and respective crustal boundaries. It was useful for understanding

the evolution and formation of passive margins.

5.1.1 East India

The interpreted sections are distributed from southern most Cauvery Basin to Krishna-Godavari Basin in

northern end.

5.1.1.1 Crustal Architecture

The crustal architecture is best visible in the interpreted section ION-GXT-1000 (Figure 3.15), which is a

NW-SE trending dip-oriented profile in the south of the Krishna-Godavari Basin. The interpretation on

this line is quite an eye opener as it avoids crossing any of the prominent magmatic features present in

Bay of Bengal. This profile effectively displays distribution of different crustal types, including thinned

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upper continental crust, thinned lower continental crust, proto-oceanic crust (exhumed mantle), and

oceanic crust.

The brittle upper continental crust has a transparent character with very low amplitude chaotic

reflection patterns. It is characterized by the presence of rift-related deformation, which includes

faulting, fracturing and block tilting, formed in the brittle deformational environment as it can be seen in

several other profiles in the Krishna-Godavari and Cauvery basin (Figures 3.1, 3.2, 3.3, 3.11, and 3.12).

The rift faults usually die out at the middle crustal levels (Figure 3.1, 3.12) or near the top of the lower

crustal levels in the proximal part of margin. In certain places, the upper crust includes older

deformation structures, possibly from its pre-rift tectonic history including Precambrian orogenic events

(Figures 3.1, 3.2). The data suggest that the upper continental crust thins down to almost zero at the

transition between continental and oceanic domains. In this region, the upper crustal faults sole out on

top of the thinned lower crust or on top of the exhumed mantle. In this context, it is to be noted that

the seismic profile example of the Palar segments connecting Krishna-Godavari and Cauvery segments

(Figures 3.5, 3.6, 3.7, 3.8) representing the Coromandal dextral strike slip transfer zone (Nemčok et al.,

2013b) contains semi-transparent upper continental crust with some low amplitude reflections in

comparison to the extensional Krishna-Godavari and Cauvery basin.

The lower continental crust in seismic images is characterized by relatively high reflectivity, parallel

reflectors, in contrast to the upper crust and uppermost mantle, which both appear almost transparent.

The lower crustal reflection package wedges-out towards the ocean continent transition represented by

a zone of exhumed continental mantle (proto-oceanic crust) (Figures 3.1, 3.2, 3.3, 3.8, 3.11, 3.15). This

highly reflective seismic package is more or less continuous from the shelfal region to the ocean-

continent boundary area. It is recorded along strike over large distances from the Mahanadi through

Krishna-Godavari to Cauvery basins in interpreted seismic profiles. This seismic signature of lower

continental crust is more prominent in the Cauvery Basin segments in comparison to the Krishna-

Godavari segment. This may be due to the different Precambrian basement types and associated

rheological differences in the Cauvery and Krishna-Godavari segments. As mentioned in section 2.3, the

Cauvery basement can be characterized by the high-grade metamorphic rocks (charnokites) belonging

to the Southern Granulite terrain (SGT) (Ramakrishnan et al., 1998), in comparison to the Eastern Ghats

mobile belt with high grade metamorphic rocks in the region adjacent to the western portion of the

Krishna-Godavari Basin (Naqvi and Rogers, 1987; Gopalakrishnan, 1998; Ramakrishnan et al., 1998). The

differences in basement type along the margin also contribute to the rheology and thermal structure of


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the crust, which significantly influences the nature of extension along the margin. The absence of the

discrete pre-rift and syn-rift deformational features in the lower crust is attributed to its ductile

behavior. The characteristics features include the high-reflectivity parallel structures anastomosing

detachments surfaces, boudinaged structures, material flow etc. (Figures 2.8, 2.10). All these

observations indicate a ductile character of the lower crust, which developed at certain stage of rifting.

The boundary between the upper and lower continental crusts is not an abrupt termination but rather

transitional. Consequently, it is much more difficult to map this boundary both in time and space in

contrast to mapping a top of continental mantle. Given the distinctively different seismic images of

upper and lower crusts, the boundary between them was roughly interpreted as a transitional zone

dividing two different reflection domains. Thus, it roughly divides the underlying reflective layer with

sub-horizontal layering from the overlying semi-transparent layer with various brittle deformation

structures.

Being a petrological boundary, also truncated by a mineralized fault plane developed at exhumation

phase of hyper-extension process (Manatschal et al., 2007), the top of continental mantle in this profile

is well imaged as a high-amplitude reflective surface. The boundary between the underlying high-density

and high-velocity mantle rocks, and overlying relatively low-density and low-velocity crustal rocks,

produces a sharp impedance contrast in the reflection seismic data, which images the Moho distinctly

beneath the thinned continental crust (Figure 3.15). Once the interpretation was constrained on profile

ION-GXT-1000, the same principle is applied to interpret the same in the other seismic profiles along the

East India margin as shown in Figures 3.1, 3.2, 3.3, 3.4, 3.5, 3.6, 3.7, 3.8, 3.9, 3.10, 3.11, 3.12, 3.13, 3.16

and 3.17). The same principle was applied to interpret a good oceanic Moho, where the impedance is

produced by the density velocity contrast between the basaltic oceanic crust and peridotite dominated

mantle rocks.

It appears that continental mantle was unroofed from beneath the continental crust and formed a zone

of exhumed continental mantle eastward of the continental margin in profile ION-GXT 1000(Figure

3.15). In this profile, a transitional region 80-85 km wide occurs between the last continental fault block

and the first appearance of a reflection from oceanic Moho, with a similar pattern of basement

morphology with similar crustal reflectivity. Similar transition zones ranging from 70 to 100 km in width

are also interpreted in ION-GXT and RIL profiles in the Cauvery, Godavari and Mahanadi basins (Figures

3.1, 3.2, 3.3, 3.7, 3.8, 3.9, 3.10, 3.11, 3.12, 3.13, 3.16 and 3.17). However, it is to be noted that in


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contrast to the similarities in thinning of continental crust, the seismic zonation of this crustal type is

highly dissimilar in different profiles. It exhibits a much broader region across the Krishna-Godavari basin

compared to the profiles in the Cauvery basin.

In all the profiles, the zone of exhumed continental mantle (i.e. proto-oceanic crust), which is located

between the thinned continental crust, and normal oceanic crust is characterized by highly chaotic

reflection pattern with occasional presence of high amplitude events. While the top of the proto-oceanic

crust is imaged by a high amplitude reflective zone, its base does not have any distinct signature in most

cases as the proto-oceanic crust made of the exhumed continental lithospheric mantle has poorly

imaged base in the available seismic datasets. This observation, again, matches with observations made

at different passive margins around the world, where seismic velocities and densities of the proto-

oceanic crust are not characteristic of oceanic crust (Meyers et al., 1996 and 1998; Wilson et al., 2003;

Rosendahl et al., 2005). This is in agreement with observed densities, which can be seen in the present

study area and described in subsequent section 5.2.

In some of the profiles (Figures 3.12, 3.16, and 3.17) particularly in the Krishna-Godavari and Mahanadi

basins, the top of interpreted exhumed continental mantle is covered by a thick volcanic cover. In the

Mahanadi Basin, due to the proximity of the 85° E Ridge, the image of the proto-oceanic corridor is not

very clear but the presence of gravity anomalies identical to those located further west suggests the

presence of a similar high-density mantle material.

The best indication of oceanic crust is the presence of the prominent sub-horizontal Moho reflector

(Figure 3.15). The oceanic crust is characterized by relative low amplitude, high frequency, sub-

horizontal and parallel reflectors, and apparent lack of any internal deformation unlike the neighboring

zone of exhumed continental mantle further to its west. The Moho depth in depth migrated seismic

profiles can be further calibrated by thickness and the Moho depth refraction documented from the

experiments in the margins of just by comparing oceanic thickness values (Mooney and Meissner, 1992;

Christensen, and Money, 1995; Rosendahl and Groschel-Becker, 1999; Rosendahl et al., 2005). In the

current study both the interpretation and calibration are similar to observations made at various

continental margins developed by slow extensional rates indicating both mantle exhumation and

subsequent serpentinization (Manatschal, 2004). As a result, seismic velocities and densities of the

proto-oceanic crust are not characteristic of oceanic crust (Meyers et al., 1996 and 1998; Wilson et al.,

2003; Rosendahl et al., 2005). This is in accordance with densities, which can be observed in the study


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area. The thickness of the normal oceanic crust is usually uniform ranging from 4 km to 7km with an

average thickness of 5.4 km. However, the thickness of oceanic crust underneath the isolated hot-spot

volcanoes and main hot spot track is considerably thicker. Isolated volcanoes are associated with crustal

thickness ranging from 7.8 to 8.5 km. The main hot spot track is underlain by an oceanic crust with a

thickness of 8.3-13 km, having an average thickness of 10.5 km (Nemčok et al., 2013b).

It is to be noted that the crustal thinning along the tectonic transport direction is more gradual across

the narrow continental slope with a strong signature of moderate sea-ward dipping break-away fault,

which demarcates the proximal and the distal margin and eventually dies out near the top of the

continental mantle as seen in ION-GXT profile 1000 (Figure 3.15). The thicker continental crust in shelfal

region is thinned down to zero thickness near the ocean-continent transition within a distance of 100

km. The thinning of continental crust in a similar fashion can also be observed in the northern profiles in

the Krishna Godavari Basin, where the transition occurs within an average distance of 80-100 km (Figure

3.12). The nature of crustal thinning is alike in the Cauvery basin in the south, where the transition

occurs through a moderate seaward dipping break-away fault demarcating boundary between the

proximal and distal margins, but the crust thins down to zero within an average distance of 100-120 km.

Although in profiles like ION-GXT 800, there is some crustal inversion observed with very sharp ocean-

continent transition as shown in Figure 3.5. In the region, between the Krishna-Godavari and Cauvery

basin, the Pennar-Palar segment shows a somewhat different crustal transition. Here, the thick

continental crust rather abruptly thins down to zero towards the ocean-continent transition through the

steep seaward dipping break-away fault. The transition also occurs within an average of distance 50-60

km from the initial thickness in proximal margin to the zero thickness near the ocean-continent

transition.

This crustal architecture of the East India, therefore, typically indicates the hyper-extended crustal

architecture model in analogy with other passive margin observations particularly at the Iberia-

Newfoundland margins (Davis and Kusznir, 2004; Huismans and Beumont, 2005; Manatschal et al.,

2007). The same crustal interpretation has been combined with interpreted fault maps to further

constrain the nature of rifting style and nature of continental breakup as described in next section.


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5.1.1.2 Rifting style, fault geometry and nature of continental breakup

The interpretation of an integrated model of East India passive margin is carried out in conjunction with

crustal interpretation and the rifting style and nature of continental breakup, which eventually leads to

understanding of different phases of extensional process. This basic interpretation sets the identification

of important building elements of hyper-extended passive margin model.

The typical syn-rift faults consist of both normal faults and strike-slip faults depending upon the nature

of extension. Normal faults usually control the overall extension, whereas the strike slip faults are acting

as linking transfer faults between zones of orthogonal extension. The normal faults have typically listric

geometries as it can be seen in profiles (Figures 3.1, 3.2, 3.3, 3.9, 3.10, 3.11, 3.12, 3.13, 3.16 and 3.17).

They can be interpreted in seismic data as major discontinuities in most cases against which truncations

of seismic reflectors are observed. The faults are well interpreted through the given seismic cross

sections from the Cauvery Basin to the Krishna-Godavari Basin as shown in profiles (Figure3.1, 3.2, 3.3,

3.4, 3.5, 3.6, 3.7, 3.8, 3.9, 3.10, 3.11, 3.12, 3.13, 3.16 and 3.17). However, in some exceptional cases,

heavily mineralized normal faults are directly imaged as it can be seen in some profiles (Figure 3.16). The

fault core mineralization is possibly caused by hydrothermal activity and associated mineralized damage

zones are thick enough to be seismically imaged directly as a prominent reflective zone. A large volume

of mineralizing fluids produces a diffuse zone interpreted as altered rock mass developed as a result of

hydrothermal cementation.

Based on observations from rift faults across the Cauvery and Krishna-Godavari Basins, it is apparent

that their controlling faults are apparently dip-slip to slightly oblique-slip normal faults. There is

generally little or no evidence of significant strike-slip faulting in areas of high-angle normal faulting.

However, a minor lateral slip occurs where a NE-SW propagating fault tip is deflected by a preexisting

basement shear zone oriented at a high angle to the propagating rift fault as shown in the tectonic

element map (Figure 5.1). The broad zones of lateral shears are accommodated by a combination of

minor vertical axis rotations and distributed oblique-slip normal faulting in which the dip-slip component

is dominant. Similar situations are observed in the East African Rift System (Morley, 1999) and the Rio-

Grande rift (Chamberlin, 2000).

Strong horizontal axis rotation and dip-slip separation of numerous sub-parallel fault blocks characterize

the highly extended Cauvery and Krishna-Godavari regions. In the southern most Cauvery Basin the

continental proto-oceanic boundary was initiated by normal faulting, as it can be seen in interpreted


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seismic profiles (Figures 3.1, 3.2, 3.3). The same is valid for the continental margin segment including the

offshore Krishna-Godavari and Mahanadi basins, as it can be seen in seismic profiles in Krishna-Godavari

basin and in ION- GXT line 1200 (Figures 3.12, 3.13, 3.15, 3.17). They were both controlled by parts of

the sub-parallel rift zones, which made it to continental breakup. Considering this typical rifting style and

geometry, both Krishna Godavari and Cauvery basin has been interpreted as extensional margin

segments. Eventually, in these regions the continental breakup is dominantly controlled by normal

faulting, which gradually becomes oblique slip towards the propagating ends. On the same note,

combining the interpretation of crustal architecture and fault geometry, it is inferred that the Pennar-

Palar continental margin segment, located between the Cauvery and Krishna segments indicates

breakup controlled by the large dextral strike-slip fault zone, as it can be seen in profiles as shown in

Figures (3.6, 3.7, 3.8, 3.9).

In spatial sense, it has been inferred that the lengths of normal faults vary between few kilometers to

tens of kilometers particularly at proximal margin. Therefore, it is difficult to connect each discrete

normal fault zone, given a relatively large spacing of studied seismic cross sections. So, based on

analogy, it has been interpreted as segmented arrays of linked, en echelon faults, which control overall

displacement distributions and resulting hanging wall accommodation patterns as shown in the tectonic

element map (Figure 5.1). This interpretation is also in accordance with global observations along major

at passive margins (e.g., Peacock and Sanderson, 1991; Anders and Schlische, 1994; Wu and Bruhn,

1994; McLeod et al., 2002; Peacock, 2002; Gawthorpe et al., 2003; Peacock, 2004)


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Figure 5.1: Tectonic element map of East India passive margin (modified after Nemčok et al., 2012b). The map

shows the tectonic segments of the study area and structural features of the adjacent onshore. The map shows

different rift faults, its trends, the trends of the pre-existing Precambrian structural inheritance. The geometry of

the crustal boundaries and different crustal domains are also provided. There are six major segments: (a) the NE–

SW-trending Cauvery rift zone; (b) the NNW–SSE-trending dextral Coromondal transfer zone; (c) the NE–SW- to

ENE–WSW trending rift units of the Krishna–Godavari rift zone; (d) the NNE–SSW-trending North Vizag transfer

zone between the Krishna–Godavari and Mahanadi rift zones; (e) the NE–SW-trending Mahanadi rift zone; and (f)

the NNW– SSE-trending dextral Konark transfer zone.


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Rift basins in the proximal margin in extensional margin segments like the Krishna-Godavari and Cauvery

basins can be interpreted using typical rift indicators. This includes wedge shaped bodies, which are

truncated and down-thrown against a normal fault. The primary rift bounding faults are detaching at

mid-crustal levels or near top of the lower crust. This type of rifting style typically exhibits decoupled

deformation, where the brittle upper continental crust is separated from the brittle upper continental

mantle by the ductile lower continental crust representing a jelly-sandwich model (Manatschal, 2004;

Husiamans and Beaumont, 2005; Manatschal et al., 2007). The visible brittle deformation is usually

localized within upper crust and the brittle deformation is also localized within the upper continental

mantle, which is not very well imaged in the given arrays of seismic profiles.

Typical rift indicators also include presence of an angular unconformity, which seals the wedge shaped

seismic reflection package on the top. Based on these criteria, several rift blocks have been interpreted

in the available seismic cross sections (Figures 3.1, 3.2, 3.3, 3.9, 3.10, 3.11, 3.12, 3.13, and 3.16). Thus,

the most prominent structures in the proximal East India margin are fault-bounded rift basins. The most

spectacular examples of fault-bounded rift basins can be observed in the profiles in Krishna-Godavari

Basin that dip toward the basin center displaying an assembly of moderately tilted blocks (Figures 3.10,

3.11, 3.12). In the Cauvery Basin the proximal margin rifting has been seen as small rift segments

initiated by normal faulting as shown in interpreted seismic profiles (Figures 3.1, 3.2, 3.3).

Distal parts of the Krishna-Godavari Basin are characterized by NW-SE trending tilted fault blocks of

continental basement, overlain by relatively thin syn-rift and thicker post-rift sequences in comparison

to the proximal margin. The distal margin rift structures inherit the older rift structures of proximal

margins but the faults developed in distal margins are younger fault structures formed by different

extensional processes sets, relatively later. This is due to the fact that the initial subsidence patterns in

proximal margin during the stretching phase are later replaced by uplift during thinning phase. The most

prominent younger-generation rifting structure is represented by low-angle concave up detachment

faults (Manatschal et al., 2007) as shown in Figures 3.15 and 3.17 . Therefore, syn-fill fill in these

accommodation spaces is developed with reduced thickness and not displaying any divergent wedges.

These tilted blocks are underlain by a prominent reflection, which can be correlated to the top of the

continental Moho.

Uneven stretching of brittle upper and ductile lower continental crusts caused variations in the

rheological layering of the various distal margin segments. It is also observed that the thinning of the


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ductile lower crust is much more homogeneous than the thinning of the upper brittle crust. The ductile

lower crust is more-or-less wedging out ocean-wards in a homogeneous manner in comparison to the

brittle upper crust due the necking process (Davis and Kusznir, 2004; Huismans and Beumont, 2005;

Manatschal et al., 2007), which is more inhomogeneous and the thinning is mostly controlled by upper

crustal faulting.

Few profiles (Figures 3.12, 3.15, 3.16, 3.17) in the Krishna-Godavari and Mahanadi basins show that the

lower crust thins down almost in the same place where the brittle upper crust thins down to zero

thickness. However, in some other profiles (Figures 3.5, 3.10) in the Cauvery and Krishna-Godavari

basins, the lower crust pinches out quite far landwards from the point, where the upper continental

crust thins down to zero thickness. This represents a variation of necking geometry along the margin.

The rift faults in the above mentioned case, usually detach over the top of the significantly thinned

lower crust or continental mantle depending on the break-away fault location. This type of rifting style,

where the brittle upper crust is directly underlain by brittle upper continental mantle is displaying a

coupled deformation. Their contact is represented by a detachment fault accommodating the unroofing

of the mantle. A comparison of necking geometry along the margin further indicates a varying degree of

coupling between the upper continental crust and the upper continental mantle. This indicates that

rheology of the lower continental crust varied during the rifting until the continental breakup, which can

be caused by spatially varying lithology, by spatially varying thermal regime or, most likely, by

combination of both.

In the Krishna-Godavari and Mahanadi basin, most of the distal margin syn-rift fills shows a low

frequency and high amplitude seismic character, which is interpreted as volcanic material in the syn-rift

section. Some hydrothermally altered sequences are also interpreted within the distal margin syn-rift

fills. This may be due to the enhanced magmatic activity during the final continental breakup at the end

of hyper-extension, where excess magma is supplied from the infiltrated fertile mantle (Manatschal and

Karner, 2012).

The distal part of the Cauvery Basin is somewhat different in comparison to the Krishna-Godavari Basin.

The southernmost segment of the Cauvery Basin shows similar distal margin architecture as that

described in the Krishna-Godavari basin but the in profiles like ION-GXT 800, where near the ocean-

continent transition large inverted grabens are interpreted with considerably thick sediments and large

graben-bounding faults (Figure 3.5). Further eastward the crust thins down very abruptly to zero near


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the ocean continent transition. This type of crustal transition is notably different in comparison to the

Krishna-Godavari and Mahanadi Basin and other passive margins examples around the world. This

evidence also indicates a propagation of the dextral Coromandal system went past through this distal

located block of the extensional Cauvery Basin. This has been also evident in the tectonic element map

(Figure 5.1)

The distal margin of the Pennar-Palar basin also shows distal margin grabens as being relatively small.

However, the thickness of the syn-rift fills is considerably thicker in comparison to those in the Krishna-

Godavari and Cauvery basins. A typical nature of the syn-rift fill also indicates a rapid subsidence during

a relatively short period of time. The continental breakup is controlled by the dextral strike slip faulting

acting as hard-linkage between the extensional margin segments in north and south (Nemčok et al.,

2013b). Therefore, these basins are interpreted as a network of small pull-apart basins formed in the

distal margin located eastward of the Coromandal dextral strike slip fault, which is acting as the main

break-away fault in this region. It is to be also noted that the width of the distal margin is narrower than

that of the extensional rift margins. It varies between 30 and 40 km in this region compared to 50-70 km

in Krishna-Godavari Basin.

The combination of interpreted upper crustal normal faults with necking geometry of ductile lower crust

indicates the hyper-extensional deformation of the continental margin. It is characterized by the

localized extension of the upper crust during stretching phase followed by the broadly distributed

extension of the lower crust during thinning phase. The process eventually leads to the exhumation of

the upper continental mantle with continued extension and finally continental breakup and initiation of

localized sea-floor spreading centers as a result of passive rise of asthenospheric mantle diapir in the

ultimate stage (Davis and Kusznir, 2004; Manatschal, 2004; Huismans and Beaumont, 2005; Manatschal

et al., 2007).

5.1.1.3 Margin segmentation

The nature of rifting and nature of continental breakup is highly variable along the margin. As discussed

in the pervious section, it depends on both crustal composition and nature of rifting style, which

determine the nature of continental breakup. The rifting vector at the breakup stage with respect to the

future passive margin controls the geometries of the two end-member scenarios for passive margins:

a. Orthogonal margin: Here the rifting vector is roughly perpendicular to the basin controlling and

breakup controlling fault systems, which are characterized by dominant dip–slip fault kinematics


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(Figure 2.12). The Krishna-Godavari Basin and the southernmost Cauvery Basin corresponds to

this type of margin. (Nemčok et al., 2013b),

b. Transform margin: Here the rifting vector is at an acute angle to the basin controlling and

breakup controlling fault systems, which are characterized by dominant strike–slip fault

kinematics (Figure 2.12), which represents the Coromandal segment of East India (Nemčok et

al., 2013b).

This can be again combined with the interpretations as described in the previous sections. It can be

further concluded that while the extensional segments of the Krishna-Godavari and Cauvery Basins are

later hard-linked by the dextral Coromandal strike slip fault, the Krishna-Godavari and Mahanadi Basins

is never hard-linked with such a fault system. The Krishna-Godavari -Mahanadi linkage system cannot be

as well understood as the Coromandal system due to lack of data. This is rather interpreted as a soft-

linked system developed through a ramp-flat-ramp geometry. Now, it is important to understand, what

conditions controlled the different nature of linkage and segmentation of the entire East India margin.

The current scope of the study could not detail this problem; however, it can be attributed to the

rheological and associated thermal changes along the margin during the continental breakup, which is

most likely responsible for such segmentation.

In comparison to normal faults in extensional margin, strike-slip faults in Coromandal segment also have

also slightly concave up geometry in seismic profiles and also detach at top of the ductile lower crust

(Figure 3.6, 3.7, 3.8, 3.9). Majority of profiles indicate this detachment at top of the lower continental

crust in most parts of the continental margin and top of the continental upper mantle in the very distal

margin. The Coromandal dextral strike-slip faults underwent apparently a significant amount of

displacement during rifting and subsequent continental breakup. This is also to be noted that

Coromandal segments obvious parallelism with the pre-existing Proterozoic Cuddapah orogenic system

pattern (Figure 2.7, 2.8, 2.9) indicates its development influenced by the easternmost pre-existing thrust

system. The Coromandal system is not just a transfer strike slip fault but also controls the hyper-

extension of crust leading to continental breakup.

Furthermore, the Coromandal system forms two horse-tail structures at its ends, which are connecting it

with the two rift zones. It is to be noted that the nature of extension is seen in Cauvery Basin in ION-GXT

profile 800 (Figure 3.5), it apparently suggest further extension of Coromandal system into the system of

Cauvery Basin as mentioned in previous section. This is because of the fact that the Coromandal system


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is kinematically linking the extensional Krishna-Godavari in the north and Cauvery Basin in the south via

its northern and southern horsetail. During this linkage it effectively propagates into the existing

Cauvery rift zone. During this process, it was interacting with the relatively older rift related faults.

5.1.1.4. Hyper-extended crustal architecture model

The hyper-extended passive margin model has been apparently applies to extensional margin segments

(Figures 3.1, 3.2, 3.3, 3.5, 3.12, 3.15, 3.16, 3.17) as well as strike-slip segments (Figures 3.6,3.7,3.8,3.9).

This model is similar to models developed in Alps and Iberia-Newfoundland margin (Lavier and

Manatschal, 2006; Manatschal et al., 2007; Péron‐Pinvidic and Manatschal, 2010; Whitmarsh and

Manatschal, 2012). Characteristic profiles through the extensional margin segment are located in

northern (Figure 5.2) and southern parts of the Krishna-Godavari Basin (Figure 5.3). The characteristic

profiles through the strike slip segment cuts through the central part of the Coromandal northern

horsetail structure (Figure 5.4).

The hyper-extended model of the northern Krishna-Godavari basin (Figure 5.2) shows that the proximal

margin is represented by an aborted H block, with basin bounding faults detached on top of the lower

crust and thicker syn-rift sediments. The aborted H block (hanging wall block) is basically a piece of

relatively undeformed continental crust. The H block is represented by an aborted rift system, which did

not manage to reach the continental breakup (Lavier and Manatschal, 2006; Manatschal et al., 2007;

Péron‐Pinvidic and Manatschal, 2010). This block represents the stretching domain.

The distal margin is represented by a residual H block, which in contrary to the aborted H block

managed to reach the continental breakup (Lavier and Manatschal, 2006; Manatschal et al., 2007;

Péron‐Pinvidic and Manatschal, 2010). The residual H block’s counterpart resides at its conjugate

margin. This interpretation was used as a fundamental guiding tool for India-Elan Bank seismic profile

marriage carried out to reconstruct pre-breakup margin architecture. The boundary between proximal

and distal margins is demarcated by a break-away fault, where crust is thinned very rapidly to zero

thickness followed by a transition into zone of exhumed continental mantle (Proto-oceanic crust). The

main characteristics of this domain formed during the thinning phase. The deformation within thinning

domain is mostly of decoupled character with a little amount of coupled deformation in the western

part. The transition between zone of exhumed continental mantle (proto-oceanic crust) and the oceanic

crust is not very well constrained due to enhanced magmatism in this region.


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Figure 5.2: Hyper extended crustal architecture model of the Krishna-Godavari Basin (northern rift zone). The stylized section shows a proximal margin with stretching related deformation represented by an aborted H block. The distal margin with mostly thinning related deformation is represented by a residual H block. Note that the lower crust thins down to zero much more landwards the crustal breakup location and the deformation in the distal margin is coupled compared to decoupled deformation in proximal margin. Further outboard, we see the zone of mantle exhumation and, finally, the oceanic spreading domain.

In the hyper-extended model of the southern Krishna Godavari Basin, the crustal architecture is almost

similar to previous model except the width of distal margin, which is significantly wider. Additionally, the

transition from zone of exhumed continental mantle to oceanic crust is well documented as a result of

apparent lack of magmatism in this region (Figure 5.3).

The third model, representing the central part of the Coromandal sheared margin is somewhat different

from the models of extensional margins (Figure 5.4). Here, the width of the proximal margin is almost

negligible. This margin is characterized by rapid thinning in its distal part towards the zone of exhumed

continental mantle. The distal margin, again, is interpreted as a residual H block. The thicker syn-rift

section in distal margin represents a pull-apart basin associated with strike-slip deformation. Although,

the strike-slip faults in most of the profiles are detached at Moho level. The outer high representing the

limit of zone of exhumed continental mantle is well documented in comparison to the extensional


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margin models. Notably, the deformation pattern in distal margin is quite different from those of

extensional margins due to presence of thicker lower continental crust. Therefore, the deformation has

a decoupled character. The spreading domain representing the normal oceanic crust is, again, well

recognized due to an apparent lack of volcanism in this region.

Figure 5.3: Hyper extended crustal architecture model of the Krishna-Godavari Basin (southern rift zone). The stylized section shows a proximal margin with stretching related deformation represented by an aborted H block. The distal margin with mostly thinning related deformation is represented by a residual H block. Note that the lower crust thins down to zero much more landwards the crustal breakup location and the deformation in the distal margin is coupled compared to decoupled deformation in proximal margin. Further outboard, we see the zone of mantle exhumation and ,finally, the oceanic spreading domain.


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Figure 5.4: Hyper-extended crustal architecture model of the Palar Basin (Central part of the Coromandal strike-slip fault zone). It is to be noted that the faults are probably cutting the entire crust and part of the brittle upper mantle. The break-away fault is showing a large offset and we also see an exhumation domain before spreading domain. Another important observation is that there is not much decoupling in the distal margin. It indicates that the Coromandal segment is a strike-slip system, but it is not a simple straight forward strike-slip fault but it is transtensional and kind of hyper-extended and then probably hard linked one.

5.1.2 Elan Bank

Given the quality and quantity of the seismic and other data on Elan Bank compared to East India the

interpretation is not as well constrained as that in East India. Nevertheless, the basic crustal

architecture, rifting style and nature of the continental breakup can be interpreted relying on gravity

constrains and seismic profile marriage technique. The published velocity structure of Elan Bank and

refraction surveys (Charvis and Opetro, 1999; Borissova, et al., 2002) are also used to constrain the

interpretation of reflection seismic data.


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5.1.2.1 Crustal architecture and rifting style

The profile-1 of the Elan Bank survey cuts the Southern Kerguelen Plateau (Figure 3.19). This profile is

tied to ODP well-1138. It has a good control to tie the basalt top. The interpreted base of the basalt and

top of the oceanic crust are not well constrained although they have been calibrated using refraction

survey (Borissova et al., 2002). The interpreted crustal architecture does not show any typical rift-

related geometry and deformation. Therefore, the crustal architecture looks simpler than that of Elan

Bank.

The crustal architecture of the Elan Bank can be best observed in Profile-2, which is located in its north-

western margin. The overall geometry of the crustal architecture can be characterized by crustal

thinning towards the west. One advantage of this profile is that the interpreted post-rift basalt thickness

is considerably smaller than the thickness interpreted in profile-1. The seismic signatures demarcating

the boundary between upper and lower continental crust are not very well developed. However, this

can be interpreted observing the terminations of the brittle faults at a common detachment level

(Figure 3.20). It is to be noted the rift faults are soling out along a relatively flat detachment surface. This

detachment surface can be a candidate for either top of the lower crust or any mid-crustal detachment

(MCC/LCC in Figure 3.20). However, the detachment surface becomes the top of the interpreted

continental mantle further westward. This can be explained by the coupled deformation near the ocean-

continent transition. Based on this thinning geometry, the Moho or top continental mantle is

interpreted along a prominent reflective surface below the interpreted detachment surface, particularly

towards the east (LCC/CM in Figure 3.20). Although, there is another prominent reflective surface that

can be interpreted below this surface, it can be an alternative candidate for Moho (CM in Figure 3.20).

Both the LCC/CM and CM surface eventually exhume towards the ocean-continent transition. Therefore,

the final Moho interpretation has been further constrained by seismic profile marriage technique, which

is discussed later. However, both alternative interpretations of Moho suggest that the lower crust is

significantly thinner in the eastern part and absent in the western part towards the ocean-continent

transition. The syn-rift sequence is also thinner and bounded by an unconformity on top. The

truncations at the unconformity can be identified very well to represent the syn-rift top surface.

The interpreted crustal architecture in this profile is typical of distal margin. It can be characterized by a

significantly thinner continental crust. A clear break-away fault that demarcates the proximal and distal

margin is absent in this profile. As mentioned earlier, the rift-faults are possibly detaching at the Moho


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westward near the ocean-continent transition. This coupling character can be only observed in the distal

margin. In addition, the large syn-rift grabens with thicker sedimentary fill typical of proximal margin are

not present in this profile. On the contrary, the interpreted small rift structures filled with thinner

sedimentary fill are the indicative of the distal margin architecture. The detachment surface is highly

reflective. Therefore, it is expected that the surface is heavily mineralized due to lot of fluid movement

during late stage of extension. This feature also indicates a typical distal margin behavior. It is to be

noted that the ocean-continent transition is not very clear. However, a presence of an outer high may

indicate the exhumation along the margin. There is no typical seismic signature of oceanic crust

identified in this profile.

Another profile through the north-western margin of the Elan Bank also indicates atypical distal margin

character (Figure 3.21). The overall geometry of the crustal architecture here can be characterized by

crustal thinning towards the west similar to the previous profile. Notably the interpreted basalt

thickness in this profile is again, thinner than that in central Kerguelen Plateau profile (Figure 3.19). This

profile shows a series of distinct small tilted fault blocks with very thin or absent sedimentary fill. Its

brittle rift-faults detach at a flat and reflective detachment surface, which assists with crustal thinning

towards west. Similar to previous profile (Figure 3.20), the detachment surface may correspond to the

top lower crust (LCC/CM) or top of continental mantle (CM). However, it appears from seismic

interpretation that the deformation style shows a coupled deformation character near the ocean-

continent transition towards the west.

Uncertainties of interpretations here have been, again, constrained using seismic profile marriage

technique and gravity curve interpretation, which is discussed later. Any clear presence of break-away

fault demarcating the proximal and distal margin is characteristically absent. There is no clear seismic

evidence of oceanic crust in this profile.

The profile through the south-western margin of the Elan Bank shows serious volcanic additions, as

interpreted in profile-4 (Figure-3.22). The upper crustal thinning is evident in this profile towards south-

east represented by the series of smaller tilted fault blocks. The half-grabens are characterized by very

thin or absent syn-rift sedimentary fill. The rift-bounding faults do not terminate at the top syn-rift

unconformity. They also affect the post-rift sequences. These faults are sealed by the interpreted intra-

basalt reflector (IB1). This feature may indicate some kind of reactivation of older rift faults along the

margin but their propagation is mostly sealed by the Kerguelen plume related volcanic rock. Like in the


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previous profiles, the tilted blocks are detached at a relatively reflective flat detachment surface, which

can be a candidate for top lower crust or Moho (LCC/CM). Another prominent reflector may be a

candidate for top continental mantle (CM). There are some characteristic volcanic mounds with a

possible root observed towards south east in the Enderby Basin. However, oceanic crust of normal

thickness cannot be delineated based on seismic character alone. The interpreted crustal architecture in

this profile indicates a distal margin deformation character affected by volcanic additions.

The profile -5 through the Elan Bank is located at its southern margin (Figures 3.23) and profile-6 is

located within the Elan Bank continental core (Figures 3.24). The profile-6 is also tied with ODP well-

1137. These two profiles intersect to each other. The profile-6 shows characteristic tilted fault block

geometry with relatively thicker syn-rift fill compared to all previous profiles (Figure 3.24). The syn-rift

top can be identified by the presence of a well-developed erosional unconformity (SR in Figure 3.24).

These faults are also detached at shallow detachment level, represented by the middle or even upper

crust level. The nature of sediments, as observed in the ODP-1137, clearly indicates that fluvial

sediments, which occur here were derived from proximal source. All these afore-mentioned

observations suggest that this profile represents the proximal margin of the Elan Bank.

In profile-5 (Figure 3.23), it is clear that the crust thins very rapidly over a short distance. It quickly steps

down from the continental Elan Bank into the oceanic Enderby Basin, across a major break-away fault. It

is to be noted that the break-away fault divides the proximal margin in the east from the distal margin in

the west. The break-away fault is relatively steep. The geometry of the fault suggests that this is a strike-

slip fault. This conclusion can be further supported by the negative flower structures occurring further

ocean-ward from the break-away fault (Figure 3.24). Therefore, the distal part of the southern margin is

characterized by fewer tilted blocks within a broad zone bounded by two prominent strike-slip faults. It

appears that these faults detach at much deeper crustal level, which cannot be determined in this

profile. This is not a pure strike-slip system but it is rather a transtentional system as the continental

crust thins over a distance of 60-80 km in a stepwise manner. Such strike-slip fault system is apparently

similar to the Coromandal transform system identified in East India.

The seismic signatures are not good enough to separate the upper and lower crustal domains. However,

a prominent reflective surface has been identified in the distal margin, which rises up towards the

ocean-continent transition (CM in Figure 3.23). The overall basalt thickness is more comparable to the

profiles through the north-western margin but less comparable to the south-western margin. It is the


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only transect in Elan Bank, where the normal oceanic crust could be interpreted. The top of the oceanic

crust is relatively sub-horizontal. A series of prominent parallel internal reflections are interpreted

within relatively transparent package below the sub-horizontal top of the oceanic crust.

The profile-7 within the survey is the longest profile but it cuts through the Enderby Basin in the west,

then through the entire southern Kerguelen Plateau and, finally, through the Labuan Basin in the

eastern-most part (Figure 3.25). This profile has a very little part cutting across the Elan Bank. The

internal rift structures cannot be identified within the southern Kerguelen Plateau. However, some high-

angle normal faults can be interpreted. The Labuan Basin is characterized as the high-angle normal fault

bounded basin, where the detachment level is much deeper down in the profile. The Labuan Basin has

been identified as a sedimentary basin bounded by high-angle strike-slip faults. The volcanic thickness

here must be quite large although volcanic load cannot be mapped in reflection seismic image.

5.1.2.2 Nature of continental breakup and margin segmentation

The interpretation of these few seismic profiles through the Elan Bank brings out the fact that being a

microcontinent; the Elan Bank shows quite variable nature of continental breakup along its margins. This

observation helps to separate out its different margin segments within a microcontinent.

The north-western margin of Elan Bank shows a typical normal fault-controlled continental breakup,

where the continental crust in distal margin thins down to zero within the zone 50-80 km wide or wider.

The proximal margin is located in the interior of the microcontinent, which is about 70-100 km from the

interpreted distal margin. The mantle exhumation is further indicated by potential field method on top

of the fact that the seismic interpretation indicates the presence of zone of exhumed continental mantle

eastward of the extended continental crust. This makes the north-western margin hyper-extended. It’s

potentially the extension took place in a zone of over 150-160 km wide before mantle exhumation and

the creation of new oceanic crust initiated (Figures 3.20, 3.21).

The south-western margin also exhibits typical normal fault-controlled breakup. Notably, the faults

appear to be reactivated later. However, this reactivation is limited within the basalt sequence erupted

during Kerguelen volcanism. In this area, the width of the distal margin is around 80-90 km or more

before crust thins down to zero (Figure 3.22). Again, presence of the zone of exhumed continental

mantle cannot be established from seismic alone. This has been further constrained by the potential

field data interpretation. Considering the fact that the proximal margin is still in this margin interior, the


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total extended zone probably ranges from 150 to 160 km. This interpretation suggests that the south-

western margin is also hyper-extended and breakup is normal fault-controlled.

The southern margin is notably different as compared to other margin segments in Elan Bank. The

breakup here is typically controlled by strike-slip deformation. Here the crust thins down to zero rapidly

towards the ocean-continent transition (Figures 3.23, 3.24). This transition occurs within a distance of

over 80-100 km, in comparison to 150-160 km at north-western and south-western margins. It is to be

noted that even if the breakup was controlled by strike-slip faulting here, the crust is still extended over

a considerable distance. Therefore, the nature of strike slip is not a pure one but rather transtensional.

Additionally, a presence of normal fault-controlled tilted blocks at proximal margin also suggests that

the strike-slip faulting is acting as a transfer fault between normal fault-controlled rift segments. This

seems to be analogous to the Coromandal strike-slip system in East India. The presence of the zone of

mantle exhumation in the southern margin of Elan Bank is not very clear in seismic. I interpret that

margin is characterized by hyper-extended strike-slip controlled breakup.

Figure 5.5: Hyper-extended crustal architecture model of Elan Bank(North western margin). The proximal margin architecture is not visible. The distal margin is represented by a residual H block. The exhumation domain is identified further outboard western side.


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Summing up the afore-mentioned interpretations, it is apparent that the Elan Bank is continental margin

is hyper-extended and typically segmented. The interpreted hyper-extension model of crustal

architecture also shows several building blocks including residual H block and possible zone of exhumed

continental mantle (Figure 5.5). The north-western and south-western margins are categorized as

extensional margin segment whereas the southern margin is considered to be a of strike-slip margin

based on its crustal architecture, rifting style and nature of continental breakup.

5.2. Potential field data interpretation

5.2.1. Qualitative Interpretation

5.2.1.1 East India

The Free-air gravity map of East India shows a distribution of anomalies representing the crustal

transition along the margin. As described in the section 4.2, the Free-air gravity directly correlates to the

water depth. The map shows a band of sharp negative anomalies at the shelf-slope break, which is

known as the ‘edge-effect’. This ‘edge effect’ is a result of change in gradient at the shelf break as well

as change in water column height. This phenomenon in East India is characterized by a linear gravity

high of about 30 mGal at the shelf break, which is juxtaposed against a linear gravity low with a

minimum value of -70 mGal over the slope. Beyond this, the map shows a relatively rising negative

anomaly towards the oceanic basin (Figure 3.26). The Free-air gravity map of East India does not allow

mapping of any oceanic fracture zones in comparison to the other offshore areas around the world,

notably in South Atlantic (Sandwell and Smith, 1997; Bird, 2001; Muller et al., 2008; Mohriak and Leroy,

2013). This can be explained by a thick sedimentary load provided by the Bengal fan from Tertiary

onwards. Another reason could be the volcanic addition during Kerguelen magmatic activity

immediately after continental breakup altering the original ocean floor configuration. The data,

however, nicely image the 85oE Ridge in the central part of Bay of Bengal as a system of several large

elongated north-south trending disconnected negative anomalies with an average anomaly value of

about -90 mGal. It has to be noted that there are few isolated circular gravity lows of similar values,

which correspond to isolated volcanic loads when correlated with seismic interpretation (Figure 3.26).

The Bouguer anomaly data contains the gravity response of the crust-mantle interface. The Bouguer

gravity data alone cannot indicate the crustal architecture very well because of ultra large wavelength


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regional anomalies, which are generated as a result in variation in Moho geometry over large area. The

Bouguer gravity high pass (10 km high pass) map shows a seaward increase in anomaly values. A sharp

increase can be noticed from near zero to 10-15 mGal close to the expected continent-ocean boundary.

Further eastward of this boundary, a gradual increase in anomaly values is observed (Figure 3.29). The

ocean-continent boundary is interpreted close to the place, where the positive anomaly over

continental side is juxtaposed over the negative anomaly on the oceanic side. The landward part of

these negative anomalies represents the crustal thinning. Here the extension process replaces the

higher density basement rocks with relatively lower density sediments during the stretching and

thinning phases of the crustal extension. A sharp rise in gravity value oceanward may indicate the

exhumation phenomenon, where the unroofed mantle rocks replace the relatively lower density crustal

rocks as a result of hyper-extension. However, a caution should be taken at this point that this cannot be

a unique interpretation. During the exhumation, the serpentinization of mantle peridotite may

significantly decrease the densities over a considerable distance. However, the width and the thickness

of the damage zone cannot be identified without forward modeling. Both Bouguer and Free-air anomaly

maps in the Bay of Bengal indicate the gradual reduction of continental crustal thickness as expected in

a typical passive margin setting.

Unlike free-Air and Bouguer gravity maps (Figures 3.26, 3.29), the isostatic residual gravity map of the

study area (Figure 3.31) allows a bit finer distinction of imaged features. The continental crust is

characterized by a broad zone of positive anomaly ranging from 40 to 80 mGal, showing the extended

nature of continental crust. In this map, the gravity highs represent the basement highs and lows

represent the basins. The continent-ocean boundary can be inferred as coinciding with a positive

anomaly on the continent side adjacent to the negative anomaly on the proto-oceanic crust side. The

85°E Ridge is represented by long wavelength gravity minima. The two larger circular anomalies further

west of the ridge itself can be identified with isolated volcanoes. The oceanic crust values range from

about 80 mGal through about 20 mGal, to, ultimately, zero indicating a homogeneous and equilibrated

nature of the oceanic crust.

5.2.1.2 Elan Bank

Apart from crustal architecture interpretation, the Free-air gravity anomaly distribution (Figure 3.27)

along with bathymetry (Figure 3.28) plays an important role in defining the physiographic boundary of

the Elan Bank. In the Free-air anomaly map, the Elan Bank is shown as an elongated east-west trending


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anomaly. The anomaly ranges from positive anomaly of 96 mGal in the central part and towards the

Enderby basin, to the negative anomaly value reaching up to -79 mGal. It is to be noted that in the Free-

air gravity anomaly map, the Elan Bank is right-laterally displaced, along an interpreted fault in the

central part (Figure 3.27). This may be an effect of the pre-exsisting Precambrian structures present in

the shield region of India-Elan Bank-Antarctica. The southern margin of Elan Bank shows a quite sharp

change in Free-air anomaly in comparison to the northern and western boundaries of the Elan Bank.

This is due to the ‘edge effect’ described earlier. It has been shown that the southern margin has a steep

slope with a typical shelf-slope-basin geometry, which is characteristic of its transform nature.

In Bouguer gravity map (Figure 3.30), the positive anomaly in central part coupled with negative

anomaly in the basin part gives a strong indication of uprising Moho at ocean-continent transition. It is

to be noted that the change in gravity value is quite gradual in the western margin of the Elan Bank in

comparison to its other margins. This may be a response of either exhumation of continental mantle at

the end of hyper-extended margin or thicker volcanic additions during Kerguelen magmatism. The right

lateral displacement of the central Elan Bank has been captured in the Bouguer gravity map.

Seismic-Gravity Correlation

At Elan Bank margins, it has been observed that there is a general seaward increase in gravity anomaly

values. However, there is no edge effect observed along the profiles. In north-western margin profiles,

there is a decrease in anomaly values from 40 mGal to near zero mGal, towards the west near ocean-

continent transition. Beyond this, a gradual increase in gravity values is observed further westward

(Figure 5.6, 5.7). A sudden rise in positive anomaly at the continent-ocean boundary is caused by the

Moho rising up towards the oceanic side as a result of exhumation. During this process, the less dense

crustal rocks were replaced with the higher density mantle material. Here, notably, there are two small

gravity highs at the ocean-continent transition. They represent a replacement of the low-density crustal

material with high-density mantle material. This signature is very common at other passive margins

where continental mantle exhumes near ocean continent transition as a result of hyper-extension

(Meyers et al., 1998; Rosendahl and Groschel-Becker, 1999; Odegard, 2003; Dehler and Welford, 2013;

Nemčok et al., 2013b). Interestingly, the interpreted seismic data and gravity curve match quite well

with each other here. The absence of any ‘edge effect’ also supports the seismic interpretation that

these regions show more of a distal margin character. This also verifies the seismic interpretation in the


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north-western margin, which is interpreted as the hyper-extended margin with a characteristic presence

of the zone of exhumed continental mantle.

Similarly, the south-western margin also shows such decrease in gravity values towards the ocean-

continent transition (Figure 5.8). Again, any characteristic ‘edge effect’ due to slope break is completely

absent in this profile. In case of normal thickness oceanic crust, a steady state rise in gravity anomaly

value towards the younger oceanic basin should be observed (Meyers et al., 1998; Rosendahl and

Groschel-Becker, 1999; Odegard, 2003; Dehler and Welford, 2013; Nemčok et al., 2013b). The gravity

data in this case, however, do not show any such distinctive presence of the normal thickness oceanic

crust. Here there are two very characteristic shallow gravity highs (Figure 5.8). They have been

interpreted as effects of volcanics, as the seismic images suggest that there are considerably thicker

piles of volcanic rocks in this region. In comparison to the north-western margin, the seismic

interpretation and gravity anomaly correlation do not give any typical match to decipher the crustal

architecture in the south-western margin of Elan Bank. However, they support an overall rapid crustal

thinning along the margin from continental domain to oceanic domain.

Figure 5.6: Qualitative seismic-Gravity showing possible crustal architecture in the NW margin of the Elan Bank. The gravity data is Bouguer gravity (Sandwell and Smith, 1997). It shows distal-margin architecture with considerable thinning in NW margin. The interpreted syn-rift sections are under thin volcanic cover. Moho interpretation is uncertain as there are two possible alternatives. No normal thickness oceanic crust signatures are observed in seismic or gravity. The westward gravity rise is possibly related to exhumation related anomaly.


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Figure 5.7: Qualitative-seismic-gravity integrated interpretation showing crustal architecture of the north-western margin of the Elan Bank. The gravity data is Bouguer gravity (Sandwell and Smith, 1997). The interpretation shows distal margin architecture with considerable thinning in NW margin. The interpreted syn-rift sections are under thin volcanic cover. No normal thickness oceanic crust signatures are observed in seismic or gravity. The rise in the gravity in the west is possibly related to exhumation related anomaly.

The gravity anomaly response in strike-slip dominated southern margin of Elan Bank is quite typical in

comparison with other strike-slip dominated margins (Figure 5.9). This margin is characterized by a

sharp decrease in anomaly from 75 mGal to -40 mGal from shelf to slope. This represents a typical ‘edge

effect’. This decrease in anomaly values as a result of ‘edge effect’ is followed by a steady-state rise up

to 20 mGal towards the oceanic region (Figure 5.9). This is, again, a common phenomenon as a result of

rise of Moho towards ocean as described earlier. Notably, there are no intermittent gravity highs. The

absence of such gravity highs are indicative of no or insignificant mantle exhumation along the margin.

This is fairly common for strike-slip margins. The gravity and seismic interpretations here can be

correlated almost one-to-one and provide a good chance to interpret the crustal architecture of the

southern margin of Elan Bank.


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Figure 5.8: Qualitative-seismic-gravity integrated interpretation in south-western margin of the Elan Bank. The gravity data is Bouguer gravity (Sandwell and Smith, 1997) This is not a dip-oriented profile. The COB region is volcanically thickened due to hotspot activity. The shallow gravity anomaly may be related to volcanic mounds. The COB transition is not very clear. No normal thickness oceanic crust signatures are observed in seismic or gravity.

Figure 5.9: Qualitative-seismic-gravity integrated interpretation showing proximal to distal margin architecture in southern margin of the Elan Bank. The gravity data is Bouguer gravity (Sandwell and Smith, 1997). The gravity anomaly looks like shallow anomaly may be related to volcanic mound. This is typically a strike-slip margin characterized by large offset faults and flower structures. The strike-slip but is transtensional. The gravity is also showing the edge effect and then rising steadily due to Moho uprising. No exhumation related anomaly is present in the profile. The inflection point indicated in this figure is the interpreted location of COB.


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5.2.2 Quantitative Interpretation

The quantitative interpretation was carried out in East India to validate the crustal architecture

interpreted in seismic. However, the same cannot be carried out on Elan Bank. The lack of depth

imaging on Elan Bank does not allow a detailed forward modeling.

5.2.2.1 East India

The first model was created exactly based on seismic interpretation of ION-GXT profile-1000. This profile

exhibits crustal architecture best, showing a distribution of upper and lower continental crusts,

continental mantle, oceanic crust and their respective boundaries. This section shows hyper-extended

character of the continental margin with a presence of a zone of exhumed continental mantle (proto-

oceanic crust) between the oceanic and extended continental crust (Figure 5.10). Such exhumed mantle

is usually composed of serpentinized peridotite, lava flows and occasional presence of continental

allochthons (Meyers et al., 1998; Odegard, 2003; Nemčok et al., 2013b). It can be noticed that both the

observed and calculated gravity anomaly value match well in this model. The gravity high immediately

near the ocean-continent transition can only be explained by the presence of a high-density material

equivalent to density between 2.99 and 3.13 gm/cc. These densities match well with observed densities

of serpentinized peridotite mantle rocks as observed other regions of the world (Sibuet et al., 2007).

This scenario further illustrates a need for assuming serpentinized blocks in several areas of the proto-

oceanic crust. The serpentinized mantle (proto-oceanic crust) is denser than the upper and lower

continental crusts, and the normal oceanic crust. The highest density values are very close to the

densities of peridotite. The proto-oceanic crust density distribution shows an oceanward decrease. The

oceanic crust area shows a three-layer density variation. This usually corresponds to tripartite sequence

of gabbro layer, sheared dyke layer complexes and a layer composed of basaltic lava flows mixed with

sediments. The interpreted upper and lower continental crust also shows typical density distribution as

described in Table 4.1.


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Figure 5.10: The gravity and magnetic models along the reflection seismic section ION-GXT-1000 showing a good match between observed and calculated gravity and magnetic data (Bird 2009). The thick and thin lines show the observed and calculated curves, respectively. The profile shows a high-density body, with the density of 3.13 g/cc, as situated just east of the extended continental crust, indicating unroofed mantle. The areas with a density lower than the continental mantle but higher than the normal oceanic crust occur towards the eastern side of the proto-oceanic crust, which can be attributed to serpentinization of the unroofed mantle. A strong magnetic anomaly is present over the proto-oceanic crust zone.

The alternate model was a modification of the previous model. It was generated by incorporating the

distribution of continental and oceanic crust only. This model without zone of exhumed continental

mantle creates a significant gravity mismatch of about -25 mGal (Figure 5.11). This mismatch cannot be

reduced by any alternate geological interpretation of the modeled architecture (Nemčok et al., 2013b).


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Figure 5.11: Gravity-magnetic modeling along Profile IONGXT-1000 (without POC). It is same as shown in Figure 5.10. This time showing a continental–oceanic transition model without proto-oceanic crust (Bird 2009). It shows a significant observed–modelled gravity curve mismatch at the continental–oceanic transition, indicating that the match can be achieved only with a model containing the high-density material of the proto-oceanic crust formed by unroofed mantle.

The presence of a corridor of exhumed continental mantle at the ocean-continent transition can be

verified through the correlation between observed and calculated gravity and magnetic anomalies along

profiles ION-GXT-1200 and 1600 (Figures 5.12, 5.13). The geometries of modeled bodies were not

altered too much from their initial definition taken from the interpreted seismic image. However, the

forward modeling results along these profiles indicate that the alternate model (without presence of

exhumed mantle) is fairly acceptable (Figures 5.14, 5.15). However, the mismatch between observed

and calculated gravity can be still observed in the alternate model of ION-GXT-1200 (Figure 5.14). Here,

the match can only be achieved if a very thin continental crust is introduced near the ocean-continent

transition. As a result, the mantle reaches almost the surface to reach the match. This is nothing but a

variation of hyper-extension model with minor density tweaking and disregarding the seismic

observation. In profile ION-GXT-1600, the alternate interpretation is valid only on condition of heavily

underplated oceanic crust introduced in the model (Figure 5.15). This option may be valid as the profile

is in close proximity to 85°E Ridge, which may contribute to the underplating phenomenon. However,


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this option is negated based on timing argument discussed in later sections. A similar interpretation

indicating underplating is also valid for ION-GXT profile-1290 (Figure 5.16).

Hence, it can be seen that both the alternate situations are more or less possible in East India along all

selected forward modeling profiles except ION-GXT profile-1000. However, if seismic interpretation and

forward modeling results are integrated, there are no geological cross sections, which logically disprove

a presence of zone of the exhumed mantle or proto-oceanic crust corridor in East India. On the contrary,

the mismatch is quite severe in profile ION-GXT -1000, where a presence of exhumed mantle is required.

Additionally, this profile is free of any later volcanic addition in East India, whereas all the other profiles

are through the areas significantly affected by Kerguelen volcanism. Similar situation also can be

observed for ION-GXT profile-1200. Therefore, the non-uniqueness of modeling results can be explained

by the volcanic addition. Based on arguments discussed above, it is quite logical and geologically

plausible to incorporate a presence of a corridor of exhumed continental mantle along East India

margin. This phenomenon also directly substantiates the hyper-extended crustal architecture model in

East India.

Figure 5.12: Gravity-magnetic modeling along Profile IONGXT-1200 (Bird, 2009). High density crustal blocks, 2.99 and 3.13 g/cc, which are located between oceanic and continental crust are thought to be unroofed and hydrated upper mantle peridotites (serpentinized upper mantle = proto-oceanic crust). The 2.63 g/cc source body located at the outboard end of the seismic line, between basement and the deepest sedimentary layer, is thought to be the volcanic 85°East Ridge.


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Figure 5.13: Gravity-magnetic modeling along Profile IONGXT-1600 (Bird, 2009). The slightly higher density crustal block, 2.9 g/cc, located from 500 to 590 km along the model is thought to be unroofed and hydrated upper mantle peridotite (serpentinized upper mantle = proto-oceanic crust). The volcanic 85°East Ridge may extend into the same location as the interpreted proto-oceanic crust.

Figure 5.14: Gravity-magnetic modeling along Profile IE-1200 (without POC) (Bird, 2009). Constant oceanic layers require a thin transition between oceanic and continental crust. The 2.63 g/cc source body located at the outboard end of the profile, between basement and the deepest sedimentary layer, is thought to be the volcanic 85°East Ridge. The crustal architecture supports both exhumation and no exhumation scenarios.


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Figure 5.15: Gravity-magnetic modeling along Profile IE-1600 (without POC) (Bird, 2009). High density crustal blocks, representing unroofed and hydrated upper mantle peridotite may not exist along this section of the margin. This because of the transition from continental to oceanic crust does not require anomalously thin crust (see Models 1000 and 1200). Modeling a transition from continental to oceanic crust, without incorporating proto-oceanic crust, is possible along this transect. 85°E Ridge is close tothe margin, suggesting that this area may not be magma-starved.

Figure 5.16: Gravity-magnetic modeling along Profile IONGXT-1290 (Bird, 2009). The thick and thin lines show the observed and calculated curves, respectively. The profile shows high-density bodies with densities are situated just SE of the extended continental crust, indicating unroofed mantle or proto-oceanic crust. Both blocks have a density lower than the continental mantle but higher than the normal oceanic crust located further to the SE. This can be attributed to the serpentinization of the unroofed mantle. A strong magnetic anomaly is present over this zone.


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5.3 Interpretation of Petrological Data

5.3.1 Textural and compositional analysis

It has been found through several studies (Nicolaysen et al., 2001; Weis et. al, 2001; Ingle, 2002) that the

conglomerate clasts recovered from fluvial sequence in drilled well ODP-1137 in Elan Bank, have a

variable lithology including alkali basalt, rhyolite, granite-gneiss and other metamorphic rocks (Figure

3.37). The seismic and well interpretation suggests that the location of this site represents a proximal

margin setting just above the interpreted syn-rift unconformity (Figure 5.17). The lithogical and clast size

variation invariably depicts that it is polymictic extra-formational conglomerate. The depositional

environment analysis suggests a near source fluvial system, in which these conglomerates were

deposited.

Figure 5.17: High resolution seismic-geological interpretation of ODP-Well -1137 site on the Profile-6 at Elan Bank. The well is located in a proximal margin setting in the central part of Elan Bank. The conglomeratic litho-units are above the syn-rift unconformity sandwiched between two thin layers of basalt units. The top massive basalt does not show any characteristic reflection pattern. While the bottom one showing some reflectivity. The reflectivity may be due to brecciated nature of basalt as well data suggests. The seismic section just above the syn-rift unconformity appears to be sedimentary in nature.


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However, a presence of fluvial system in an oceanic domain is rather rare. Additionally, these types of

conglomerates are typical product of tectonically active continental crustal source in close proximity. As

a result, the clasts formed by typical basement lithology, have been eroded from active region and

transported by a braided river system over a short distance. Hence, it can be interpreted that the

uplifted footwall blocks are acting as the primary provenance, which are nothing but Precambrian

basement that has been involved in India-Elan Bank-Antarctica rifting. Although, the short distance

fluvial systems are possible in a sub-aerial environment in long-lived LIPs like Kerguelen Plateau as a

result of dynamic uplift. However, in such a case, the extra-formational conglomerate with presence of

clasts with variable lithology like granitioids, gneisses and metamorphic rocks of Precambrian age is

difficult to explain. Additionally, the unaltered granite clast indisputably indicates a nearby presence of a

continent. Therefore, a presence of continental crust and its vicinity to East India-Antarctica

Precambrian province can be ascertained.

5.3.2 Geochemical analysis

The geochemical analyses including major element chemistry, trace element geochemistry and rare

earth element analysis of the Elan Bank rocks demonstrate volcanic enrichment, which is sometimes

post-Albian (Coffin et al., 2002; Ingle et al., 2002). While the volcanic clasts including alkali basalt, and

trachyte show a high level of enrichment, the non-volcanic clasts show a relatively lesser enrichment in

mantle normalized incompatible trace element variation diagrams (Ingle et al., 2002) (Figure 5.18). This

strongly indicates a common provenance for all clasts.

Several alternate hypotheses can explain these trends. However, it is most likely that Kerguelen plume

magmas were contaminated by continental crustal rocks due to partial melting of hotter continental

crust. This phenomenon empirically directs towards a condition where continental crust is not only

stretched but considerably thinned. The thinning due to extreme extension leads to initiation of the

partial melting in the crust. A partial melting of crust during advanced stage of rifting is often associated

with magmatism, which contains a crustal contamination (Manatschal and Karner, 2012). This also

indirectly points towards the hyper-extended crust at Elan Bank, which was in place during the India-

Elan Bank-Antarctica rifting.


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Figure 5.18: Multi element variation diagram for volcanic clast (top) and non-volcanic clast (bottom) normalized to primitive mantle composition (Ingle et al., 2002). The average composition of upper and lower continental crust are shown for reference.

5.3.3 Isotopic and radiometric analysis

Isotopic analysis was carried out to compare the granite gneiss clasts, sampled from the Eastern Ghats,

Indian craton, East Indian shield region and Elan Bank (Li and Powell, 2001). It suggests a correlation

between India and Elan Bank. The felsic volcanic clasts from the Elan Bank have intermediate isotopic

characteristics, which indicate a continental source (Ingle et al., 2002). Ages of depleted mantle from

samples taken from Elan Bank, Antarctica and East India (Nicolaysen et al., 2001; Rickers et al.,, 2001;

Weis et. al, 2001) indicate the addition of juvenile mantle-derived material during remobilization and

reworking of the Precambrian continental crust, mostly during Neoproterozoic Pan-African orogeny.

Notably, the Eastern Ghats mobile belt is formed during Pan-African orogeny (Naqvi and Rogers, 1987).

Mentioned data prove a similarity of crustal types present through all these continental masses. The

Eastern Ghats of India and Rayner complex of East Antarctica show a direct isotopic correlation. Based

on the isotopic composition, the entire correlated belt can be sub-divided into four separate domains

(Rickers et al., 2001) (Figures 5.19). It can be seen in the diagram that the isotopic composition of Elan


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Bank samples falls into domain-2 field, which represents mostly the Eastern Ghats mobile belt domain in

East India. Interestingly, the basement map of East India (Figure 2.8) shows that the Eastern Ghat mobile

belt, dominantly composed of high grade metamorphic rocks, is located in the region adjacent to the

western portion of the Krishna-Godavari Basin. It also proves by samples from wells that these rocks

primarily form the basement in offshore Krishna-Godavari Basin (Bastia, 2009). This forms another

interesting evidence for proximity of Elan Bank to the Krishna-Godavari basin.

Figure 5.19: Isotopic analysis to compare the granite clasts, sampled from Eastern Ghats, Indian carton and Indian shield region. The felsic volcanic clasts from Elan bank have intermediate isotopic characteristics for a continental source (ingle et al., 2002). (a) Plate reconstruction at 130 Ma, just before early Kerguelen plume activity (Li & Powell, 2001), (b), εNd (T) vs (87Sr/86Sr) expanded to compare the gneiss clasts and crustal reservoirs of Eastern Ghats and Indian cratons (Rickers et al., 2001). The sample analysis by Ingle et al., (2002) indicates the clasts have intermediate isotopic characteristics between the gneisses and alkali basalts and could represent either a mixture between mantle derived magma and the gneiss or partial melts of a younger, less evolved continental crustal source. (c) Variation of 207Pb/204Pb vs 206Pb/204Pb with superimposed plumotectonics curves for the upper and lower crust (Ingle et al., 2002). Gneiss samples are plotted against age corrected to 109 Ma, which is volcanism age at site 1137, Coffin et al., 2002 and 550, which is age of thermal metamorphic overprint (Nicolayesen et al., 2001).


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The sample analysis from ODP well-1137 in Elan Bank indicates that the gneisses and sandstone and

conglomerate matrix contain zircons and monazites with Neoproterozoic and Archean ages (Weis et. al,

2001). The same age with similar error margin characterizes its East Indian counterpart (Rickers et al.,

2001; Chetty et al., 2003)

The interpretation of discussed petrological data together with the data produced in this study leads to

following conclusions:

1. Presence of continental crust in Elan Bank;

2. The age of the Elan Bank conglomerate clasts ranges from Neoproterozoic to Archean;

3. The continental crust in Elan Bank is considerably stretched and thinned as a result of hyper-

extension;

4. The textural and compositional nature of the conglomerates indicates a nearby source

represented tectonically active area; i.e. rift shoulder;

5. The crustal contamination signatures of magma indicate a time delay between hotspot activity

and lithospheric extension, and

6. The provenance of granite clasts indicates proximity of Elan Bank to Krishna-Godavari Basin and

Eastern Ghats of East India.

5.4 Tectonic reconstruction of Elan Bank and East India

As described in Chapter-4, the reconstruction between India and Elan Bank did not follow any

conventional global plate reconstruction methods. It is rather a derivative of the global model using the

interpretative results from regional datasets as its constraints. The global plate reconstruction model

(Reeves, 2008) has been taken as a basic tool, to which several regional constraints have been applied

with. This resultant model, thus, provides more detailed regional model.

The initial geometric fit was achieved using the global plate reconstruction models by Reeves, (2008)

and Műller et al., (2008). It includes the magnetic anomaly track lines and the rotation poles, which were

used therein. The initial fit places the Elan Bank in front of the Krishna-Godavari margin segment (Figure

5.20). In this tight geometric fit, the north-west margin of Elan Bank ties against the southernmost part

of the Krishna-Godavari Basin. As a result, the western margin of Elan Bank matches with the dextral

Coromandal accommodation zone, which is roughly striking north-south segment of the East Indian

margin. The south-eastern and southern margins of Elan Bank form a roughly NW-SE trend, which


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matches the eastward continuation of the Cauvery rift zone. It is to be noted that the presence day

southern margin is a conjugate to the East Antarctica margin. The northern side of the Elan Bank

matches with the orthogonally rifted Krishna-Godavari Basin. The eastern margin of the Elan Bank

roughly matches with the north-Vizag transfer zone (Figure 5.1) connecting Krishna-Godavari and

Mahanadi basins. It is to be noted that the easternmost boundary of Elan Bank is limited by the NW-SE

trending pre-existing Precambrian shear zone, which is known as Nagavali Shear Zone (Chetty et al.,

2003) (Figure 5.1).

Figure 5.20: India–Elan Bank reconstruction. In this reconstruction, the north-west margin of Elan Bank ties against southern most part of Krishna-Godavari basin. As a result, the western margin of Elan Bank then matches with the East India segment, which is roughly north-south striking dextral Coromandal accommodation zone. The south-eastern margin and southern margin of Elan Bank fits in a trend along the eastward continuation of the Cauvery rift zone. It is to be noted that, the southern margin in present day is a conjugate to East Antarctica margin. The northern side of the Elan Bank matches against the orthogonally rifted Krishna-Godavari basin. The south-east side of the Elan Bank roughly matches with the north-Vizag transfer zone connecting Krishna-Godavari and Mahanadi basin. The shape of the Elan Bank is modified after Borissova et al., (2003).


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Described initial fit is not accurate because it is a pure geometric fit that has been achieved through

rotation pole calculation and magnetic track lines fitting the isobaths. Therefore, this fit does not

account for any hyper-extended crustal architecture at conjugate margins. Hence, this model is further

constrained by interpretations of large amount of seismic data in East India as well as limited amount in

Elan Bank. It can be concluded, after interpretation of East India and Elan Bank data that both margins

were hyper-extended before the continental breakup was reached. The seismic and potential field data

interpretation also proves the mantle exhumation phenomenon on East India.

The regional constraints for plate reconstruction include:

a. Structural architecture of different margin segments;

b. Crustal architecture of margin segments and geometry of crustal boundaries;

c. Cross verification with seismic marriage using 2D seismic lines and

d. Determination of the tectonic timing using borehole data (biostratigraphy, sedimentology

etc.).

5.4.1 Structural architecture of different margin segments

Interpreted normal fault systems characterize extensional rift segments represented by the Cauvery,

Krishna-Godavari and Mahanadi rift zones, roughly trending in NE-SW direction. On the contrary, the

Coromandal segment is characterized by the strike-slip transfer system, hard-linking the Krishna-

Godavari and Cauvery rift zones. The north Vizag transfer zone is not a transform but rather a

gradational ramp-flat-ramp transition zone. Apart from fault geometries imaged by reflection seismic

data, the nature of rifting at different margin segments is also indicated by calculated paleo-stress

tensors from meso-scale fault striae and joints data from outcrops along the East India passive margin

(Shah et al., 2007, Srivastava and Shah, 2007) (Figure 5.21).

The seismic images of north-western and south western margins of Elan Bank indicate normal fault

controlled breakup (Figure 3.20, 3.21) while the southern margin displays the strike-slip controlled

breakup (Figure 3.23). Unfortunately, the nature of northern margin cannot be inferred due to a lack of

seismic coverage.


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It is important for this reconstruction that the rifting style matches on both the conjugate margins. The

north-western and northern margins can fit against the Krishna-Godavari, Cauvery and Mahanadi basins

(Figure 5.20).

Figure 5.21: The calculated paleo-stress analysis from meso-scale fracture data from outcrops all along the East India passive margin (modified after Shah and Srivastava, 2007)

It has to be noted that as we move south, the Krishna-Godavari margin gradually becomes oblique

before linking with Coromandal transfer zone (Figure 2.12). A similar situation exists in Elan Bank where

the north-western margin becomes more oblique down south (Figure 3.21, 3.22). Therefore, it is more

likely that Elan Bank fit is close towards the propagating end of extensional margins in East India, where


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there is a prominent oblique rift signature. The interpretation in northern margin is clearly uncertain.

This leads to a broad range of possible correlation between two conjugate margins.

Based on rifting style, the initial geometric fit can be constrained with two reasonable options to place

Elan Bank against East India. The first option is to fit Elan Bank against the Mahanadi Basin bounded by

the North-Vizag transfer zone in the south and the Konark transfer zone in the north. The second option

would be to place it against the Krishna-Godavari Basin bounded by the Coromandal transfer zone in the

south and north-Vizag transfer zone in the north. Simply based on rifting style, it is also possible to place

Elan Bank against the Cauvery Basin. However, the presence of the Sri-Lankan continental block will

create a problematic overlap. Therefore, this option is not considered. The India-Elan Bank fit using

rifting style is further constrained by margin segment architecture and geometry of crustal boundaries.

5.4.2 Crustal architecture of margin segments and geometry of crustal boundaries

The interpreted geometry of crustal boundaries in East India gives a clue to India-Elan Bank fit. The

easternmost limit of extended continental crust and westernmost limit of normal oceanic crust are

represented by continental-proto-oceanic crust and proto-oceanic crust-oceanic crust boundaries

respectively (Figure 5.22). The zone bounded by these two boundaries represents the proto-oceanic

crust corridor in East India. The along-strike variation of the width of this zone is strongly linked with the

variations in kinematic control of the continental breakup at different segments of the margin (Figure

5.22). This variation is also directly linked to variable rifting style prior to continental breakup along the

future passive margin. The fit between crustal boundaries is a good constraint to fit the hyper-extended

conjugate margins. This is crucial information in this study, which provides more local control in the

reconstruction between India and Elan Bank.

It is to be noted here that the original shape of Elan Bank was taken from Borissova et al., (2003).

However, to fit Elan Bank in this reconstruction, the original Elan Bank continental mass has been re-

shaped based on seismic and gravity interpretation (Figure 5.23). The reinterpreted shape of Elan Bank

represents the continental crustal domain. The other domains interpreted in Elan Bank, using seismic

and gravity interpretation, include the zone of exhumed continental mantle and oceanic crust domain.

The ocean-continent boundary in Elan Bank is also interpreted based on above mentioned constraints in

different segments of Elan Bank margins (Figure 5.23). It is to be noted that the shape is well

constrained at north-western, south-western and southern margins, but the northern and eastern

margins remain almost the same as those interpreted by Coffin et al., (2002) and Borissova et al., (2003).


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Figure 5.22: The East Indian margin crustal boundary map showing the extent of continental, proto-oceanic and oceanic crusts and their boundaries overlaying the isostatic residual anomaly map (Nemčok et al., 2012b). Note that the width of the proto-oceanic crust varies along the margin. The width depends on the breakup mechanism and extension rate. It is narrow in the strike-slip-controlled margin segments and wide in the normal fault-controlled segments. See the caption to Figure 3 for the six margin segments.

The crustal boundary map prepared for East India shows a clear variation of proto-oceanic crust corridor

along its margin (Nemčok et al., 2013b) (Figure 5.22). This corridor is widest close to the extensional

margin segments, having an average width of 85 km. It is less wide next to oblique segments and

narrower next to the sheared margin segment (Table 5.1)s. The proto-oceanic crust corridor is the


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narrowest next to the strike-slip dominated Coromandal sheared segment having an average width of

50 km.

A similar function characterized the Elan Bank. The north-western margin has a widest proto-oceanic

corridor in front of it with an average width of 70 km. The proto-oceanic corridor is narrower next to the

obliquely rifted south-western margin and almost absent next to the strike-slip dominated southern

margin. Apart from the rifting style, the fit between India and Elan Bank can be constrained by the

measured width of the proto-oceanic corridor. This justifies the placement of the Elan Bank against the

Krishna-Godavari and Cauvery rift zones in a more logical manner. However, in this placement, the

timing of separation cannot be well resolved.

Figure 5.23: Tectonic element map of Elan Bank prepared by summarizing all the observations and interpretation. The Free-air gravity anomaly map (Sandwell and Smith, 1997) is the backdrop. The distribution of the proximal, distal and proto-oceanic or exhumation and oceanic domains are demarcated. The continental-proto-oceanic boundary is only present in the north-western margin. The proto-oceanic-oceanic boundary or simply continent-ocean boundary (COB) is interpreted in north-western, south-western and southern margins. Different margin segments show the geometry and width of the different crustal architectural elements and their variation in respective margins. Note that the thicker volcanic cover is only present in the SW margin.


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The maximally extended crust characterized in the extensionally rifted Krishna-Godavari Basin (Figure

5.1). Its width from the onshore basin margin fault to proto-oceanic –oceanic crust boundary is about

220 km (Table 5.1). The same measurement cannot be as accurate for Elan Bank conjugate due to

limited imaging of the proximal margin architecture. The best estimate of the width of the distal margin

one can make from the images is about 65 km in north-western margin. If it is extrapolated to

continental core of Elan Bank, it reaches around 200-250 km (Table 5.2). These estimates indicate a

highly extended crust in north-western margin of Elan Bank, which makes it similar to the Krishna-

Godavari margin. This strongly advocates the placement of the Elan Bank against the Krishna-Godavari

Basin. However, this placement may be inaccurate due to imperfect restoration of the strong

extensional asymmetry along this margin. Despite of all above mentioned inaccuracy, it can be

concluded that the Elan Bank crustal block shifts the oldest continent boundary southwards about 200-

250 km when fitted into the area between the Krishna-Godavari and Cauvery basins.

Indian margin segments Proximal margin

width

Distal margin

width

ZECM

width

Cauvery rift zone 80-100 40-50 60-70

Coromandal strike-slip fault zone 15-20 25-40 40-70

Krishna-Godavari rift zone 50-100 40-70 70-100

North Vizag transfer zone 30-40 30-50 40-60

Mahanadi rift zone 60-100 60-80 60-80

Table 5.1: Table showing the width of the different interpreted crustal domains in East India from integrated seismic interpretation and potential field modeling. Measurements are in km.

It is also possible that several other small continental pieces could have been released along with Elan

Bank during its separation from East India. These unaccounted continental pieces may remain as

unidentified microcontinents, somewhere in the Indian Ocean or Kerguelen Plateau itself (see also

Borissova et al., 2002). It would take to identify them and piece them together with Elan Bank and India


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to reconstruct a precise initial match. This reconstruction, therefore, leaves few gaps in current model.

These gaps are potential places, where such unidentified microcontinents can be accounted for.

Elan Bank margin segments Proximal margin

width

Distal margin

width

ZECM

width

North-western margin ? 65-70 60-80

South-western margin ? 65-70 20-25

Southern margin 40 35-40 ?

Table 5.2: Table showing the width of the different interpreted crustal domains in Elan Bank from integrated seismic interpretation and potential field modeling. Measurements are in km.

5.5 Seismic profile marriages

The reconstruction described in the previous sections, allows interpreting that the NW margin of Elan

Bank is a possible fit to the southern part of Krishna-Godavari Basin close to the northern horsetail

structure of the Coromandal strike slip system (Figure 5.22). This fit allowed to select two profiles in East

India to be paired with two profiles through the north-western margin in Elan Bank. Due to limited data

on Elan Bank, the Profile-2 of Elan Bank margin has been paired with several interpreted profiles in East

India margin based on the initial India-Elan Bank fit.

Table 3.1 and 3.2 shows that the final continental breakup took place and continents were moved away

from each other, the subsidence and depositional histories during continental drift on either conjugate

margin became different. The youngest sequence, which can be mapped across both conjugate margins,

is the interpreted horizon K30 being close to Aptian-Barremian boundary and roughly indicating the

initiation of sea floor spreading between India and Elan Bank.

The first seismic marriage is done between profile-2 through the Elan Bank (Figure 3.20) and a

composite profile through the East India consists of ION-GXT profile-1000 (Figure 3.15) and two profiles

from Reliance seismic data repository (Figure 5.24). This is the closest pairing possible for this part of

margin. This reconstruction was carried out for the time just prior to sea-floor spreading initiation

between India and Elan Bank, which is close to K30 age (Figure 5.25).


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It can be observed that the deeper reflector initially mapped as CM in Elan Bank (Figure 3.20) does not

match with the well constrained Moho interpretation in East India. It is rather the shallower reflection

LCC/CM, which is a close match for East Indian Moho interpretation. This allowed correcting and

constraining the interpretation of Elan Bank image.

Figure 5.24: Map showing the locations of the India –Elan Bank conjugate margin paired seismic sections. Transect (A-A’) is created in East India side usingIONGXT-1000, 4000 and Profile-2 of Elan Bank seismic survey. The transect B-B’ is created by using Reliance seismic line, IONGXT-4000 and Line-3 of Elan Bank seismic survey. Note that there is a small gap in transect B-B’. These transects have been used for seismic profile marriage between India-Elan Bank conjugate margin. The proto-oceanic boundaries of India and Elan Bank have used as a constraint. The shape of the Elan Bank has been modified after Borissova et al., (2003).

Another pairing was performed between Elan Bank profile-3 with two other Reliance profiles in East

India (Figures 5.26). The gap between these profiles represents some amount of uncertainty in the

interpretation. These paired seismic profiles also indicate that the interpreted continental mantle (CM)


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in Elan Bank does not match well with well constrained Moho interpretation in East India. Therefore,

this paring, again, allows reinterpreting and adjusting the Moho depth on Elan Bank. In this case, the

Moho also matches with the interpreted horizon (LCC/CM) in Elan Bank.

Figure 5.25: India –Elan Bank reconstruction at 123 Ma (Transect A-A’). This reconstruction was made by pairing seismic sections on conjugate margins. It also helps to fix the Moho interpretation in Elan Bank. The reconstruction is limited by Aptian (123 Ma) only as this is the time when the final continental breakup occurred between East India and Elan Bank. Note that the width of the distal margin in the Elan Bank is wider than its conjugate East India. The nature of extension appears to be asymmetric as there is a distinct variation in width of the coupled and decoupled domains in the distal margin. The geometry of the exhumation fault suggests that the East India margin is located on the hanging wall side, i.e. upper-plate, while the Elan Bank is located on the foot-wall side, i.e. lower-plate.

As discussed in the previous sections, the East India margin architecture clearly shows a hyper-extended

crustal architecture in both proximal and distal margins. These include aborted H block, residual H block,

and zone of exhumed continental mantle. The interpreted faults are shown in Figures 5.25, 5.26

different colours indicating stretching, thinning and exhumation related origins. The different

extensional domains including coupled and decoupled deformation are also interpreted based on

geometry of detachment faults as well as interpreted upper crust, lower crust, inherited and infiltrated

sub-continental mantle components (Figures 5.25, 5.26). It is evident that the deformation is decoupled


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in proximal margin. On the contrary, the deformation is coupled in distal margin as a result of necking

and complete removal of the lower crust during hyper-extension.

Figure 5.26: India –Elan Bank reconstruction at 123 Ma (B-B’). This reconstruction was made by pairing seismic sections on conjugate margins. Similar previous reconstruction (Figure 5.25), the reconstruction is limited by Aptian only as this is the time when the final continental breakup occurred between India and Elan Bank. The width of the distal margin is more in East India in comparison to the Elan Bank. Here also the nature of extension appears to be asymmetric of the coupled and decoupled domains in the distal margin. The geometry of the exhumation fault suggests that the East India margin is located on the hanging wall side, i.e. upper-plate, while the Elan Bank is located on the foot-wall side, i.e. lower-plate. Note that the widths of the coupled domain in this transect which less than as shown in Figure 5.25. The coupling width gradually becomes lesser as it approaches towards the Coromandal strike slip margin down south.

Available imagery of the Elan Bank margin mostly displays distal margin architecture. The crust is very

thin and the syn-rift sedimentary fill is quite thin. The rift faults do not have much throw. Elan Bank

margin shows a very thin layer of lower crust, which is present in the west. It pinches out before it

reaches the ocean-continent transition further eastward. It has to be noted that the width of the

coupled domain also varies between conjugate margins, where the width of the decoupled domain is

almost similar (Figure 5.25). The width of decoupled domain in Elan Bank is wider than that of the

conjugate East India in other paired profile (Figure 5.26).


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The widths of interpreted crustal mass (Table 5.2) and conjugate margin crustal architectures in seismic

pairs also indicate that the East Indian margin represents the upper plate. This indicates that the East

India margin is on the hanging wall side of the master detachment fault, which is responsible for mantle

exhumation and continental breakup. This can be supported by the presence of terraced geometry,

narrower distal margin width, thicker lower crust, relatively steeper break-away fault, and concave-

down geometry of detachment fault related to mantle exhumation on the East Indian margin. The

conjugate Elan Bank margin representing the lower plate margin is characterized by wider distal margin,

wider coupled domain, thinner lower crust and concave-up geometry of the master detachment fault.

The width of the coupled domain in paired profiles is different. This indicates some amount of

asymmetry on conjugate margin. This margin asymmetry is another indicator of complex breakup

localization along the margin, which is a key to understand microcontinent release mechanism.

5.6 Tectonic timing

A precise timing of continental breakup in the hyper-extended margin is always difficult to define. The

best option is direct sampling and dating of the oldest sediment deposited over the oldest oceanic crust,

which was created by the organized sea floor spreading (Manatschal et al., 2007). Unfortunately, this

data is extremely rare and almost non-existent at most of the passive margins. In the current study,

several methods using different data sets are combined to narrow down the error margins and to

determine a realistic relative timing of continental breakup along different margin segments. They have

been integrated to understand the timing of microcontinent release and associated mechanism.

5.6.1 Breakup timing

All the available Gondwana plate reconstructions (Powell et al., 1988; Gaina et al., 2007; Reeves. 2008;

Veevers, 2009) indicate that the Bay of Bengal opening initiates from northern Mahanadi Basin and

propagated southwards in a zipper-like fashion to the Cauvery Basin. The first marine incursion

advanced from northern Tethys. The style of opening of Bay of Bengal can be further supported by the

earliest marine transgression recorded in the drilled wells near north Vizag transfer zone in East India. In

the other drilled wells down south, the age of first marine transgression is later in age. This evidence is

also concurrent with the appearance of first magnetic anomaly (M11) in Valanginian (132 Ma) in the Bay

of Bengal (Ramanna et al., 1994). The M11 magnetic anomaly is again interpreted as the first magnetic


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anomaly in Enderby basin, which is located between Elan Bank and Antarctica (Ramanna et al., 1994;

Gaina et al., 2007). This provides a broad spectrum about the timing of the opening of the first ocean.

The paleo-environment indicators from wells typically show that the subsidence is quite gradual in the

inner part of proximal margin (Aborted H Block) in the extensional Krishna-Godavari margin (Figure

5.27). In this setting, the paleo-bathymetry data from the two wells (P1, P2) suggest that the continental

environment during early rifting, progressively changed into shallow marine (>20 m bathymetry) and

then deeper marine (>200 m bathymetry) during the entire passive margin evolution (Figure 5.28). In

the northern Krishna-Godavari proximal margin, the continental to shallow marine transition occurs

close to early Valanginian (137 Ma). This is confirmed from palynology and foraminifera records from

the wells. Interestingly, this change also coincides across a seismically mapped major unconformity (K20

horizon) (Figure 5.29). All rift faults are sealed by this unconformity, allowing this unconformity to be

interpreted as ‘breakup unconformity’. It is tempting to interpret breakup time as Valanginian (137 Ma),

based on this ‘breakup unconformity’ interpretation. In proximal margin setting, classical breakup

unconformity based interpretation of continental breakup can be misleading. Similar situations are also

observed in the other passive margins like erstwhile Adrian-European margin (Manatschal, 2004).

Therefore, this unconformity and associated timings cannot be conclusively interpreted as breakup

timing.

In proximal margin setting, the deep marine transition beyond 200 m occurs during Late Hauterivian

(130 Ma). The deep marine transition beyond 500m further records a rapid change in subsidence

pattern, which is close to early-Barremian-Hauterivian boundary (130-128 Ma) (Figure 5.30). The

correlated surface in this case, is not an unconformity but a conformable sequence boundary (K30).

Interestingly, this first deep marine sequence is quite localized and cannot be mapped towards the

oceanic region. Therefore, the first deep marine envelope, which can be mapped across the entire basin,

definitely postdates the initial thermal subsidence and formation of first oceanic crust (Manatschal et

al., 2007; Péron-Pinvidic and Manatschal, 2010). In this study, the equivalent time, when the first deep

marine envelope is formed has been taken as the closest proxy to estimate the true continental

breakup. Accordingly, the breakup in proximal northern Krishna-Godavari Basin is estimated as taking

place no earlier than Hauterivian-Barremian boundary (130 Ma). It can be stated that the actual breakup

must be post 130 Ma in northern end of Krishna-Godavari Basin.


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Figure 5.27: Well based paleo-environment indicators shows different responses and timings at different margin settings in orthogonally rifted northern Krishna-Godavari rift zone. The Aborted H Block in proximal margin shows deepening is transitional across an unconformity. While residual H block in distal margin shows transition across a conformable section. The deep-water transition in distal margin is relatively younger than the proximal margin.

There are five wells (T1 to T5) are located in the outer part of proximal margin in the Krishna-Godavari

Basin. These wells show sudden change in paleo environment. The continental environment directly

changes into a deeper marine one, without any presence of shallow marine sequence across the

mapped regional ‘breakup unconformity’ K20 horizon (Figure 5.27). It is to be noted that these wells

have significant missing sections, which ranges from Late Valanginian (134 Ma) to Barremian-early

Aptian (125 Ma). Hence, in this setting, the first convincing deeper marine sequence is identified as post

Barremian-early Aptian (125 Ma). This timing is estimated by detecting the change in foraminifera

count. Unfortunately, these are located on the horsts. Therefore, non-deposition as well as erosion

related to breakup cannot be ruled out. Yet, the subsidence patterns and paleo-environment changes in

this whole area indicates that the breakup age is no younger than Barremian-early Aptian (125 Ma).

As we move further towards the distal margin, the age correlation and determining the breakup timing

becomes clearer. Well D1 is located in a graben setting in the northern distal Krishna-Godavari Basin.


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Figure 5.28: Interpretation of tectonic timings from paleo-environment chart. The change in paleo-bathymetry is recorded, where the transition from continental environment changes into shallow marine and then deeper marine (>500 m bathymetry). See text for detail.

This well shows an almost complete record of paleo-environment change (Figure 5.27). Being located in

the distal margin graben with complete record of paleo-environment change, this well provides the

closest timing estimate. The paleo-environment change from continental to shallow marine occurred

around Late Hauterivian (132 Ma). The shallow-marine to deeper-marine transition occurred around

Barremian-Early Aptian (125 Ma). This is also evident in the subsidence profile, which shows a change in

subsidence pattern around 125 Ma (Figure 5.30). Besides, the correlated regional K20 horizon is a

conformable surface here, against which the deeper marine transition beyond 500m occurs. The deeper

marine sequence above this transition (K30) can be mapped across the Krishna-Godavari Basin as the

first deep marine envelope. This information provides the closest estimation of the true breakup timing

around Barremian-early Aptian (125 Ma) in this region. Therefore, the breakup between India and Elan

Bank in northern Krishna-Godavari Basin is estimated to take place close to 125 Ma.


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Figure 5.29: A drilled well shows paleo-environments and age controls in the proximal part of northern Krishna-Godavari rift zone. The reflection seismic image shows different interpreted horizons. Note, the clear cut unconformity at 137 Ma (Valanginian), correlated as K20 horizon. To interpret breakup time as Valanginian (137 Ma) by mapping the age difference along this ‘breakup unconformity’ can be misleading. It has to be noted that the major subsidence change occurs by 130 Ma. This unconformity marks the end of rifting (stretching and thinning phase) process and onset of passive margin development.

The discrepancy in breakup age in inner and outer proximal margin simply remained non-determinant

due to lack of more well records in different tectonic settings. This may also relate to interpreted rift-

jump from its initial location to further eastward (Figure 5.27). The basement and tectonic element

maps together show two clear fault trends (Figure 5.31). While the inner proximal margin shows roughly

NW-SE trend, the outer proximal margin shows roughly east-west trend. Both fault trends can be linked

through a divergent overlapping transfer zone separated by a high relief accommodation zone (Figure

5.31).

The southern part of Krishna-Godavari Basin, which is close to the northern end of Coromandal horsetail

structure, represents an oblique-slip extensional setting. Here well P3 is located in proximal margin

graben setting. This well contains a complete paleo-environment change record. The well data show the

continental to shallow marine transition that happened sometimes during Valanginian-Hauterivian (135-

132 Ma). The transition from shallow to deep marine took place close to the Aptian time (123 Ma).

Seismically, there is no recognizable ‘breakup unconformity’ is present (Figure 5.32). On the contrary to


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northern Krishna-Godavari basin, the correlated regional K20 horizon in this proximal graben is a

conformable surface instead of an unconformity. Although most of the sequences are quite localized in

the proximal graben, the timing of continental to shallow marine transition is quite similar to the timing

estimated (123 Ma) in the proximal northern Krishna-Godavari Basin. This proximal shallow marine

sequence, which mixed with significant continent derived materials, probably represents a sag phase

deposit in proximal margin. The mixing of sediments has been supported by mineralogical, textural and

palynology evidences. Therefore, it can be stated that in the extensional margin north of the

Coromandal sheared margin, the extension in distal margin continued up to early Aptian (123 Ma).

Figure 5.30: Subsidence profiles in northern rift zone calculated at 1D location of wells T2, P1 and D1. The subsidence curves are quite gradual in all profiles. However, the well D1, located in distal margin is showing a steeper gradient. Note, the effect of water depth is corrected in all profiles.

The two other wells (T7 and T8) are drilled at the outer proximal margin on horsts (Figure 5.28). These

wells are also located in a terraced architecture similar to the northern Krishna-Godavari Basin setting.

Therefore, the shallow marine sections in these wells are eroded or not deposited, just like in the case of

their northern Krishna-Godavari counterparts. These wells show that the enveloping deep marine

transition took place near late Aptian time (123 Ma). Therefore, the estimated timing of continental

breakup in this region is close to 123 Ma. It is to be noted that the breakup is almost 2 Ma younger than

that of the northern end of Krishna-Godavari Basin. However, such precision cannot be guaranteed as


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error margin of this estimation can vary based on each drilled well data analysis. It would be reasonable

to mention that the breakup in southern Krishna-Godavari Basin is either coeval or relatively younger

than that in the northern Krishna-Godavari Basin.

Figure 5.31: The structural element map in north Krishna-Godavari basin is showing two clear fault trends. While the inner proximal margin shows roughly NW-SE trend, while the outer proximal margin shows roughly east-west trend. Both the fault trends can be linked through a divergent overlapping transfer zone separated by a high relief accommodation zone. The age of synrift sections in inner and outer trends are different as shown in Figure 5.28. Comparing structural and age information, the rift jump event is documented.

In the central part of the Coromandal horsetail structure the single well (T9) was drilled on a horst in the

distal margin. The shallow marine section is not present, which may be due to either erosion or non-

deposition. The paleo-environment data show a rapid transition from continental to deeper marine

within a very short period of time (Figure 5.28). Being located at sheared margin, this is also reflected in

a steep subsidence curve (Figure 5.33). The timing of the transition comes close to 120 to 122 Ma, which

can be closest to the exact breakup timing in this area. This means that the breakup is relatively younger

in the strike-slip dominated Coromandal segment in comparison to extensionally rifted Krishna-Godavari


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Basin. However, it is associated with an uncertainty due to abnormal presence of arenaceous benthonic

foraminifera in the sediments deposited during the transition.

Figure 5.32: A drilled well shows paleo-environments and age controls in the proximal part of the southern Krishna-Godavari rift zone. The major seismic sequences are interpreted in the reflection seismic image. The well located in proximal margin graben setting is showing complete paleo-environment change record. Here also the based on subsidence and transition to deep-water is occurring close to within Aptian (123 Ma). Seismically, there is no recognizable unconformity is present. The correlated region K20 horizon is a conformable surface. The other wells located in outer proximal margin in horst setting are also showing the deep water transition near late Aptian (123 Ma). This is matches with the subsidence pattern.

Situation in the Cauvery Basin is a bit different. A unique tectonic setup in Cauvery Basin needs to be

considered carefully. In its northern part, the rift-structures filled with thicker syn-rift sediments are

present near to ocean-continent transition (Figures 3.4, 3.5). The rift structures filled with thicker syn-

rift sediments, however, are typical for the proximal margin setting. So, the presence of proximal margin

setting close to the ocean-continent transition is anomalous for any hyper-extended passive margin.

Here the normal fault controlled breakup is observed in southern part of the Cauvery Basin only (Figures

3.1, 3.2, 3.3). Unfortunately, there is no well control there.

There are two wells (P4 and T10) present in the northern region. Similar to the Krishna Godavari Basin,

the interpreted seismic ‘breakup unconformity’ (K20) represents a bathymetric transition from

continental to shallow marine in both inner and outer proximal margins in the Cauvery Basin (Figure

5.28). In the northern Cauvery Basin, the shallow-water to deep-water transition has happened close to

the early Valanginian (137 Ma) (Figure 5.28). This age is a little bit older in comparison to the Krishna-


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Godavari margin while similarity in age can be drawn between Krishna-Godavari and Cauvery Basin

when the first deep marine enveloping surface is considered. The first deep marine transition that

envelopes the entire basin occurs during early Aptian (122 Ma) only. The entire section from 135 to 120

Ma is missing in one well (P4). In another well (T10), the missing section ranges from 130 to 122 Ma.

Figure 5.33: Subsidence profiles in the central Coromandal strike slip system calculated at 1D location of well T9. The subsidence is quite rapid in strike slip dominated domain. Note, the effect of water depth is corrected in all profiles

Analyzing the tectonic setting, in which wells are located, it is interpreted that much of the sections are

eroded probably due to post breakup isostatic uplift. It is well established that the extensional margins

undergo significant post breakup uplift due to isostatic rebound of the whole lithosphere as a result of

thermal relaxation (Buck, 1986; Weissel and Karner, 1989; Nemčok and Rosendahl, 2006b). However, at

sheared margin, the isostatic uplift appears to be less significant. Therefore, it is expected that there

would be a significant margin uplift at Krishna-Godavari and Cauvery extensional margins and less uplift

at Coromandal sheared margin. This can be expressed by the sediment thickness, which is a function of

isostatic uplift due to flexural unloading (Weissel and Karner, 1989) and dynamic topography change as

a result of the change in mantle convection process (Gurnis et al., 2000; Conrad and Gurnis, 2003). The

rejuvenation and rearrangement of post-rift drainage system would increase the topography-driven

sedimentation in the deeper water environment. As a result, one should expect a thicker sediment cover

during the early breakup (K20-K30) interval (Figure 5.34). However, the thickness map clearly indicates


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that the sediment thickness is quite high in Krishna-Godavari Basin in comparison to the Cauvery Basin.

The Krishna-Godavari system is clearly providing more sediment towards the deeper basin in

comparison to the Cauvery Basin. As expected for the sheared margin, the Coromandal strike-slip zone

does not show much sediment thickness variation. The time thickness map of the next interval (K30-

K60) shows a thicker sediment cover in both Krishna-Godavari and Cauvery basins towards the deeper

water environment (Figure 5.35). The Coromandal segment is still showing not much variation in

sedimentary thickness.

Figure 5.34: Time thickness map K20-K30 is showing relative change in post breakup dynamic topography change as a result of isostatic uplift. It shows that the Krishna-Godavari system dumping more sediment towards the deeper basin as compared to the Cauvery basin. The Coromandal strike-slip zone in between does not show much sediment thickness variation or extent as expected in a strike-slip setting. See text for detailed explanation.


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Figure 5.35: Time thickness map K30-K60 is showing relative change in post breakup dynamic topography change as a result of isostatic uplift. The time thickness map of the next interval (K30-K60) is showing thicker sediment cover in both Krishna-Godavari and Cauvery Basin. The Coromandal segment is still showing not much variation. The deeper extent of sediment distribution is also observed. See text for detailed explanation.

If these observations are combined with tectonic element map and East India-Elan Bank reconstruction,

it appears that the Coromandal strike slip fault plays a major role in the breakup localization and final

breakup of Elan Bank from India. The Cauvery region is also affected by the same process, at least in the

outer proximal margin. Eventually, this means that the post-breakup topography driven sedimentation

was not well developed in the Cauvery Basin during K30-K20 interval. As per biostratigraphic records

from wells (P4 and T10), this time shows a missing section. Therefore, it can be inferred that the

isostatic uplift and change in dynamic topography must have taken place bit later in the Cauvery Basin.


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However, as per the tectonic reconstruction, the Cauvery rift zone was not really a conjugate to Elan

Bank. It formed the south-westward continuation of the southern margin of Elan Bank. This segment

was actually conjugate to Antarctica (Figure 5.20). Now, as there is a microcontinent release associated,

the Cauvery-Antarctica breakup must pre-date the microcontinent release.

Therefore, as indicated in seismic interpretation, the normal fault-controlled breakup between the

Cauvery Basin and Antarctica took place earlier than India-Elan Bank breakup. This leads to an inference

that the Cauvery Basin underwent two continental breakups. The first one took place between the

south-eastern normal fault-controlled Cauvery rift zone and southern end of Enderby Basin of

Antarctica. This breakup can be probably dated as taking place at around Valanginian (132 Ma) as per

the oldest observed M11 in the Enderby Basin (Ramanna et al., 1994; Gaina et al., 2007). The second

breakup took place while the Elan Bank was clearing the margin. This breakup was controlled by the

Coromandal fault zone. This took place sometimes around early Aptian (122-120 Ma). Therefore, this

region must have suffered two isostatic uplifts. The effect of the first isostatic uplift can be interpreted

based on several seismic unconformities mapped in the Cauvery basin seismic data and

biostratigraphically indicated missing zones. The second isostatic uplift is indicated by the dynamic

topography change. The presence of inverted proximal rift setting in the northern Cauvery Basin was

possibly developed during second continental breakup. During the second continental breakup, it was

already a failed rift and the Coromandal strike-slip fault was chipping parts of the failed rift away while

clearing the margin.

The situation can be summarized in following way;

1. There are two breakups and associated isostatic uplifts in the Cauvery Basin;

2. There is a single breakup and associated isostatic uplift in the Krishna-Godavari Basin;

3. The oldest breakup took place between the south-eastern Cauvery Basin and Antarctica during

latest Valanginian (132 Ma);

4. The breakup along the southern Krishna-Godavari zone (123 Ma) is slightly younger than that

along the northern Krishna-Godavari zone (125 Ma); and

5. The breakup is youngest along the Coromandal fault zone (120-122 Ma), which is associated

with the second breakup event in Cauvery Basin as Elan Bank slide past the Indian margin using

Coromandal strike-slip system.


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5.6.2 Hot spot timing

The role of hotspot in Elan Bank microcontinent release is a crucial issue needed to be understood. It

was the activity of the Kerguelen plume related magmatism, which was believed to be almost

synchronous with Elan Bank separation (Műlller, 2001; Gaina et al., 2007). The extent and timing of the

Kerguelen plume magmatism is well established in the Indian Ocean region, Bay of Bengal and onshore

India. The Rajmahal trap and Shyllet trap volcanic province in onshore East India are products of the

Kerguelen Plateau volcanism (Baksi et al., 1995; Kent et al., 2002; Ghatak and Basu, 2011). These

volcanic rocks, sampled from different outcrops and boreholes, were radiometrically dated. 40Ar/39Ar

dating method carried out on samples from Rajmahal trap delivers an age, close to 117.4 ± 0.2 Ma

(Coffin et al., 2002; Kent et al, 2002; Ghatak and Basu, 2011). The basalt samples from Bengal and

Mahanadi Basin boreholes deliver an age, close to 116.2 ± 0.3 Ma (Baksi et al., 1995; Kent et al, 2002).

The lamprophyre dykes in onshore Bengal Basins are also dated by same method, giving an age range of

114.9 ± 0.3 Ma (Baksi et al., 1995; Kent et al, 2002).

The offshore expression of Kerguelen plume magmatism is represented by volcanic 85° E Ridge, which is

considered to be the hotspot track (Choudhuri et al., 2013). The 85° E Ridge is not directly dated through

any recovered samples, apart from a single well in the offshore Mahanadi Basin, which provides data

from recovered volcanic sample. This well is located close to continental-proto-oceanic boundary, giving

the age of basalt as 112 .2 ± 1 Ma by 40Ar/39Ar dating method. It is exactly the place, where the hotspot

track traverses from continental to oceanic region. The age data control points obtained from exsisting

onshore and offshore samples of volcanic rocks is shown in Figure 5.36. Interestingly, all e data points

can be linked to offshore 85° E Ridge to show the continuation of linear hotspot track from onshore to

offshore.

It is very clear from these dates that the Kerguelen volcanism came at least 3 to 8 Ma later than the

continental breakup between India and Elan Bank. It has to be noted that this age gap is valid only for

the Bengal and Mahanadi basins, which are further north from the reconstructed Elan Bank. Therefore,

down south, the age of hotspot related volcanism is expected to be younger than that of its northern

portions. The date of oldest basalt volcanism sampled from ODP-1137 well in Elan Bank also goes back

to 107.7 ± 0.5 Ma. This suggests at least a 12 Ma difference between East India-Elan Bank breakup and

the Kerguelen magma output. Therefore, a direct role of Kerguelen mantle plume in Elan Bank breakup

can be ruled out.


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Figure 5.36: Hotspot track created in Bay of Bengal as a result of migrating the plate over Kerguelen plume (Backdrop: Free air gravity, Snadwell and Smith, 1997). The age constraints obtained from radiometric dating (Coffin et al., 2002) have been plotted on top of the track. Note, there is no definite age control is present in the deep offshore region. The main age constraints are from onshore outcrops and subcrop data. Only single data point exisit in deep offshore on 85° E ridge.

5.7 Geodynamic evolution and release of Elan Bank microcontinent

The hyper-extended crustal architecture on India and Elan Bank indicates that the passive margin

evolution is a result of multi-phase extensional process. The lithospheric rupture mechanism and

kinematics behind passive margin evolution from its pre-rift stage to breakup stage included four major


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processes similar to the other passive margins in the world (see Manatschal et al., 2007; Péron-Pinvidic

and Manatschal, 2010), which includes;

a) Stretching (a broad distribution of high angle brittle faults in the upper crust);

b) Thinning (lithospheric necking and detachment faulting);

c) Exhumation (exhumation of continental mantle at continental breakup location); and

d) Sea floor spreading (an irrevocable localization of tectonic and magmatic processes within the

plate boundary).

The model of the Elan Bank microcontinent separation reached by this study shows a complex

interaction of all four processes. These processes were in effect in a sequential manner and ultimately

release the microcontinent during continental breakup (Figure 5.37). The data along the margin suggest

that these processes varied in time and space.

The stretching phase initiated with the rift phase initiation during Toarcian-Alenian (180-176 Ma). It was

characterized by broadly distributed strain over a large region encompassing India-Elan Bank-Antarctica-

Australia. The upper crustal high-angle normal faults (blue faults in Figure 5.37) were the primary

features accommodating the stretching phase. These faults later acted as the primary basin bounding

faults, controlling the rift geometry. At present, they are mostly limited to the proximal margin setting

within H blocks (Figures 5.2, 5.3, 5.4). During the stretching phase, rifting was largely decoupled and not

yet localized. However, due to the effect of pre-existing anisotropies, it was effectively localized close to

the pre-existing Precambrian mobile belts, which surround the major cratons. This is shown in 168 Ma

reconstructions between India-Antarctica and Elan Bank (Figure 5.38). Jurassic rift systems were

developed in Mahanadi, Krishna-Godavari and Cauvery basins. During this phase, rifting took place at

Elan Bank.

The stretching phase was followed by thinning phase, where the broadly distributed extension was

localized within narrow focused zones. In the Krishna-Godavari and Cauvery basins, the thinning stared

no earlier than during Beriassian (140 Ma) and continued no later than until Valanginian-Hauterivian

(132 Ma) (Figure 5.39; 5.40). This can be implied from the well data, which document bathymetry

variation between 20-200m indicating a gradual deepening of the basin during this interval. As the

extension continued in this region, the deformation started to localize along a few major faults (green

faults in Figure 5.37), forming the future necking zones between India-Elan Bank and Elan Bank-


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Figure 5.37: Schematic kinematic evolution model for India-Antarctica passive margin and microcontinent release. This model shows a schematic kinematic restoration of the conjugate margins. This model is prepared based on competing rift zone hypothesis followed by Péron-Pinvidic and Manatschal (2010).

Antarctica (Figure 5.37). It has to be noted that in subsequent time from Middle Beriassian (135 Ma)

onwards, the deformation became localized at the edges of relatively stronger lithospheric blocks.

During this phase, a conjugate system of thinning faults separated the relatively stronger lithospheric

blocks in their footwall. This caused individualization of the Elan Bank as a strong and relatively less

deformed block of crust. This individualization was the primary step in creating a future microcontinent

(Figure 5.37). However, at this point, it was not sure if it would become a microcontinent or remain as a

continental ribbon attached to India. The intense strain localization was an irreversible process that was

mostly concentrated within the H-blocks in the hanging wall of these conjugated pairs of thinning faults.


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Figure 5.38: The Elan Bank reconstruction at Bajocian (168Ma). Reconstruction (modified after Reeves, 2008).It shows the Initial rift between India-Antarctica-Elan Bank. These rift signatures indicate stretching phase. Note that the rifts are localized within crustal weak zones. TBSZ : Terrain boundary shear zone; NVSZ: Nagavalli-Vamshadhara shear zone; SCC: Singhbhum cratonic complex; BCC: Bastar cratonic complex; DCC: Dharwar cratonic complex; VC: Vijayan complex; HLC: Highland cratonic complex; RC: Rayner complex; NC: Napier complex; EB: Elan Bank

The hyper-extended crustal architecture model in the region between Antarctica and Elan Bank is

schematically interpreted based on published seismic profiles in Stagg et al, (2005) and the tectonic

element map in O’brien and Stagg (2007). The interpreted margin architecture suggests that the rate of

extension between India-Elan Bank and Elan Bank-Antarctica was asymmetric. Therefore, it is

interpreted that the exhumation process between Elan Bank and Antarctica starts relatively earlier than

that between East India-Elan Bank (Figure 5.37). This exhumation was taken place in Beriassian, when

thinning phase still continued between East India and Elan Bank. This step further works on the

development of Elan Bank as either future microcontinent or future continental ribbon attached to


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India. The change in the extension rate is probably an effect of highly anisotropic lithosphere in the

India-Antarctica region. The highly different rates of extension are also the key to understand the early

breakup and the sea floor spreading initiation between the Mahanadi rift zone and Antarctica (Figure

5.40). Here the lack of data does not allow identifying the timing of each phase, but it seems that the

rifting is very fast in this region, which ends up in lithospheric breakup by 132 Ma. A combined rate of

extension in the Krishna-Godavari and Mahanadi basins and its western and eastern end translates to

movements of about 27 and 52 km per million years (km Ma-1) (Nemčok et al., 2013a), which is quite

fast in a comparison with other margins of the world.

Figure 5.39: The Elan Bank reconstruction at Beriassian-Valanginian (140 Ma). The Reconstruction is modified after Reeves 2008). Sea floor spreading data (Muller et al., 2008). This reconstruction shows initiation of thinning in Krishna-Godavari (KG), Cauvery (CY), Elan Bank (EB) and Antarctica (ANT). Initial sea-floor spreading is from north and break up between Mahanadi Basin and Antarctica. Note rifting continues between KG basin- Elan Bank-Enderbyland and Cauvery basin and Antarctica. TBSZ : Terrain boundary shear zone; NVSZ: Nagavalli-Vamshadhara shear zone; SCC: Singhbhum cratonic complex; BCC: Bastar cratonic complex; DCC: Dharwar cratonic complex; VC: Vijayan complex; HLC: Highland cratonic complex; RC: Rayner complex; NC: Napier complex; EB: Elan Bank


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Figure 5.40: The Elan Bank reconstruction at Valanginian-Hauterivian (132 Ma). The Reconstruction is modified after Reeves 2008, Reliance internal report. Sea floor spreading data (Muller et al., 2008). Antarctica COB (Obrien and Stagg, 2007); Elan Bank structure (modified after Borissova, et al., 2003). Note that the breakup occurs between Cauvery-Antarctica and Elan Bank Antarctica. A major rift-jump occurs in the Krishna Godavari basin as observed as major change in fault geometry and orientation. The KG rift continues to propagate westward. This phase represents little amount of stretching but mostly thinning related deformation localized in the future distal margin. Note that a triangular gap between India-Elan bank at this stage. TBSZ : Terrain boundary shear zone; NVSZ: Nagavalli-Vamshadhara shear zone; SCC: Singhbhum cratonic complex; BCC: Bastar cratonic complex; DCC: Dharwar cratonic complex; VC: Vijayan complex; HLC: Highland cratonic complex; RC: Rayner complex; NC: Napier complex; EB: Elan Bank

The sea-floor spreading was initiated from north when the arm of Tethys entered into future Bay of

Bengal and a new ocean was progressively opened in a zipper-like fashion towards south (Muller et al.,

2008). Although not well constrained, it is assumed that the continued exhumation between Elan Bank

and Antarctica finally caused building up of the asthenospheric anomaly at the future continental

breakup location. The passive rise of magma chamber and adiabatic melting was about to set in to break


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the lithosphere. Therefore, the early oceanization was about to begin by this time. The reconstruction

shows that early sea floor spreading was initiated between Elan bank and Antarctica by Hauterivian (132

Ma) causing the final continental breakup (Figure 5.40). Apart from biostratigraphic interpretation, the

first breakup between Antarctica and eastern part of the Cauvery rift zone is also interpreted based on

the Mesozoic magnetic anomaly data (M11 to M0) in the south of present day Sri-Lanka (Desa et al.,

2006; Műller et al., 2008). Post, M0, these magnetic anomaly patterns shows major spreading

reorganization. The oldest magnetic anomaly (M11) in the Enderby Basin, which was formed during

Hauterivian (Ramanna et al., 2001; Gaina et al., 2007; Gibbons et al., 2013) is a supportive evidence of

the early breakup.

However, the data in East India further suggest that the extension never ceased between India and Elan

Bank to delink itself from the entire kinematic system. The extension between India and Elan Bank was

reorganized simultaneously as a result of eastward rift jump. Therefore, the along-strike rifting

trajectory changed its orientation (Figure 5.40). This eastward rift jump in the northern Krishna-Godavari

basin continued to propagate westward. The rift propagation along the Krishna-Godavari rift zone did

not stop even after the breakup in the overstepping Cauvery rift zone, where the breakup was already

completed by then. In addition to the complex lithospheric heterogeneity between East India and Elan

Bank, the intense strain localization can be the output of slow asymmetric spreading between Antarctica

and Elan Bank. The asymmetric extension helped to localize the strain at the edge of an H block.

As a result of the progressive extension, the lower continental crust in India and Elan Bank thinned down

to almost zero at about Hauterivian time (132 Ma) in the future continental breakup location. The crust

was thinned up to a limit where no ductile lower crust remains. As a result of that, the deformation in

the brittle upper crust and brittle upper mantle became coupled and the exhumation fault (red fault)

started to form (Figure 5.37). This fault eventually exhumed the sub-continental mantle along the

breakup trajectory. It was the geometry of the exhumation fault, which defined the upper plate-lower

plate architecture for East India and Elan Bank conjugate.

As the extension continues between India and Elan Bank, the western propagating end of the Krishna-

Godavari rift zone became eventually hard-linked with western portion of the Cauvery rift zone through

the Coromandal dextral strike-slip fault, sometimes after Barremian (123 Ma) (Figures 5.41). The

transfer fault hard-linked the normal faults systems of both rift zones via its northern and southern

horse-tail structures. At that point, an effective kinematic linkage got established between the failed


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Cauvery rift zone and the portion of the Krishna-Godavari rift zone that underwent a crustal breakup

and undergone the exhumation phase. This kinematic linkage continued until Early Aptian (123 Ma) in

Krishna-Godavari rift zone. The first oceanization and continental breakup must have started by this

time. A new spreading center has been formed between Elan Bank and East India as a result of

continued fast extension (Figure 5.41). It is possible that both asthenospheric anomalies between East

India-Elan bank and Elan Bank-Antarctica were simultaneously active for a while, competing for the final

continental breakup by Early Aptian (123 Ma).

Figure 5.41: The Elan Bank reconstruction at Aptian (123 Ma). The Reconstruction is modified after Reeves 2008. Sea floor spreading data (Muller et al., 2008). Antarctica COB (Obrien and Stagg, 2007); Elan Bank structure (modified after Borissova, et al., 2003). Note that the breakup occurs between East-India and Elan Bank. Due to intense strain localization between India-and Elan bank with continued rift propagation, a new spreading center in formed between India and Elan Bank. The older spreading center between Elan bank-Antarctica eventually deactivated and active ridge jumped towards west During the rift jump and western propagating end of KG is hard-


237

linked with western portion of the Cauvery rift zone. Kinematic linkage established between the failed Cauvery rift zone Krishna-Godavari rift zone that underwent a continental breakup. TBSZ : Terrain boundary shear zone; NVSZ: Nagavalli-Vamshadhara shear zone; SCC: Singhbhum cratonic complex; BCC: Bastar cratonic complex; DCC: Dharwar cratonic complex; VC: Vijayan complex; HLC: Highland cratonic complex; RC: Rayner complex; NC: Napier complex; EB: Elan Bank

The Krishna-Godavari rift zone, which did not stop competing for breakup captured the final breakup

location in the region to the north of the future Elan Bank microcontinent. This resulted in a ridge jump.

The new spreading center was then active between East India and Elan Bank and it started to propagate

southward with continued movement of Coromandal system. As a result, the active ridge is also jumped

from the initial Elan Bank-Antarctica breakup corridor towards the Elan Bank-East India breakup

corridor. This northern spreading system was eventually connected to the southern spreading system

(Figure 5.42). Thus, the older spreading canter in Antarctica got abandoned as a result of the final

breakup localization. At that point, the Elan bank was finally became a microcontinent and it was

actively transferred to the present day Antarctic plate. The Coromandal transform was active for the

entire time necessary for the Elan Bank microcontinent and the sea-floor spreading center to the north

of it laterally clearing the Coromandal margin. Only by late Albian-Aptian time, the organized sea-floor

spreading system has been established. The Coromandal fault activity is stopped at that time. Therefore,

as a result of competing rift zones and breakup localization, the Elan Bank microcontinent was finally

released.


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Figure 5.42: The Elan Bank reconstruction at Albian-Aptian (112 Ma). The Reconstruction is modified after Reeves 2008, Reliance internal report). Sea floor spreading data (Muller et al., 2008). Antarctica COB (Obrien and Stagg, 2007); Elan Bank structure (modified after Borissova, et al., 2003). Active kinematic linkage between different spreading segments established. Elan Bank is transferred to the Antarctic plate and moving away The typical organized sea floor spreading established in early Indian Ocean. The Coromandal transform activity slowly dies out at this time. TBSZ : Terrain boundary shear zone; NVSZ: Nagavalli-Vamshadhara shear zone; SCC: Singhbhum cratonic complex; BCC: Bastar cratonic complex; DCC: Dharwar cratonic complex; VC: Vijayan complex; HLC: Highland cratonic complex; RC: Rayner complex; NC: Napier complex; EB: Elan Bank

5.8 Match of Elan Bank with hot-spot track in Bay of Bengal

The 85°E Ridge being a hotspot trail (Choudhuri et al., 2013) related to Kerguelen plume provides a track

for hotspot movement over Indian plate (Figure 5.36). The 85°E Ridge can be correlated with Rajmahal

trap in onshore India through, Bengal basin volcanics, then offshore volcanism as the ridge itself,

trending north to south and it is finally connected to Afansy-Nikitin seamount in the south of Sri-Lanka,

where the track bends towards south west (Krishna, 2003). It records the movement of Indian plate over


239

the Kerguelen hotspot. Now if close comparison is drawn from the Elan Bank displacement tracks, which

has been interpreted from the geodynamic evolution of Elan Bank as described previously, the Elan Bank

displacement vector is showing roughly towards south east with occasional bends due to tectonic

adjustments during extension. Interestingly, both tracks eventually intersect each other at quite oblique

angle. This provides evidence that the plume related volcanism did not affect the microcontinent

separation (Figure 5.43).

Figure 5.43: The Elan Bank reconstruction at Albian-Aptian (112 Ma) shows relationship between Elan Bank drift path and hotspot track. The Reconstruction is modified after Reeves 2008, Reliance internal report). Sea floor spreading data (Muller et al., 2008). Antarctica COB (Obrien and Stagg, 2007); Elan Bank structure (modified after Borissova, et al., 2003); Age of volcanics (Coffin et al., 2002). The Elan Bank drift path and hotspot trail are intersecting at a high angle. Note that synchronous volcanic emplacement age in Elan Bank and Mahanadi basin, which are almost 600 km apart.


240

241

Chapter 6

Discussions

The study indicates that the evolution of the Elan Bank microcontinent is a crucial part of India-

Antarctica rifting history, without which the geodynamic understanding of Eastern Gondwanaland

breakup and evolution of Indian Ocean remains incomplete. Furthermore, the East Indian margin

architecture has been interpreted as a major guiding tool to understand the architecture of its Elan Bank

conjugate. This study also addresses the several previously unanswered topics, including the crustal

architecture of the Elan bank to constrain its shape and size using different methods. The position of

Elan Bank with respect to East India has been addressed in reconstruction, which relies on regional data

constraints rather than global plate reconstruction models. The study further specifies, although not

very exclusively, the necessity for restoration of hyper-extended margins to its pre-kinematic positions.

The proper restoration of margin architecture is a basic requirement for understanding its evolution.

There are several important issues addressed in this study including the importance of tectonic timing.

This study allows one to say that the evolution of Elan Bank as a microcontinent is a result of complex

continental breakup processes, which involve not only hyper-extension but also breakup localization

along the margin that drives the ridge jump. These processes need to be discussed for a comprehensive

understanding of the process of microcontinent evolution.

6.1 Elan Bank microcontinent in India-Antarctica plate reconstruction models

Despite of their importance in kinematic evolution of passive margins, microcontinents are

conspicuously absent in most of the publicly available global plate tectonic reconstruction models,

Chapter-6 Discussions

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including geometric reconstructions (Reeves, et al., 2002; Reeves and DeWit, 2004; Reeves, 2009),

paleomap projects (Scotese et al., 2003) integrated geological reconstructions (Harrowfield et al., 2005),

geological and paleo-magnetic reconstructions (Collins and Pisarevsky, 2005). As mentioned in Chapter-

1, these global plate tectonic models do not incorporate Elan Bank microcontinent into their models.

Delineation of the pre-rift extent of continental crust is a necessary constraint to obtain a meaningful

full-fit plate reconstruction. A number of plate tectonic models have used various physical and

geophysical markers as proxies for the pre-breakup extent of continental margins, particularly in the

case of India. These includes the 1000m isobaths (Reeves, 2008), prominent Free-air gravity anomalies

(Ramanna et al., 2001; Subramanyam and Chand, 2006) and horizontal gradient of gravity anomalies

(Schettino and Scotese, 2005). This assumption to use either of the criteria as proxy for the pre-rift

extent of continental crust has no clear justification for why such parameters will represent the ocean

continent transition. As a result, a direct comparison is challenging between the synthetic

reconstruction carried out in this study and approximations used in other models.

Another reason for mismatch of this study results with published plate reconstruction can be the fact

that the traditional plate-tectonic reconstruction models are often based on modern day geometric

boundaries of the margin and simple rotation poles of larger plates, and usually do not take into account

of the individual rotation poles of smaller plates, including microcontinents. Therefore, rotation pole

based geometric reconstructions end up with an erroneous result creating serious gaps or overlaps,

particularly if microcontinents are involved. As most of the tight-fit reconstructions show partial gaps

(Scotese et al., 2003; Reeves and DeWit, 2004), the reconstruction by Gibbons et al., (2013) shows an

unusual overlap for Sri-Lanka-Madagascar and India in their reconstruction (Figure 6.1). Similar

problems exist not only for the Elan Bank but also for the Jan Mayen microcontinent reconstruction as

described by Gernigon et al., (2012). Rotation poles important in global reconstruction cannot be

ignored completely but need to be modified with more available local constraints. Therefore, in this

study rotation poles calculated in the reconstruction of Reeves and DeWit, (2004) and; Reeves, (2009)

have been used as a fundamental framework to create a synthetic reconstruction, but without

completely relying on it. It has to be clarified that the presented reconstruction is not a global plate

reconstruction. This synthetic reconstruction model with focused approach calibrates the global plate

reconstruction models providing a detailed fine-tuned local model for East India-Elan Bank-Antarctica.


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Figure 6.1: 125 Ma reconstruction of the Indian Ocean with the Enderby Basin model (Gibbons et al., 2013). Note that a ~150 km overlap exists between Sri-Lanka -Madagascar and India. Coastlines are in grey, continent-ocean boundary is in yellow, thin red lines indicate location of seafloor-spreading magnetic anomalies with best fit for the Enderby Basin isochrons including M4 (126.7Ma) magnetic anomaly picks, which are indicated by small black circles, while pale black circles indicate the older Enderby anomalies M9 (130.2Ma). Continental fragments including the Elan Bank are shown in green. Note that the Elan Bank has been constructed against Mahanadi basin following Gaina et al., (2007).

Additionally, major problem with the individual plate reconstruction models (e.g. paleomap projects by

Scotese et al., 2003; Reeves et al., 2009) is created on a present day plate configuration, assuming the

rigid plate models. In reality, the plates have been stretched and thinned by extension during rifting

throughout their kinematic evolution to form passive margins. Therefore, unusual overlaps or gaps

occur when they are reconstructed into their pre-kinematic positions particularly in the case of

asymmetric extension. It has been described so far that the East India rifting is a hyper-extended

asymmetric rifting. This is the reason why the reconstruction based on 1000 m isobaths shows gaps and

overlaps in the models (Reeves, 2009, Figure 6.2). These ambiguities are resolved in the current

attempt by using crustal architecture of hyper-extended basins, particularly using a more precise map of

ocean-continent boundaries. Therefore, several modifications were made to India-Antarctica


244

reconstruction, including incorporation of the Elan Bank microcontinent into the kinematic and

paleogeographic reconstruction in this study as shown in Figures (5.38 to 5.42).

Figure 6.2: 145 Ma tight rigid plate reconstruction of the Indian ocean (modified after Reeves et al., 2008). The reconstruction is based on 1000 m isobath used as a proxy for ocean-continent transition. Note that a overlap exists between India and Antarctica in the present day Krishna-Godavari (KG) along Mahanadi region. Similarly, a gap exists between Cauvery Basin (CY) of India and Sri-Lanka and between Antarctica and Sri-Lanka. Coastlines are in grey, the Precambrian cratons are in pink. Precambrian sedimentary basins are in yellow.

It is important to note that the rates of plate movements are also inconsistent in different published

models (Scotese el al., 2003; Reeves and DeWitt, 2004; Collins and Pisarevsky, 2005; Reeves, 2009;

Seton et al., 2012). Overall, all of them address the spreading vectors and adjustments including major

plate reorganizations without capturing the minor adjustments due to regionally localized extension. For

example, in all the above mentioned models, there have been five major plate reorganization episodes

with a very broad age ranges identified. These include:

1. Separation of eastern Gondwanaland from western Gondwanaland (180-170 Ma),

2. Antarctica rift (134-130 Ma),

3. Antarctica drift (130-125 Ma),

4. Rotation of Sri-Lanka with respect to India (120 – 114 Ma) and


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5. Again another Rotation of Sri-Lanka with respect to India (114 – 100 Ma).

Although, these plate reorganizations occur in this study, their timing are durations are modified.

Particularly, the Antarctic rift and drift is constrained by extra datasets, which are discussed later in

more detail. Furthermore, Elan Bank is included.

Additionally, some of the crucial regional deformation events are missing in all the above mentioned

models. In East India, none of the plate tectonic models (Scotese el al., 2003; Reeves and DeWitt, 2004;

Seton et al., 2012) captured the movement caused by the prominent strike-slip deformation along the

Coromandal strike-slip fault (Figure 6.3). However, it is a critical component for understanding the

geodynamic evolution of Indian passive margins and Elan Bank microcontinent release mechanism from

India. This conspicuous absence may be either related to non-availability of seismic data or weak

magnetic anomaly signals in the Bay of Bengal.

Figure 6.3: 145 Ma tight rigid plate reconstruction of the Indian Ocean (Scotese el al., 2002; Reeves and 2008; Seton et al., 2012). All the displayed models show different rates and plate vectors direction. Note that none of these models capture the movement caused by the prominent strike-slip deformation along the Coromandal fault. The microcontinents are conspicuously absent in all models.


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Although, several workers including Gaina et al. (2007), Veevers (2009), Gibbons et al (2013) incorporate

the Elan Bank into their respective reconstructions. However, the position has been incorporated either

as purely based on magnetic anomaly interpretation (Gaina et al., 2007; Gibbons 2013) or to fill the gaps

remaining in the palinspastic restoration (Veevers, 2009).

Detailed magnetic anomaly studies in the Enderby Basin (Gaina et al., 2007) and India (Ramanna et al.,

2001) identifies the asymmetric nature of spreading between India and Elan Bank during early

Cretaceous ridge jump, which was probably a trigger for the separation of Elan Bank from East India. The

ridge jump as a result of plate reorganization in early Cretaceous is postulated in this model as based on

magnetic anomaly interpretation, which is subsequently mentioned by Gibbons et al., (2013) as well.

Gibbons et al (2013) reinterpreted the magnetic anomalies and proposed a new reconstruction model

for the Indian Ocean region. This model shows almost 50 km overlap between the Elan Bank and East

India ocean-continent boundary, which is interpreted from magnetic anomalies organized before the

ridge jump, which is considered to occur around 115 Ma (Figure 6.4). In this model, the earliest seafloor-

spreading anomalies are located close to Elan Bank, which is around 127 Ma.

Veevers (2009) used a more robust palinspastic reconstruction with quantitative elimination of the

intervening ocean floor. This reconstruction model uses seismic and gravity magnetic data to determine

the continent ocean boundary (COB) and thickness of the extended continental crust. This model also

provides geochronological connections across the India-Antarctica breakup trajectory. This

reconstruction involves a spreading system located east of 55°E in the Enderby Basin with initial

direction of NNW, and different from the N–NNE direction west of 55°E in the spreading system

between Antarctica and India. In this model, the Elan Bank was not reconstructed against East India.

However, it mentions that there is a gap in the palinspastic reconstruction between latitude 15°N to

20°N, which covers an area approximately of 550 km * 50 km (Figure 1.10). Veevers (2009) suggested

that this gap could be the location of Elan Bank microcontinent. Again, this gap is mostly located

northeast of the Krishna-Godavari Basin and close to the Mahanadi Basin and does not match with the

reported geometry of the Elan Bank microcontinent (Coffin et al. 2000; Wise et al., 2000; Coffin et al.

2002; Borissova, et al., 2003) as well as with the reinterpreted geometry of Elan Bank microcontinent in

the presented work. Therefore, the shape of Elan Bank as interpreted here and geometry of the

conjugate margin architecture provide a better control to position the Elan Bank between India and

Antarctica near the Krishna Godavari Basin. This placement further implies that the gap near the


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Mahanadi Basin as shown in both Veevers (2009) and presented work is needed to be further

investigated with more constrained data on Antarctica side.

Figure 6.4: 120 Ma reconstructions of the southern Indian Ocean (Gibbons et al., 2013). This model shows almost 50 km overlap between the Elan Bank and the ocean-continent boundary interpreted from magnetic anomalies before the ridge jump. Coastlines are in grey, continent-ocean boundary is in yellow, thin red lines indicate location of seafloor-spreading magnetic anomalies with best fit for the Enderby Basin isochrons including M4 (126.7Ma) magnetic anomaly picks, which are indicated by small black circles, while pale black circles indicate the older Enderby anomalies M9 (130.2Ma). Continental fragments including the Elan Bank are shown in green. Note that the Elan Bank has been constructed against the Mahanadi Basin following Gaina et al., (2007).

The plate reconstruction of Gaina et al. (2007) places the Elan Bank into the area in front of the

Mahanadi Basin and continuing into the northern end of the Krishna-Godavari margin segment, which

was also used in the latest model of Gibbons et al (2013) (Figure 6.4). The interpretation of magnetic

anomaly and ridge jump timing (Gaina et al., 2007) is also consistent in the presented work. In this

study, the improved physiographic boundary and the shape of the Elan Bank in combination with the

continental and proto-oceanic crustal boundary for the Eastern Indian offshore (Figures 5.38 to 5.42)


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and plate movement data obtained from Reeves (2008), allows to fit the Elan Bank further south close

to the Krishna-Godavari Basin. The north-western margin of the Elan Bank fits to the southern part of

the Krishna-Godavari basin (Figures 5.38 to 5.42). The south-western and western margin of the Elan

Bank fits against N-S striking dextral Coromandal strike-slip fault system. The southern side of the Elan

Bank is located exactly along the eastward continuation of the Cauvery rift zone. The northern side of

the Elan Bank matches with the proto-oceanic crust-continental crust boundary defined along the whole

Krishna-Godavari Basin. The differences from previously published information are also based on more

constrained timing arguments, and they will be discussed in detail in the following sections.

The reconstruction in current study also includes the global sea floor spreading datasets created by

analysis global magnetic anomalies to create an ocean floor age map Műller et al., (2008). An attempt

has been made here to restore the hyper-extended margins to their pre-kinematic position using the

crustal architecture of conjugate margins (Figure 5.37). The merits of the restorations will be discussed

in the subsequent sections. Interestingly, this process leads to several gaps between India and Elan

Bank, as this is not a tight-fit geometric reconstruction.

It has to be noted that presence of many gaps in reconstruction also potentially indicate a possibility of

more microcontinents present in the Indian Ocean, particularly around the Kerguelen Plateau. The

crustal structure of the Kerguelen Plateau indicates that perhaps include more microcontinents within it

representing a composite microcontinent as mentioned earlier (Rotstein et al., 2001; Borissova et al.,

2002; Bénard et al., 2010). This subject remains a good candidate for future research. The recalculation

of rotation poles has not been attempted in the presented study. This remains an important future area

of research, where new sets of rotation poles can be created to the fit Elan Bank a bit better as a

microcontinent separated during India-Antarctica rifting in global reconstruction models.

6.2 The physiographic boundary and crustal architecture of continental Elan Bank

microcontinent

The physiographic boundary of Elan Bank has been described previously based on potential field studies

(Frey et al., 2000; Coffin et al., 2002; Gaina et al. 2003; Bénard et al., 2010) seismic velocity structures

from refraction survey (Charvis et al., 1995; Operto and Charvis, 1996; Charvis, and Operto, 1999;

Borissova et al., 2003 Bénard et al., 2010) and reflection seismic interpretations (Borissova et al., 2003).


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In general, the Elan Bank has been identified as a promontory of the southern Kerguelen Plateau,

although the shape of the physical boundary varies a bit (Figure 1.3). These smaller ambiguities of the

shape of Elan Bank may not affect the overall understanding Elan Bank microcontinent separation

through a ridge jump but it significantly affects its conjugate fit with East India.

Commonly, the shape of Elan Bank can be derived from the bathymetry ranging from 1000 m to 2000 m

coupled with positive gravity anomaly signature in both Free air and Bouguer gravity maps against the

surrounding negative gravity anomaly (Frey et al., 2000; Coffin et al., 2002; Gaina et al. 2003; Bénard et

al., 2010) as shown in Figure (1.3), which is common for the present study as well. The Elan Bank

continental margin represents a hyper-extended rifted margin currently submerged at a 1000 m water

depth. Hence, the mass deficit due to water column between more or less uniformly thick oceanic crust

and adjacent microcontinent produces the contrasting high to low gravity anomaly. The mass excess due

to presence of shallow mantle below oceanic crust and deeper mantle depth beneath the continental

crust also contributes to the contrasting anomaly. While the Free-air data are indicative of density

contrasts within the crust, the Bouguer anomaly data tend to highlight the location of sedimentary

basins and flexurally compensated features such as rift shoulders. It should be noted that the

architecture of rift shoulders cannot be well understood as most of the seismic profiles are located on

distal margin except a single profile in the southern Elan Bank.

The positive anomaly on continental side coupled with negative anomaly on oceanic side has been used

as a representative for physiographic boundary of the Elan Bank microcontinent. This has been used for

interpretation along its entire margins. However, the described method to determine the

oceanic/continental boundary vis. a vis. the physiographic boundary of continental Elan Bank is a

simplified method, as it takes a uniform density value for oceanic and continental crust without

considering the effect of sediment loading. It is known that Elan Bank has almost 375 m of sediments

drilled in ODP well 1137 (Coffin et al., 2002; Ingle et al., 2002). However, in interpreted seismic sections

one can see almost more than 3-4 seconds (TWT) of sediments and volcanic rocks, which can be easily

translated into 2-3 km of sediments or more (Figures 3.20 to 3.23). The total volcanic thickness is highly

variable on the Elan Bank. This further contributes to the shape of the positive gravity anomaly. Despite

of the correction, described in the methods, the Bouguer anomaly map cannot indicate the presence of

exhumed mantle in front of the hyper-extended margin of Elan Bank, the external boundary of which

overlaps with relatively higher Bouguer gravity. Therefore, in this study, the gravity interpretation was

intergrated with seismic interpretation (Figures 5.6 to 5.9).


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Previously, the crustal structure of the Elan Bank has been described using the simultaneous inversion of

travel times of seismic arrivals (Borissova et al., 2003; Bénard et al., 2010). The results show mostly

almost 14km thick lower crust, with velocities as low as 6.6kms-1. This observation indicated that this is

not consistent with oceanic plateaus, which usually have high velocity due to heavy underplating.

Coverage the seismic lines are located only in the distal margin. As a consequence, they cannot address

total crustal thickness and question of underplating. However, seismic imaging suggests that the central

part of the Elan Bank is quite undeformed (Figure 3.24). Hence a thicker undeformed crust is expected in

the core of microcontinent similar to the situation in other microcontinent development models such as

that for the Jan Mayen microcontinent (Péron-Pinvidic and Manatschal, 2010). The detailed crustal

architecture interpretation of the margin has been summarized in a tectonic map of the Elan Bank,

which shows clear limits of different crustal architectural elements (Figure 5.23)

In the seismic interpretation, one of the typical problems was Moho calibration as data quality was not

quite good at deeper level, which also contained multiples (Figure 5.20 and 5.21). Therefore, the seismic

interpretation should be constrained with forward gravity modeling. However, due to the absence of

depth converted sections in Elan Bank, this could not be attempted in the present study. Therefore, the

seismic interpretation has been constrained by with gravity values extracted along individual seismic

profiles (Figures 5.6 to 5.9). Such correlation has been successfully applied to East India (Nemčok et al.,

2013b) and other passive margins (e.g. Rosendahl and Groschel-Becker, 1999; Odegard, 2003; Nemčok

et al. 2006; Dehler and Welford, 2013).

The seismic interpretation allows placing the Moho surface on the lowest strong reflector to

differentiate the oceanic, continental crusts and unroofed mantle according to their thickness and

seismic nature. While it is possible to end up with several possible scenarios using only gravity, the

unique scenario comes from constraints provided by interpreted seismic data. The synthetic approach

indicates that one needs to honor lateral changes of lithologies such as those caused by preferential

serpentinization, to sometimes explain the geophysical images of crustal transitions. Examples come

from studies done at different parts of the world (e.g. Meyers, 1996; Rosendahl and Groschel-Becker,

1999; Odegard et al., 2002; Odegard, 2003; Nemčok and Rosendahl, 2006a, b; Sinha et al., 2010;

Nemčok et al., 2012a). It has to be noted that the synthetic interpretation would be more effective if a

forward gravity modeling on Elan Bank attempt was made in this study. This would require a

reprocessing of Elan Bank seismic data filtering out the multiples and during the velocity modeling to


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convert the time-migrated sections to depth-migrated ones to get a more precise result. This remains an

open problem for future research.

In this study, the interpreted crustal type distribution and crustal architecture is further constrained by

the seismic profile marriage technique (Figures 5.25 and 5.26). The reflection seismic profiles from

conjugate margins have been chosen based on conjugate margin architecture and merged their external

proto-oceanic ends or external continental ends together (Figure 5.24). It has been taken care that the

lines have to intersect at conjugate positions near the merged oceanic/proto-oceanic or proto-

oceanic/continental crustal boundaries (see Rosendahl et al., 2005; Nemčok and Rosendahl, 2006a and

b). This technique has been quite helpful in the segments with roughly orthogonal rifting. However, it

was complex in the case of the Coromandal transform margin.

It is quite interesting to note that the hyper-extended crustal architecture model and gravity

interpretation correlate quite well. Furthermore, the profile marriage technique has been proved quite

useful in the study to define the gaps in crustal architecture, Moho surface and to reshape the

physiographic boundaries as per the margin architecture and rifting history. This correlation has been

further constrained by petrological information from the Elan Bank and conjugate East India (Figures

5.17 to 5.19). This kind of correlation was not previously attempted. Therefore, the preliminary

physiographic boundary and crustal architecture interpreted by Borrisova et al. (2003) has been

modified according to the margin architecture. The more precise physiographic shape of continental

Elan Bank provides a better control for understanding its geodynamic evolution and its release from the

East India.

6.3 The importance of the East Indian crustal architecture and ocean-continent transition in

reconstruction with conjugate Elan Bank

The tectonic interpretation of basement structures in seismic reflection profiles of hyper-extended

margins is particularly difficult due to its complex nature of ocean continent transition (Manatschal,

2004; Manatschal et al., 2007; Reston, 2007).

Previously, for East India, some authors (Rao et al. 1997; Radhakrishna et al. 2000; Subramanyam, and

Chand, 2006) have interpreted a limited amount of seismic images indicating the position of the ocean

continent transition, assuming that it is a classical ocean-continent transition based on the model by


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White and McKenzie (1989). Their interpretation primarily shows that the transition between the

continental and oceanic crust is a broad transition known as transitional crust. The subsided continental

crust is deformed by normal faults that dip seaward. This assumption was previously justified by the fact

that most profiles with shallow record length and basement were not clearly imaged. The seismic

interpretation of the ION\GXT grid with a record length of up to 16 sec two way travel time, along with

high resolution Reliance seismic profiles in East India shows significant variations in the orientation,

geometry and deep crustal structure along the passive margin (Figures 5.1 to 5.4 and 5.22).

The crustal architecture of East India has been interpreted through an integrated synthetic approach

(see Rosendahl et al., 2005, Nemčok and Rosendahl, 2006a and b for example). Using this approach, the

combination of interpretation of deep seismic reflection data both with forward gravity modeling

provided a better control in East India to determine crustal type distribution and crustal boundary

identification. It has been described that entire rifting history from rift initiation to continental breakup

went through several phases, including stretching, thinning, exhumation and spreading. The example of

the regional seismic profile shows various stratigraphic and structural features that define the passive

margin architecture. It particularly illustrates the seismic expression of these deeper structures below

the interpreted basement (Figures 3.1 to 3.13; 3.15 to 3.17). Such observations are critical as it leads to

a hyper-extended crustal architecture model for East India. This kind of attempt has been used in East

India and Elan Bank conjugate margin study for the first time.

The main features of hyper-extended passive margins include a multi-phase rifting history, distinct

structural and stratigraphic architecture in proximal and in distal margin, a corridor of exhumed mantle

as a result of extreme extension, presence of continental allochthons, a set of kinematically related

faults developed during different phases of deformation and finally the lithosphere beneath passive

margins that thins down to zero at the ocean continent transition seaward. (Manatschal, 2004;

Manatschal et al., 2007; Karner, 2008). The interpretation of seismic profile ION-GXT-1000 is consistent

with this interpretation, and it is the best section documenting this mechanism in the study area (Figure

3.15). The interpretation of all other ION-GXT and Reliance profiles also suggests that this mechanism

applies to nearly the entire East Indian margin. However, all of these features, for example, continental

allochthons are not present everywhere in East India. The reason is that the East India being the fast

extension system underwent a very quick continental breakup (Nemčok et al. 2012a). Another reason

may be the location of East India on the upper plate margin, whereas the continental allochthons are


253

mostly observed in lower plate margins in other parts of the world (Péron-Pinvidic and Manatschal,

2010; Sutra et al., 2013).

In current interpretation, the rifting style clearly shows the typical proximal-distal margin architecture

characteristics of a hyperextended margin. Overall, the proximal part of the continental margins can be

characterized by extensional deformation that results in the formation of rift units connected through

transfer zones or pull-apart basins with marginal highs. The distal parts are divided into segments where

the upper-crustal break-up was controlled by normal faulting and strike-slip faulting. Both can be

observed in the tectonic element map (Figure 5.1).

At this point, the rift architecture has been further discussed through the geometry and genesis of the

faults. The first-order basin-bounding faults effectively delineate the limits of proximal to distal margins.

The high-angle SW-NE trending faults are crustal-scale and represent the stretching phase of

deformation. These faults cut through the entire syn-rift stratigraphic section and are passively

onlapped and buried by a large amount of post-rift sedimentation (Figures 3.1 to 3.3, 3.5, and 3.10 to

3.13). Additionally, evidence for small normal offsets of the picked top basement reflector can be

observed as well. This small offset is best observed within the more prominently tilted fault blocks, or

within the deeper graben areas (Figures 3.10 to 3.13). Low-angle intra-basement reflections are

interpreted as detachment faults (Figure 3.10, 3.15) based on the hyper-extended crustal architecture

model developed for the Iberia-Newfoundland margin (Reston 1996; Manatschal and Bernoulli 1999;

Manatschal et al, 2007). Although, the overall margin architecture that varies from rifted margin to

sheared margin has been described by previous authors (Biswas, 2003; Subramanyam, and Chand,

2006), the detail geometry of the margin provided later (Nemčok et al 2007; Nemčok et al 2013b; Sinha

et al., 2010;) and further in this study. It is similar to the most of the hyperextended margins in the

world (Manatschal et al. 2007; Karner, 2008; Péron-Pinvidic and Manatschal, 2010; Mohriak and Leroy,

2013; Nemčok et al 2012a; Sutra et al., 2013).

Different margin segments in East India range from orthogonal rift dominated margins (e.g. Krishna-

Godavari and Cauvery segments) to strike slip dominated margins (Coromandal segments) (Figure 5.1).

The margin segmentation is the result of pre-existing crustal inheritance and kinematic variation in

rifting style along the margin prior to continental breakup (Nemčok et al 2013b) (Figure 5.1). The

complexity further increases due to presence of volcanics. The geophysical studies earlier proposed the

existence of two end-member margin types viz. “volcanic” type where magmatism was widespread


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during continental breakup, and a “non-volcanic” type where magmatism was much more limited

(White et al., 1987; Kelmen and Holbrook, 1995).

Conceptually the ocean continent transition at volcanic margins is commonly a simple transition

between continental crusts through increasing degrees of intrusion, until the point is reached where the

crust consists entirely of intrusive igneous material of mostly basaltic composition. This initial oceanic

crust is typically thicker than normal one and may reach a thickness of approximately 25 km (White and

McKenzie, 1989). This transition is typically 20-100 km wide and may be characterized geophysically by

changes in P-wave velocity structure, density and magnetization (Minshull, 2009). The seafloor-

spreading magnetic anomalies are identified over the earliest-formed oceanic crust. The seaward limit

of continental crust may be very difficult to define at such margins, because heavily intruded upper

continental crust may have a bulk composition very similar to oceanic crust and, therefore, physical

properties similar to upper oceanic crust.

The typical non-volcanic margins are usually characterized by a lack of magma in the ocean continent

transition. The most common phenomenon of magma-poor margins is the exhumation of mantle rocks

to the seafloor during the continental breakup (Manatschal and Bernoulli 1999; Manatschal, 2004;

Husimans and Beumont, 2005; Manatschal et al., 2007; Karner, 2008; Minshull, 2009; Nemčok et al.

2012a & b). At these margins, the ocean-continent transition may be less well-defined. There a certain

agreement exists regarding the formation the non-volcanic margins based on observations from seismic

and field analogues in the Alps. However, there is no general agreement in definition of the volcanic

margin. It is extremely difficult to explain its genesis and what truly distinguishes it from its non-volcanic

counterpart at crustal and lithospheric scales (Geoffroy, 2005). Therefore, the classical concept of

volcanic and non-volcanic margins seems more problematic to explain as observed in other parts of the

world (Gernigon et al., 2012; Manatschal and Karner, 2012).

This study, however, differs in concept of end member paradigms of volcanic and non-volcanic margins

with using the new observations, particularly around East India. Potentially, it infers that passive margin

formation involves around complex transition between rifting, breakup and seafloor spreading. The new

insight indicates that the both kinematics of rifting and additional magma, required to break the crust in

combination develops the final passive margin architecture (Manatschal and Karner, 2012). This rifting

dynamics is therefore, not leads to a single end-member of volcanic or non-volcanic margin. In this

study, one can see the complex margin architecture is a result of interaction of different phases of

breakup and multiple ridge jumps associated with. It is possible that the rifting initiates by hyper-


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extension. However, the Breakup is linked with excess magmatic event overprinting hyper-extended

crust. Therefore, this study advocates for a common margin type, where the ocean-continent transition

can be classified as magmatic or non-magmatic based depending on magma required to break the crust.

Most interestingly, the microcontinent association has been proved to be critical to define volcanic and

non-volcanic margins as new end member paradigms.

The deeper structures interpreted in the profiles also need a special focus in the current discussion. It

can be observed in most of them that the lithosphere thins down to zero near the ocean-continent

transition as a result of hyper-extension (Figures 5.2 to 5.4). The upper crust-lower crust boundary

interpretation is based on relative change in reflectivity. The upper crust-lower crust boundary imaging

in East India is similar to that at other margins (Christensen, and Mooney, 1995; Rosendahl and

Groschel-Becker, 1999; Odegard, 2003; Nemčok et al. 2006; 2012a; Manatschal et al., 2007; Dehler and

Welford, 2013; Mohriak and Leroy, 2013). In East India, it is particularly characterized by relatively

transparent upper continental crust with listric normal faults soling out at middle to lower crustal levels

(Figures 3.1 to 3.13, 3.15-3.17). The lower crust is usually highly reflective, with sub-horizontal

reflectors, and pinches out in an oceanward direction. However, the lower crust in this margin is thicker

than that at other margins, which is reported by Nemčok et al (2012a).

Clear imaging of continental as well as oceanic Moho significantly contributes to the improved

interpretation (Figures 3.15, 3.16). The crustal type distribution and boundaries are, therefore,

restrained by the Moho interpretation. In seismic reflection profiles, the ocean-continent transition

commonly lacks a clear Moho reflection, in contrast to the oceanic and continental crusts, where the

Moho reflections are quite clear. It is identified by looking for the evidence of tilted, syn-tectonic

sediment packages, low basement relief and continental allochthons as observed in typical Iberia-

Newfoundland margins (Reston 1996; Manatschal and Bernoulli 1999; Manatschal et al, 2007; Reston,

2007; Minshull, 2009; Whitmarsh and Manatschal, 2012). Similar observations are also noted in the

current interpretation. It has to be noted that the Moho interpretation is clearer in profiles, which are

free from post-rift hotspot related volcanism (Figures 3.1 to 3.5, 3.12, 3.15, and 3.16). The Moho

imaging in few other profiles with this effect is disturbed due to thick pile of volcanics (Figures 3.15 and

3.17). However, due to presence of a relatively intertwined seismic grid, it is possible to correlate Moho

throughout the region.

The distribution of different crustal types deterministically indicates the presence of a corridor of

exhumed mantle or proto-oceanic crust in the ocean-continent transition. This is clearly visible in


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seismic imagery as well. This feature has been a matter of debate for a long time due to its variable

characters. However, the proto-oceanic zone can be uniquely characterized by deeply subsided,

anomalous crust at ocean continent transition. Its thermodynamic behaviour is more similar to normal

oceanic crust than the continental crust. In reflection seismic data, it appears as a chaotic package with

numerous discontinuities, which may be due to the fracturing and shearing of the mantle during

exhumation (Figure 3.15). During exhumation phase when the mantle is unroofed at the continent-

ocean transition, it subsides quickly and comes directly in contact with sea water. The mantle mostly

composed of the peridotite can be transformed into serpentinite owing to chemical reactions with

downward-percolating sea water (Manatschal and Bernoulli 1999; Odegard et al. 2002; Odegard, 2003).

The percolation is allowed by intensive fracturing, known from areas with identified unroofed

continental mantle (Manatschal, 2004).

Presence of low-amplitude magnetic anomalies is quite common in the ocean continent transition at

magma-poor margins and it may be weakly linear (White and McKenzie, 1989; Kelmen and Holbrook,

1995; Srivastava et al., 2000; Manatschal, 2004; Sibuet et al., 2007; Minshull, 2009). The presence of

these anomalies indicates the presence of rocks with higher magnetisation than that of the adjacent

continental crust. The process behind the formation of magnetic anomaly has a number of contrasting

interpretations. These anomalies were either originated during formation of normal oceanic crust as a

result of sea floor spreading. Another explanation suggests that during the serpentinization of the

peridotite, olivine is transformed to magnetite. However, the extent of magnetite transformation may

vary due to the migration pathways of altering fluids, which may have some consequences on the

magnetic anomaly occurrence (Srivastava et al., 2000; Sibuet et al. 2007). These contrasting hypotheses

clearly indicate that the presence of magnetic anomaly cannot be the only determinant factor for

oceanic crust identification.

The lithospheric thinning in East India is accompanied by upwelling of the underlying asthenosphere,

which initiates the decompression melting, followed by formation of oceanic crust (Figure 5.2 to 5.4,

5.25 and 5.26). The ocean-continent transition may be defined in a variety of ways. In this study, a

relatively narrow definition has been adopted that focuses on the region between unequivocal thinned

continental crust and oceanic crust formed by “normal” seafloor spreading processes. There is a lack of

consensus on to definition of normal oceanic crust. In case of East India, however, the average thickness

of normal oceanic crust formed by organized sea-floor spreading is ranging from 4 to 7 km, which falls


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within global average (see Christensen, and Mooney, 1995; Kearey et al., 2009). However, the thickness

of the oceanic crust with hotspot volcanic addition increases up to 10 km near the 85E Ridge.

The detailed analysis of all geologic and geophysical data provides two plausible solutions to explain East

Indian crustal architecture. Two alternative hypotheses have been proposed for the nature of the ocean

continent transition along the East India margin. The first hypothesis is concordant with hyper-extended

crustal architecture interpretation where the ocean continent transition occurs through exhumation of

lithospheric mantle forming the proto-oceanic corridor. The second one assumes that the transition

from continental to oceanic crust is classical and there is no need to interpret a proto-oceanic corridor in

between.

The forward modelling results based on only gravity data (Bird 2009) indicate that two thirds of the

modelling results support both alternative hypotheses. The rest of the modelling results are mostly

indeterminate to reach a clear conclusion. Overall, the Free-air and Bouguer gravity data across the

margin (Figures 3.26 and 3.29) have following common features. In the modelled profiles (Figures 5.10

to 5.16), the shelf break region would overlap with a positive anomaly, reaching from about 0 mGal to

27 mGal. Here, all models indicate a thinner continental crust. However, the profile along ION-GXT 1000

convincingly allows interpreting a presence of the proto-oceanic crust to fit the calculated gravity curve

to the measured one (Figure 5.10). If, the same profile uses the classical oceanic crust model, it grossly

does not fit these curves while keeping the geological interpretation intact. The other models (Figure

5.11, 5.12 and 5.16) show similar observations. Some interesting observations can be pulled out if one

can observe the modelled profile without proto-oceanic crust (Figure 5.13). The IONGXTprofile-1000

(Figure 5.13) shows that a gravity minimum on the oceanic side of the crustal boundary reaches only

about -80 mGal, while the gravity minimum shown in modelled profile 1290 (Figure 5.16) reaches to

about -105 mGal. The curve further tries to climb up by -25 mGal and -35 mGal in profile-1000 and 1290

in respectively. This indicates the rising of Moho as a result of crustal extension. So, any presence of a

high-density body, possibly mantle, would be imaged as a hump on a curve climbing from its minimum

near the continent–ocean boundary to the mentioned slightly negative values on the oceanic side.

Using the modelling, it is not possible to determine whether the mantle reaches the surface, particularly

in modelled profile 1290 (Figure 5.16). Such interpretation has to rely on seismic interpretations only.

These two tested end-member alternatives, each with several favorable and unfavorable arguments,

have profound implications on the width of the proto-oceanic corridor, geometry of the ocean-

continent transition boundary and their role in placing the Elan Bank against East India.


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Interpretation of the crustal boundaries from all of the interpreted seismic profiles produces an

interesting outcome (Figure 5.22). It shows a notable along-strike variation in the width of the proto-

oceanic corridor. It is interesting to compare variations in width of this corridor with kinematic control

over the break-up for different segments of the continental margin (Figure 5.1). Such a comparison

roughly indicates that the width of the interpreted proto-oceanic crust corridor varies along the margin,

possibly due to the variable rifting style prior to break-up (Nemčok et al 2013b).

The map of the zone of exhumed continental mantle or proto-oceanic crust corridor in the East Indian

offshore (Figure 5.1). This was not previously reported in the maps interpreted from gravity data,

magnetic anomaly pattern, and a very limited amount of relatively shallow reflection seismic data (Rao

et al. 1997; Radhakrishna et al. 2000; Ramana et al. 2001; Veevers 2009), This would not be possible

without the ultra-deep reflection seismic ION-GXT grid, which allows to image the hyper-extended

crustal architecture interpretation.

6.4 Time constraints for tectonic events and rift evolution at hyper-extended margins

The separation of the Elan Bank from East India has been constrained through the age determination of

different tectonic events. These timing events are also critical for fitting the Elan Bank to East India.

Especially the use of borehole data provides important timing data, which has been used in this study.

The study of rift event timing along the Iberia-Newfoundland margins leads to an interesting

observation. It shows a significant variation in age of the syn-extensional sequences across the margins

as well as along the margins (Manatschal, 2004; Tucholke and Sibuet et al., 2007, Péron-Pinvidic et al.,

2007; Sutra et al., 2013). The interpreted ‘breakup’ unconformity separating synrift sediments from

post-rift sediments coincides with the mechanical decoupling of the lithosphere rather than the first

occurrence of the oceanic crust (Péron-Pinvidic et al. 2007). This observation shows that the use of

“breakup unconformity” may not be a valid proxy for determining the age of the continental breakup. It

has been already discussed that due to various reasons, the oldest magnetic anomaly identification

cannot be a proxy for the age of continental breakup as well.

In East India, the situation is more complex than that in Iberia-Newfoundland due to various reasons.

First, the lengths of the conjugate margins are comparatively larger, and segmented to six margin

segments. Second reason is the poor data control documenting the time constraints in the conjugate


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Antarctica and Elan Bank itself. The only detailed data control comes from the East India margin

borehole data. It has good quality biostratigraphic data. However, the data is not well distributed along

the margin but rather clustered in regions with exploration activity (Figure 5.27). The data is present

from proximal to distal margin, which provides reasonable margin coverage (Figure 5.28). The third and

most important reason is the microcontinent separation from the East India, which was associated with

multiple breakup events along the margin.

Age constraints for rift activity in the proximal margin are domains are based with identification of

typical rift indicators including growth packages against rift faults. The onset of rifting is poorly

constrained along the entire margin except the Krishna-Godavari Basin. The onset of rifting in the

Krishna-Godavari Basin is dated as Toarcian (174 Ma). However, the Cauvery Basin shows a thicker

synrift section, which is completely undrilled. The oldest drilled sediment in Cauvery Basin is reported as

Late Jurassic in onshore (Chaudhury et al., 2010). Therefore, it has been assumed that brittle faulting

had to initiate in both regions after Toarcian. Implying from this, the onset of rifting in the future distal

margin is thought to be initiated in both regions in the Early Cretaceous, as indicated by the occurrence

of Barremian sediments (127-125 Ma).

The unconformity in the proximal margin is associated with an age gap ranging from Beriassian to

Hauterivian (145 to 132 Ma), which is maximally 13 Ma long age gap. This is true along the entire margin

from Cauvery through Coromandal to Krishna-Godavari segments. The Hauterivian strata represent the

youngest syn-extensional unit defined by growth structures mostly in stretching dominated decoupled

domain in the proximal margin (Figures 3.10 to 3.13). However, this relationship in the distal margin is

different. In distal margin, there is not a typical growth signature as it is mostly defined by the

decoupled domain dominated by thinning. The very small fault bounded basins shows no visible

unconformity but a conformable transition from shallow to deep water by early Aptian (123 Ma)

(Figures 5.27 and 5.28). The age determination of exhumation is doubtful due to the absence of well

information in the exhumation domain. The observations suggest that the first deep marine enveloping

surface has early Aptian age (122-120 Ma) along the entire margin (Figure 5.27). The deep marine

sedimentation is a clear indication of rapid subsidence along the margin.

The age of breakup and first accretion of the oceanic crust is always a matter of debate in conjugate

margin studies (see Manatschal, 2004; Manatschal et al., 2007; Gaina et al., 2007; Péron-Pinvidic et al.,

2007 Sibuet et al., 2007; Tucholke and Sibuet, 2007 Műller et al., 2008; Minshull, 2009; Veevers 2009;

Sutra et al., 2013). In East India, most of the previous studies indicate the breakup age using the


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magnetic anomaly patterns. The ages proposed for breakup range from Valanginian to Hauterivian with

a wide variety of age ranges. These include 132.5 Ma (Powell et al., 1988), 134 Ma (Ramanna et al.,

2001), 130 Ma (Gaina et al., 2003), 132 Ma (Reeves. 2008) and 132 Ma (Veevers, 2009). Although not

exclusively but in all these studies, the breakup is claimed to be synchronous with the first accretion of

oceanic crust. This assumption has been subsequently used as a primary guiding tool in all plate

reconstructions (see Scotese et al., 2003; Reeves and DeWit, 2004; Reeves, 2008; Veevers, 2009). The

identification of M11 magnetic anomaly (Ramanna et al., 2001) indicated the oldest oceanic crust in the

Bay of Bengal in the offshore Krishna-Godavari area, thus indicated the breakup must be around

Valanginian (136 Ma). Now the seismic and isostatic gravity anomaly study (Nemčok et al., 2013b)

shows that this anomaly is coincident with a solitary volcano (Figure 3.31) in the ocean continent

transition zone not a representative of the normal oceanic crust.

As already discussed, this assumption may create a fundamental flaw in constraining the time events at

hyper-extended margin. In this study, the true continental breakup is defined as the time when the

deformation and magmatic activity finally localized at the end of extensional deformation history and a

steady-state organized sea floor spreading center is created. From this moment onwards the lithosphere

is truly separated and continental drift is initiated. Thus, the limits of rifting and drifting are also defined

in this way.

Considering, the theoretical background of hyper-extended margins, this study, proposes the first deep

marine transgression related to rapid subsidence, which seals all syn-extensional units from proximal to

distal margin as a proxy for dating the continental breakup. This claim is based on the observations in

other margins particularly Iberia-Newfoundland (Sutra et al., 2013) and Southern Australia (Whitteker et

al., 2010). This has proved to be a good proxy, which has been proposed in this study. This particular

event honors the kinematics of rifting and passive margin development. It also satisfies the fact that the

rapid post breakup subsidence ends the extensional process, followed by isostatic uplift (Buck, 1986;

Weissel and Karner, 1989; Nemčok and Rosendahl, 2006a; Manatschal et al., 2007; Tucholke and Sibuet,

2007). As the post breakup deep marine enveloping surface is also easy to identify based on marine

fossil abundance. In reflection seismic data, this surface can be easy identified by the presence of

downlaps. The base of the downlap surface has been mapped as K30 surface (Figure 3.1 to 3.13, 3.15 to

3.17).

It has to be noted that the breakup timing in the Cauvery Basin is bit complicated. As mentioned in the

previous chapter, the Cauvery Basin has underwent two breakups, first one owing to Antarctica breakup


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during Valanginian (132 Ma) and the second one along the Coromandal transform, which subsequently

accommodated a lateral clearance of the Elan Bank which was passing though the Cauvery Basin along

the Coromandal strike slip fault and its extension to the eastern margin of Sri-Lanka during Aptian (122

Ma). Large compressional deformation structures are observed in outboard margin of the Cauvery Basin

where the synrift section itself is deformed (Figures 3.4 and 3.5). It has to be noted that the thicker

synrift section suggests that it is probably a proximal margin setting. This observation clearly indicates

the strike slip motion along the eastern margin of Cauvery basin is affecting the older synrift section. It is

possible that the Elan Bank separation along the sheared margin created local transpressional domains,

responsible for observed deformation pattern. It has to be noted that the deformation is quite localized,

which further supports the transpressional deformation along the sheared margin. The unavailability of

seismic data on the eastern margin of Sri-Lanka makes this assumption incomplete at this moment.

However, if the Free Air and Bouguer gravity map (Figures 3.26 and 3.29) are compared, it is quite

reasonable to assume the eastern margin of Sri-Lanka is a sheared margin as well.

The two breakup events indicate that two isostatic uplifts must have affected the Cauvery Basin. These

are not directly dated using fission track studies, as there are no samples available at this moment.

However, as discussed earlier, the sediment response pattern and considerable missing sections in the

stratigraphy of some wells, (Figures 5.28, 5.34 and 5.35) documents an additional possible isostatic

uplift in this region, which makes it different from the Krishna-Godavari region, where there is a single

isostatic uplift is observed (Figures 5.34 and 5.35). This observation leads to inference about the

complexity in passive margin evolution, especially when the microcontinents are involved. It has a

strong implication in thermodynamics of the margin evolution as well as the influence on economic

issues including exploration for hydrocarbons.

The use of the discussed constraint on continental breakup provides a very narrow time window for

exhumation, which is assumed to take place approximately between 127 and 120 Ma and 125 and 122

Ma. This leads to almost 3 to 7 Ma time window for exhumation. The timings are variable along different

margin segments e.g. from orthogonal to strike-slip segments. This narrow time window for exhumation

also leads to relatively fast extension along the East Indian margin, which further supports the fast

extensional model proposed for East India (Nemčok et al 2012a). Considering these observations, it is

clear that the age of the first oceanic crust formed by organized sea-floor spreading initiates during early

Aptian (between 122 to 120 Ma), depending on margin segment architecture. It has to be noted that the

Cretaceous long-normal polarity super-chron ranges from 125 Ma to 83 Ma and the breakup timing


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estimated in this study overlaps within this time span. This piece of information also justifies why clear

magnetic anomaly pattern and fracture zone geometries are difficult to recognize in front of this margin.

Therefore, diverse opinions, so far, existed in the reconstruction of the early history of drift between

India and Antarctica and Elan Bank separation. This also signifies why there is a discrepancy in Elan Bank

microcontinent placement against East India, if the reconstruction is purely based on magnetic

anomalies and rotation poles.

It has to be noted that these data constrains are mostly limited to East India only. The data constraints

in the Enderby Basin and Elan Bank region has been only based on loose correlation of magnetic

anomalies (Gaina et al., 2007; Gibbons et al., 2013) and ocean floor age (Műller et al., 2008). It is

reported that the break-up between Elan Bank and Antarctica initiated at about Hauterivian (130 Ma)

and subsequent spreading ceased at about Barremian (124 Ma) in the Enderby Basin conjugate to Elan

Bank on the Antarctic side (Ramanna et al., 2001; Gaina et al., 2007; Gibbons et al., 2013) (Figure 2.19;

Figure 6.5). The end of active spreading ridge between the Antarctica and Elan Bank in the Enderby

Basin is a crucial piece of information, which is commonly not difficult to identify in magnetic anomaly

patterns. So, this information remains as one of the key anchor points for timing. Interpretation using

these datasets in combination with dataset available for East India provides a better control in

understanding the ridge jump events during microcontinent separation. In East India, an early Aptian

situation happens to be very close to the continental breakup (122 Ma) with a permissible error margin

(Figure 5.27). The coincidence of the timing can be used for implied timing of the spreading center jump

from the region between the Elan Bank and Antarctica to the region between the Elan Bank and East

India.

The data and observations presented above allow for a logical age determination of each major event

associated with the evolution from the stretching dominated proximal margin deformation through

thinning dominated distal margin deformation to the first exhumation of mantle rocks and lithospheric

breakup in region of the future East India-Elan Bank-Antarctica passive margins.


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Figure 6.5: 1 min Satellite-derived free-air gravity map for the Enderby Basin (Gibbons et al., 2013). The magnetic isochrons are reinterpreted after Gaina et al., (2007). The isochrons M0 are shown as thick black line, M2 as thick brown line, M4 as thick red line, and the 115 Ma extinct ridges are as thick blue dashed line. Younger synthetic isochrons are shown to highlight the change in direction of sea-floor spreading 100Ma. Dark grey line represents the COB.

6.5 Elan Bank microcontinent release mechanism

One of the primary objectives of this study was to determine the microcontinent development model at

passive margins using the Elan Bank as a case study. As mentioned, at the very beginning that there are

two alternative models exist for microcontinent development, one is the plume refocusing (Műller et al.,

2001) (Figure 1.1), another one being the competing rift zone hypothesis (Péron-Pinvidic and

Manatschal, 2010) (Figure 1.2). Although, some other alternate examples of microcontinent release

mechanism include transpression e.g. Ghana ridge case in equatorial Atlantic (Nemčok et al 2013c),

repeated ridge jumps e.g. Seychelles in Indian ocean (Coiller et al., 2008; Misra et al., pers comm.),

change in stress regimes in the multiple times e.g. Lomonsov Ridge in Arctic (Minakov et al., 2013, T

Doŕe, pers comm.). All these examples and hypotheses are signifying that microcontinent development


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models can be grouped into two distinct categories. The first one that involves mantle plumes, the

second one, does not require plume or its role is quite secondary.

Another common phenomenon, which is observed during microcontinent separation, is the ridge jump.

It has been observed that under certain conditions there is reorganization in sea-floor spreading, where

active spreading is gradually overtaken by a newly-formed spreading center (Goff and Cochran, 2006).

One of the other critical aspects of the current study is to understand the geodynamics of the ridge

jump during the microcontinent separation.

The plume refocusing hypothesis is built upon the active rift mode and a hypothesis by Steckler and ten

Brink (1986). The hypothesis states that minimum yield strength occurs along the edges of a rifted

margin because the crust there is already weakened and hotter than a normal crust owing to a

conductive heat flow from the adjacent rift. Therefore, when the newly-formed passive margin moves

over a hotspot, this facilitates renewed rifting at the edge of the margin. The renewed rifting under the

influence of the hotspot causes another continental breakup. Subsequently, the active ridge moves

towards hotspot causing a ridge jump. This gradually isolates the microcontinent, which later becomes a

part of the newly formed plate. The ridge jump toward hotspot causes further asymmetries in oceanic

crustal accretion, resulting in the excess accretion and series of extinct ridges, even within the

microcontinent (Figure 1.1). This model explains a fast separation of microcontinents well and the

common association of plumes in the microcontinent formation, volcanism and asymmetric rifting,

which have been observed with many microcontinents around the world. The examples include Jan

Mayen and Seychelles (Műller et al., 2001) and Elan Bank (Gaina et al., 2007).

Several observations from different microcontinents and numerical models in passive margin suggest

that the presence of a mantle plume is not essential to invoke renewed rifting (Huismans and Beumont,

2005; 2011; Collier, et al., 2008; Simon et al., 2009; Péron-Pinvidic and Manatschal, 2010; Sutra et al.,

2013). These observations and numerical modeling results indicate existence of alternative competing

rift zone hypothesis. Depth dependent extension hypothesis suggests that the lithosphere, which

undergoes extreme thinning during passive margin formation, is already weak during the final stages of

rifting. Several competing rift zones simultaneously thin the lithosphere in various places forming

several weak zones known as crustal wounds (Nemčok et al., 2012a). The typical examples of such weak

zones or crustal wounds come from Galicia Interior basin (Hooper et al., 2007), Flemish Cap (Péron-

Pinvidic and Manatschal, 2010) and Laxmi Basin (Collier et al., 2008).


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Prior to continental breakup, several such crustal wounds are developed through a series of detachment

faults during the thinning and exhumation phases. Additionally, the deformation becomes localized

during this time at the edges of strong crustal bodies. This results in the individualization of the future

continental ribbons, referring to zones that do not further extend and thin (Péron-Pinvidic and

Manatschal, 2010). A conjugate system of faults accommodating the thinning separates relatively

stronger lithospheric blocks including the future microcontinent in their footwall from hanging wall

blocks (Figure 1.2).

The kinematic competition between each crustal wound that wanted to host the crustal break-up

eventually leads to the formation of a line of breakup by establishing a linkage. It is to be noted that the

age of the continental breakup is not the same along the segmented passive margin (see Chapter 5). The

initial breakup may be localized in segments with discrete and rudimentary sea floor spreading as

observed in Woodlark Basin, Papua New Guinea (Goodliffe and Taylor, 2007). Although, this is a three-

dimensional problem, it is shown along a profile (Figure 1.2, Figure 5.37) how the complex detachment

system helps in creating several weak zones, which serve as possible future breakup locations. Now

once an initial line of breakup is established and followed by a discrete initial sea-floor spreading, the

renewed or continued rifting within one of the competing rift zones further propagates, depending on

structural, compositional and thermal inheritance as well as on the strain rate (Péron-Pinvidic and

Manatschal, 2010). As a consequence, there is a serious reorganization in breakup localization, which is

followed by a renewed linkage between competing rift zones. Thus, some of the discrete sea-floor

spreading centers related to the initial breakup geometry eventually become inactive when the final line

of breakup and a continuous subsequent sea-floor spreading center system becomes established. The

spreading asymmetry may further accentuate a shutting of one ridge and causing the ridge jump. As the

spreading asymmetry causes an irregular ridge propagation, both the newly formed and older ridge

simultaneously remain active for a short period of time prior to ridge jump (Goff and Cochran, 1996). As

a result, the continental block, which is trapped between the initial and final breakup leaves with one of

the newly-formed plates, becoming a microcontinent (Figure 1.2 and 5.37).

There are several hypotheses explaining a ridge jump. The mian ones are plume assisted ridge jump

(Brozena and White, 1990; Mittelstaedt et al., 2008) and ridge jump due to irregular ridge propagation

(Sempéré et al., 1995; Searle et al., 1998). Numerical modeling by Mittelstaedt et al. (2008, 2011)

suggests that the ridge jump occurs due to ridge-plume interaction. The off-axis magmatism due to

hotspot causes shearing at the base of the lithosphere, below which the magma chamber acts as a


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reheating agent (Figure 6.6). This reheating of the lithosphere is controlled by the balance between the

magmatic heat flux, advection due to plate spreading and surface heat flow (Mittelstaedt and Ito, 2005;

Mittelstaedt et al., 2008). If the heat is sufficient enough with higher magmatic influx, the lithosphere

gets thinned and a new oceanic rift forms. Gradually it becomes the principal ridge axis and spreading at

the former axis ceases.

Figure 6.6: Conceptual model of magmatic heating of the lithosphere (Mittelstaedt et al., 2008). A source of melt beneath the hotspot (dashed ellipse) provides magma which passes through the lithosphere in a magma transport zone (grey box) and proceeds to thermally weaken the lithosphere. The rate of heating is influenced by magma flux, spreading rate (U, black arrows) and surface heat flow. Magmatism also heats and thermally weakens the ridge axis.

The other hypothesis suggests that spreading reorganizations are triggered by the spreading asymmetry

and irregular ridge propagation due to change in plate motion. Generally, small changes can be

accommodated by transform fault adjustment but if the changes are significant they cause a drastic

response by abandoning one spreading ridge and formation of a new one (Goff and Cochran, 1996). The

abandonment of one spreading center and creation of another new one can be achieved through ridge

propagation and coincident initiation of rifting along the length of the new axis (Sempéré et al., 1995;

Goff and Cochran, 1996). The mechanism behind ridge propagation can be well explained by fracture

mechanics (Searle et al., 1998). Faults at the active ridge grow by along-axis linkage. In the inside corner,

they also link in the axis-normal direction by inward curving to meet the next outer (older) fault. This

leads to wider-spaced faults compared to segment center or outside corner. The non-transform offset is

characterized by faults that are highly oblique to the spreading direction, and show cross-cutting

relations with ridge-parallel faults suggesting along-axis migration of the offset (Searle et al., 1998).

Where there is an asymmetric spreading, the shape of such inward-curving faults must continuously

change, but is assumed to cyclically return to its original shape by discrete inward ridge jumps (Wilson,

1990). Hence, it is also possible that the spreading asymmetry alone can trigger irregular ridge


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propagation and, therefore, ridge jump can occur without a significant magmatic thinning due to

hotspot actively.

In case of the Elan Bank, if the plume refocusing model holds true, then the initial breakup in the east of

the Elan Bank is abandoned in favour of Krishna-Godavari basin due to the arrival of the Kerguelen

plume (Figure 6.7). The process was initiated when India-Elan Bank-Antarctica was undergoing extension

as a single plate. First, Antarctica is separated from India-Elan Bank, forming two separate plates and a

new spreading center is created in between. The approximate timing of the onset of new spreading

center is Hauterivian (around 132 Ma ± 3 Ma). However, a major thermal reorganization took place

immediately after the breakup due to the Kerguelen hotspot. A new breakup area was focusing due to

thermal instability in between the weak zone between India and Elan Bank, which was already formed

during the India-Antarctica rifting. The renewed rifting west of Elan Bank in the Krishna-Godavari Basin

further continued. The renewed rifting in Krishna-Godavari Basin west of the Elan Bank caused a new

spreading center development. Therefore, the ridge jump was like due to Kerguelen hotspot

interference and the spreading center has been shifted between India and Elan Bank, thus, transferring

the Elan Bank to Antarctic Plate. Old spreading center between Elan Bank and Antarctica became

defunct and Elan Bank remained as a micro-continent.

This model is essentially an active rift model (Turcotte and Emerman, 1983), which explains the quick

and fast breakup with ridge jumps. In such case, the microcontinents and their conjugate margins are

expected to have serious syn-rift and syn-breakup volcanism as well as major underplating due to plume

effect. This, however, is lacking in East India and Elan Bank and is unknown in Antarctica. In the case of

East India, the onset of Kerguelen plume activity in onshore associated with Rajmahal trap, which is

almost 800 km away from the reconstructed Elan Bank position (Figure 5.36). Quite accurate

radiometric dating of the volcanics in the Rajmahal trap clearly indicates the emplacement timings at

about 117 ± 0.2 Ma (Coffin et al., 2002; Kent et al, 2002; Ghatak and Basu, 2011), which is clearly a post

breakup timing. As mentioned earlier, the offshore expression of the Kerguelen hotspot trail is

represented by the 85°E Ridge. The radiometric dating is not available for the ridge, but the

reconstructed plume trail (Figure 5.36) shows that the age near Mahanadi Basin is 112 Ma ± 1 Ma. The

underplating is reported beneath the ridge (Choudhuri et al., 2013) but this is beneath ridge itself, not

much underplating is reported in the rifted zone, except in place where the hotspot is directly passing

through the rifted zone in the Mahanadi Basin from onshore to offshore (Bird, 2009). Therefore, the

underplating and volcanism are the post breakup phenomena as the timing arguments suggest. It is to


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be noted that the hotspot trail, which is a physical feature represented by 85°E Ridge and the breakup

and hypothetical migration track of the Elan Bank as shown in successive reconstruction (Figure 5.43)

clearly intersect with each other at a very high angle. It is highly unlikely that the two simultaneous and

synchronous tectonic events i.e., breakup of the Elan Bank and formation of 85°Ridge caused by the

same Kerguelen mantle plume would create two different trends.

Figure 6.7: Conceptual model of the Elan Bank microcontinent formation based on plume refocusing hypothesis. It shows that the initial breakup in the east of the Elan Bank is abandoned in favour of the Krishna-Godavari basin due to arrival of the Kerguelen plume. Ridge jump likely due to Kerguelen hotspot interference and the spreading center targeted the zone between India and Elan Bank, thus transferring the Elan Bank to the Antarctic Plate. Old spreading center is become inactive and the Elan Bank remained as a micro-continent. See text for more explanation.


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The crustal architecture of the Elan Bank suggests that it is a hyper-extended margin with minimum syn-

rift volcanism. The maximum volcanism in Elan Bank is mostly on the south-western margin, where the

thicker section is visibly in a post rift position. The radiometric dating of ODP Site-1137 suggests 107.7 ±

0.5 Ma ages of the volcanics, when the peak volcanism took place. If the sampling depth of ODP-1137,

which is located within proximal margin of Elan Bank is compared with reflection seismic (Figure 3.24), it

indicates that the sample depth is clearly above the interpreted synrift section. Therefore, it can be

concluded that the major plume activity in the Elan Bank has post-breakup timing. The absence of syn-

breakup volcanism is also evident in other microcontinent examples including Jan Mayen (Kodaira et al.,

1998); Seychelles (Collier et al., 2008) and the Lomonsov Ridge in the Arctic Ocean (Minakov et al.,

2013).

No data has been interpreted on Antarctica in current study. The data are available only from

theliterature and published seismic images from Geoscience Australia. The width of the extended crust

in Antarctica and the geometry of the ocean-continent transition boundary, which is more than 250 km

as shown in tectonic element map of Antarctica (O’Brien and Stagg, 2007; Figure 2.35) possibly indicate

that the Antarctic margin is also a hyper-extended continental margin. Seismic images through the East

Antarctica, particularly the offshore Enderby Basin also show segmented margin architecture with no

clear volcanic evidence (Stagg et al, 2005; Figures 2.35, 2.36). Enderby Basin is appears to be underlain

by the oceanic crust with normal thickness and relatively clear Moho Image (Stagg et al, 2005; Figure

6.8). The M9 magnetic anomaly identification in this region also suggests that it is an early Cretaceous

oceanic crust (Ramanna et al., 2001; O’Brien and Stagg, 2007). The lack of clear information about

thickened crust due to oceanic accretion (Stagg et al., 2005; O’Brien and Stagg, 2007) and the

considerable width of extended continental crust is probably a good argument to rule out the plume

driven breakup in Antarctica.


270

Figure 6.8: The seismic interpretation shows Moho imaging in offshore Antarctica (Stagg et al., 2005). Seismic detail from line GA-228/07 shows the interpreted continent–ocean boundary zone. Note the step up from continental/transitional crust on the left to oceanic crust on the right, and the abrupt southwards termination of reflection Moho beneath the inboard edge of oceanic crust at 10 s TWT. Note also the horizontal partitioning of the oceanic crust into a thin upper layer of short, seaward-dipping flows; semitransparent upper-middle crust; lower, highly-reflective crust; and high-amplitude, continuous reflection Moho.

The plume-related renewed rifting hypothesis (Müller et. al, 2001) is also debated when the lithospheric

strength, thermodynamics and kinematics of the rifting in concerned. When a plume impinges upon the

continental lithosphere, it produces a large swell. A numerical model shows that a regional dynamic

uplift can be observed within a diameter of 2000 km from the impingement location, having a

magnitude of 1000 m (Griffiths and Campbell, 1991; Figure 6.9). The diameter of the plume head (D) is a

temperature dependent phenomenon. The temperature difference (ΔT) between ambient mantle and

the plume itself determines its buoyancy flux (Q), which drives the plume ascent. The effect of kinematic

viscosity (v) of the lower mantle also influences the height of the rise (Z) of the plume. The relation is as

follows;

D = Q 1/5 (v/g α ΔT) 1/5K2/5 Z3/5, (6. 1)

where g is acceleration due to gravity, α is coefficient of thermal expansion of mantle and K is the

thermal conductivity (Griffiths and Campbell, 1990).


271

According to this model, if the ΔT is sufficiently high, it produces enough tension to rupture the

dynamically uplifted lithosphere (Griffiths and Campbell, 1990). Other models suggest that even if a

‘hot’ plume with a diameter of 2500 km reaches the lithosphere, it will be unable to generate a

sufficient horizontal deviatoric stress to initiate a lithospheric rupture, unless the uppermost mantle is

already anomalously hot (Hill, 1991; Hill et al., 1992). However, the uppermost mantle can be dilated by

a mantle plume if a rift zone undergoing extension is already there (White and McKenzie, 1989).

The primary assumption of plume refocusing model is that renewed rifting as a result of plume

refocusing needs considerable amount of heat transfer through conduction. Now, several numerical

modeling results (Lavier and Manatschal, 2006; Simon et al., 2009; Huismans and Beaumont, 2011)

successfully prove that the conductive heat transfer is not sufficient enough to generate adequate

deviatoric stress during the rifting.

Figure 6.9: Uplift above a plume head, as predicted by Griffiths and Campbell (1991), compared to the uplift observed at the center of the Emeishan flood basalt by He et al. (2003). Predicted profiles are given for maximum uplift (t = 0), when the top of the plume is at a depth of ~250 km, and 2 Ma later (t = 2 Ma), when flattening of the head is essentially complete. The uplift for the Emeishan flood basalt province is the minimum average value for the inner and intermediate zones, as determined from the depth of erosion of the underlying carbonate rocks.


272

The alternative possible hypothesis for the Elan Bank development is the competing rift zone

hypothesis. As described already in the previous chapter, the Elan Bank was separated as a result of a

competition between the Krishna-Godavari and Cauvery rift zones to host the final continental breakup

(Figures 5.37 to 5.42, Figure 6.10). The rifting in the Cauvery rift zone ceased since its breakup. However,

the Krishna-Godavari rift zone never stopped its propagation. The younger propagation end of the

Krishna-Godavari rift zone eventually got hard-linked through the Coromandal strike-slip system.

However, this linkage is established with the inner failed portion of the Cauvery rift zone instead of

initial linkage between the end-Cauvery zone and inner Krishna-Godavari zone. It can be observed that

the older rift between India and Antarctica ceased when a rift jump occurred and rift between India and

Elan Bank continued westward and eventually kinematically linked through the Coromandal strike-slip

fault (Figure 5.41). This process ended by Aptian (122 Ma). The strike-slip faulting between India and

Elan Bank represents the final breakup. Here the ridge jump could occur due to spreading asymmetry,

which caused irregular ridge propagation. The active spreading center has been shifted to a location

between India and Elan Bank, thus, transferring the Elan Bank to Antarctic Plate as a result of both the

break-up localization and asymmetric ridge propagation. Finally, old spreading center became inactive

and Elan Bank formed a microcontinent.

Several observations support this hypothesis as a favorable alternative for the microcontinent

development in Elan Bank case. These include;

1. The hyper-extended margin architecture of East India, Elan Bank and possibly Antarctica;

2. The lack of serious syn-breakup volcanism;

3. The breakup timing arguments, which are well constrained,

4. The hotspot volcanism timing, which is well constrained through high precision radiometric

dating as having a post-breakup timing; and

5. The interrelation and geometry of hotspot trail and Elan Bank separation vectors.

One of the advantages of this hypothesis is that the kinematics of the rifting and breakup mechanism is

quite universal concerning all passive margins around the world. The strongest evidence comes from

Iberia-Newfoundland system (Reston, 1996; Manatschal and Bernoulii, 1999; Manatschal, 2004;

Manatschal et al., 2007; Tucholke and Sibuet, 2007; Péron-Pinvidic and Manatschal, 2009;

2010;Whitmarsh and Manatschal, 2012; Sutra et al., 2013). The kinematics of rifting is not only based on

seismic examples around the world but also on the outcrop examples in the Swiss Alps also the rifting


273

model (Manatschal, 2004). It has to be noted that rifting mechanism in Iberia-Newfoundland represents

are slow extensional system. However, it has been shown by Nemčok et al., (2012a) studying the Gabon

and East India that this mechanism is also applicable for intermediate and fast extension scenarios.

Several numerical modeling studies (Huismans and Beaumont, 2005; 2007; Lavier and Manatscahl, 2006;

Simon et al., 2009) advocate this model as a preferable one. The reasons for the preferences were

discussed in this study quite extensively in previous chapters.

Figure 6.10: Competing rift zone hypothesis model for the Elan Bank microcontinent formation. The Elan Bank was separated as a result of a competition between Krishna-Godavari and Cauvery rift zones. Ridge-jump can be explained as controlled by the spreading asymmetry, which caused irregular ridge propagation. The active spreading center has been shifted to a location between India and Elan Bank, thus transferring the Elan Bank to Antarctic Plate as a result of both the break-up localization and asymmetric ridge propagation. Old spreading center became inactive and Elan Bank formed a microcontinent. See text for explanation


274

It is to be noted that ridge jumps are more common, when rifting occurs directly above the mantle

plume, as the active rift migrates in an attempt to remain above the plume, which makes the rifting

easiest. It is also observed that most of the ridge jumps are prevalent during the earliest phase of the

sea-floor spreading, immediately following the breakup before the spreading patterns become fully

organized (Smallwood and White, 2002; Rosendahl et al., 2005). During advanced stage of continental

extension, the active rifting becomes the dominant process as the asthenospheric ridge rises and

focuses towards the organized spreading (Huismans and Beaumont, 2005; Manatschal et al., 2007).

Therefore, in either case, ridge jumps are possible with or without plume refocusing. Therefore, the trail

of hotspot in combination with crustal architecture and timing critically rules out the control of

Kerguelen plume on the pre-breakup margin evolution along the future East Indian margin.

These observations are not exclusive for the Elan Bank. The detailed study in Jan Mayen microcontinent

also presents similar observations. The crustal structure of the Jan Mayen microcontinent is

characterized by a sedimentary basin, a thin basaltic layer within the sedimentary section and thinned

continental crust in direction towards the Iceland Plateau (Kodaira et al., 1998; Péron-Pinvidic et al.,

2010 Figure 6.11). It is now commonly accepted that Jan Mayen microcontinent is formed due to

multiple reorganizations of ridge segments along the hype-extended Norwegian rifted margin (Gernigon

et al., 2012) (Figure 6.12). Scott et al., (2005) propose that stepwise northward propagation of the

Kolbeinsey Ridge and simultaneous northward retreat of the Aegir Ridge (Figure 6.13) cause the ridge

jump and subsequent separation of the Jan Mayen microcontinent. The ridge tips were linked by a

fracture zone known as the Jan Mayen fracture zone (Skogseid and Eldholm, 1987) that was occasionally

replaced by the new short-lived fracture zones to the north cutting through the microcontinent. This

was another factor, which contributed to balanced propagation and retreat of the spreading ridges,

causing the segmentation of the intervening oceanic and microcontinent lithospheres (Lundin and Doré,

2002; Scott, et al., 2005; Gernigon et al., 2012 ). The dominant sea-floor spreading along the Kolbeinsey

Ridge was finally achieved during Oligocene-Miocene, leading to complete separation of the

microcontinent (Rey, et al., 2003). This event was coeval with the extinction of the Aegir Ridge.


275

Figure 6.11: Crustal architecture and continental breakup of the Jan Mayen Microcontinent (Péron-Pinvidic et al., 2010). Schematic crustal architecture based on seismic interpretation and analogue models of the Jan Mayen microcontinent shows the extreme thinning of continental crust and rifted sedimentary basins in the north-west.

Roest et al. (2002) suggest that when both the Kolbeinsey and Aegir ridges were active, the

microcontinent underwent a 30°–50° anticlockwise rotation. This was caused by the northeastward

propagation of the Kolbeinsey Ridge and simultaneous reduction in the spreading rate at the

southwestern end of the Aegir Ridge (Lundin and Doré, 2002), which finally became inactive by Miocene

(magnetic anomaly C6). The Kolbeinsey Ridge reached the Jan Mayen Fracture Zone and linked with the

Mohns Ridge during Late Oligocene to Early Miocene (magnetic anomaly C6) (Lundin and Doré, 2002).

At this time, the activity of the Aegir Ridge died out and the microcontinent became part of the Eurasian

Plate. In Jan Mayen, the plume activity has been proved to be post-dating the development of

microcontinent (Lundin and Doré, 2002; Gernigon et al., 2012)


276

Figure 6.12: Kinematic evolution of the Norway Basin and the Jan Mayen microcontinent illustrated by a series of key plate reconstruction stages (Gernigon et al., 2012). Jan Mayen Microcontinent (JMMC); continent–ocean transition (COT); mid-ocean ridge (MOR); Kolbeinsey Ridge (KB); Aegir Ridge (AR). Mohns Ridge (MR). The arrows show ridge propagation direction from south. See explanation in text.


277

Figure 6.13: Schematic sequence of panels, showing the evolution of the North Atlantic ridge system and development of the Jan Mayen microcontinent (Scott et al., 2005). In the first stage (1), spreading between C24 and C18 anomalies had a constant NW-SE direction. During this stage the Aegir ridge (A) was straight and parallel to Mohns and Reykjanes ridges, located to the east of Reykjanes and south of Mohns Ridge. During the second stage (2), a change in spreading direction occurred from previous NW-SE trend to WNW-ESE trend. A stepwise northward propagation of the Kolbeinsey ridge (K) was linked by fracture zones to the northward retreating southern tip of the active Aegir ridge. As a consequence, the Jan Mayen microcontinent became progressively separated as the linking fracture zones ‘jumped’ periodically northward. Finally (3), the propagating tip of the Kolbeinsey ridge had reached the area to the west of the Jan Mayen fracture zone and spreading finally ceased on the Aegir ridge.

Although not extensively as Jan Mayen, the Seychelles also show a similar history. They underwent two

stages of rifting to isolate themselves from Madagascar and India. The Seychelles and India together

became separated from Madagascar sometimes around 85-90 Ma (Reeves and DeWit, 2000; Reeves et

al., 2002). An initial period of strike-slip movement at around 90 Ma caused the northward migration of

the Seychelles-India landmass (Plummer and Belle, 1995). At around 84 Ma, the oceanic crust started to

form in the Mascarene Basin, causing an anti-clockwise rotation of the Seychelles-India land mass

(Reeves and de-Wit, 2000). This continued until 65 Ma when a renewed rifting and subsequent ridge

jump separated the Seychelles from India, forming the currently active Carlsberg Ridge (Dyment, 1998;

Reeves and de-Wit, 2000; Collier et al., 2008) (Figure 6.14). The renewed rifting is commonly attributed

to the asymmetric spreading at the Carlsberg Ridge, located to the southwest of India (Müller et al.,


278

2001). This was caused by propagation of the ridge toward the Reunion hotspot (Dyment, 1998). The

main problem of Seychelles so far is that there was no hyper-extended margin architecture proved

either in Seychelles or its conjugate West India.

Figure 6.14: Tectonic evolution of the Seychelles microcontinent (Reeves, 2008, unpublished). Around 84 Ma, the oceanic crust started to form in the Mascarene Basin, causing an anti-clockwise rotation of the Seychelles-India land mass (Reeves and de-Wit, 2000). This continued until 65 Ma when the renewed rifting separated the Seychelles from India, forming the currently active Carlsberg Ridge. Separation of the West India and Mascarenes was completed at around 65 Ma. A new mid-ocean ridge became established between India and Seychelles.

A key finding of this study is the involvement of strike-slip faulting in the breakup localization and

release of microcontinent, which opens up a possibility for future research. In the case of East India, it

has been discussed that the kinematic linkage between propagating rift zones became eventually

established through the Coromandal strike-slip fault. Another example of this process comes from

Madagascar, although, it is strictly not a microcontinent as it is still attached to the African continent

along the Davie Fracture Zone (Figure 6.15). Thanks to that one can see which mechanism almost

succeeded in this release. The breakup first took place between the orthogonal Somalia-Majunga (north


279

Madagascar) conjugates and Zambezi (South Mozambique)-Antarctica conjugates. The final breakup was

localizing along the Davie strike-slip system as it formed a kinematic linkage between the two

orthogonal rift zones. However, a spreading reorganization in the Indian Ocean in Late Cretaceous

(Aptian) caused the migration of active ridge towards the east of Madagascar, causing the extinction of

existing spreading centers located between Africa-Antarctica and Madagascar. This event took place just

before the final release of microcontinent.

The change in thermal history and strain evolution during the release of a microcontinent can be

another future research subject. The strain evolution during rifting of conjugate margins has been

studied by previous workers (Davies and Kusznir, 2004; Lavier and Manatschal, 2006; Sutra et al., 2013).

However, the strain evolution is expected to be different if a microcontinent is released according to the

model presented in this study. The change in thermodynamics of the rifting along the associated passive

margin is also expected to be quite anomalous compared to passive margins with no microcontinent

associated. This can be a good problem for future research.

This study potentially describes the role of the microcontinents to comprehend the complex

geodynamic and tectonic histories in passive margin development. The Elan Bank case study presented

here has several advantages in comparison to other examples around the world because a high

resolution, dense grid of reflection seismic imaging provides a better control over the crustal

architecture interpretation and reconstruction of the rifting history of India-Elan Bank conjugate

margins. The presence of well data also provides better control over the tectonic timing arguments.

However, the study lacks a control over oceanic domain, where magnetic anomalies are poorly

constrained both due to lack of data and difficulty in their identification discussed earlier.

Combining all the observations, interpretation, modeling and global analogues, it is proposed that Elan

Bank microcontinent developed as a result of the competition between rift zones, breakup localization

along a strike slip system and irregular asymmetric spreading, which caused the ridge jump and eventual

final release of the Elan Bank microcontinent.


280

Figure 6.15: Tectonic map of East Africa (modified from Reeves et al., 1986; CGMW, 1990; Rapolla et al., 1995; Mbede and Dualeh, 1997; Tanzania Petroleum Development Corporation, 2009). Kinematics of rifting indicates that the rifting between Madagascar and Somalia and Mozambique and Antarctica is linked through the Davie fracture zone (DFZ). DFZ has been interpreted as a dextral strike-slip zone. The deformation of which is localized in the northern Mozambique region. Its faults are oriented roughly N-S. The Madagascar breakup was never completed due to the ridge reorganization in the eastern part of Madagascar. Note that Madagascar is still attached to the African continent. See text for explanation.

281

Chapter 7

Conclusions

7.1 Conclusions

The evolution of conjugate passive margins from initial stretching phase to sea floor spreading has been

evaluated in detail in this study. The India-Elan Bank-Antarctica case study addresses the geodynamic

interpretation of ridge jump and the Elan Bank microcontinent release mechanism. The key findings can

be broadly categorized into three principal domains:

1. The reconstruction and restoration of conjugate passive margins

2. The crustal architecture of passive margins and continental breakup

3. Microcontinents release mechanism

The reconstruction and restoration of conjugate passive margins

1. The rotation pole based rigid plate geometric reconstructions can be ambiguous while

reconstructing the hyper-extended passive margin. In the case of East India and Elan Bank this

has been proved to occur. The kinematic restoration of hyper-extended margins and

reconstruction of margins will provide a better model for plate reconstruction.

2. The global plate reconstructions based on rotation poles and magnetic anomaly patterns are

suitable only for mega-regional scale but can be misleading on regional scale. To create a

reasonable reconstruction on regional scale, a synthetic reconstruction method followed in the

study provides improved and proper understanding.

3. The first deep marine sediments, which envelops by lapping onto the margin as a result of

thermal subsidence, help to define the upper limit of the continental breakup timing.

Chapter-7 Conclusions

282

The crustal architecture of passive margins and continental breakup

1. The conjugate East India and Elan Bank continental margins appear to be hyper-extended

margins. The hyper-extended margin of India indicates stretching, thinning, exhumation and

spreading domains documenting both coupled and decoupled deformation domains.

2. The results also indicate that the margin evolution along the majority of the east coast of India

has a similarity to Iberia-Newfoundland margin (e.g. Whitmarsh et al. 2001; Manatschal 2004;

Sibuet et al. 2007), West African margin (e.g. Rosendahl et al. 2005) and Exmouth Plateau

margin (Karner & Driscoll 1999; Karner 2008) in terms of the mechanism of continental breakup,

including a lithospheric mantle exhumation.

3. This study also describes margin segmentations along East India. It shows that the East Indian

margin resulted from an interplay of orthogonal, oblique margins either soft-linked through

ramp-flat ramp geometry (Krishna Godavari and Mahanadi orthogonal rift zones) or hard-linked

though strike-slip faulting (Coromandal strike-slip system linking the Krishna-Godavari and

Cauvery orthogonal rift zones). It has also shown that the margin segmentation is a function of

crustal types and tectonic inheritance.

4. The East Indian margin architecture shows an exhumation domain represented by proto-oceanic

crust, which exists between stretched and thinned continental crust and normal oceanic crust

formed by organized sea-floor spreading. It is also shown that the width of the proto-oceanic

domain varies depending upon the nature of extension. It is narrower in front of the

Coromandal transtensional transform margin segment compared to the Krishna-Godavari

extensional margin segment.

5. The crustal architecture of the Elan Bank shows mostly distal margin and exhumation domain

architecture except the southern margin, where the proximal to distal margin architecture is

demarcated. The physical shape of the Elan Bank has been reinterpreted based on hyper-

extended margin architecture of the microcontinent.

6. The margin architecture of Elan Bank is not fairly constrained due to limited seismic data

availability. However, it is proved that the north-western margin represents an orthogonal rift

margin and the southern margin is a strike-slip dominated margin. While the south-western

margin represents a more obliquely rifted margin.

7. The profile marriage technique used for reconstruct the conjugate margin architecture indicates

asymmetric rifting between India and Elan Bank.


283

8. The Upper-plate lower plate geometry of the study area consists of the East India margin

representing an upper plate margin and the Elan Bank representing the lower plate margin.

9. Two continental breakups in the Cauvery-Krishna Godavari regions is another critical

observation in this study. The first continental breakup occurred between India-Elan Bank and

Antarctica along the eastern portion of the Cauvery rift zone around Valanginian (132 Ma). The

second one between the Elan Bank and India along the Krishna-Godavari rift zone became

finalized in Aptian (122 Ma).

10. The two continental breakups are associated with two isostatic uplifts. This claim can be

concluded from the sediment response to dynamic topography change due to uplift and the

missing sections in stratigraphic column of the drilled wells in the uplifted regions.

Microcontinent release mechanism

1. The Elan Bank microcontinent formation is a result of competing rift zone propagation and

breakup localization along crustal weak zones formed during rifting as a result of

hyperextension. The breakup localization initiated along orthogonal rifted segments and finally

culminated along the strike-slip fault zone.

2. The initial breakup between India and Elan Bank was controlled by continued rift propagation of

western end of the Krishna-Godavari rift zone followed by a rift jump in Early Cretaceous in the

northern Krishna-Godavari rift zone.

3. The final breakup between India and Elan Bank was controlled by the dextral movement along

the Coromandal strike-slip fault. This transfer was linked with normal fault systems of both rift

zones via its northern and southern horse-tail structures.

4. The final strain localization occurred as the Coromandal strike slip was linking the failed western

portion of the Cauvery rift zone with the western propagating end of the Krishna-Godavari rift

zone that underwent a continental breakup.

5. The section of the Krishna-Godavari rift zone that did not stop competing with Cauvery rift zone

for the location of the final breakup location was able to capture the break-up to the north of

the future Elan Bank micro-continent causing a ridge jump.

6. The Coromandal transfer fault zone was subsequently active for the entire time required for the

Elan Bank microcontinent and the sea-floor spreading center to the north of it to clear the

contact with the East Indian margin laterally.


284

7. The spreading center jump occurred as a result of the breakup localization and asymmetric ridge

propagation.

8. The plume refocusing model as an alternative explanation for the Elan Bank release failed to

explain all data. This model neither explains the kinematics of rifting and formation of hyper-

extended passive margin nor the age constraints for involved tectonic events. The plume related

volcanism, for example, has a post-breakup timing. It occurred in the onshore north-east India

(Rajmahal and Syllet) around 118 Ma and in the Elan Bank around 108 Ma, i.e., 5 Ma and 15 Ma

after the breakup (123 Ma).

The proposed competing rift zone hypothesis as a microcontinent release model for the Elan Bank

represents the main result of this study. This microcontinent development model can be applicable

to other microcontinents and passive margins associated with continental ribbons. Suggested

analogues include Sri-Lanka and Madagascar.

7.2 Future Works

Finally, it is to be mentioned that in this thesis, there are few topics have not been fully explored either

due to time limitations or it was beyond the scope of the present study. Some of the listed topics below

can be taken up as potential future research projects.

1. The improved reconstruction of hyper-extended architecture of India-Elan Bank and Antarctica

with additional datasets, potentially addressing those gaps mentioned in this study. It requires

the recalculation of rotation poles to incorporate Elan Bank and additional pieces of

microcontinents indicated inside the Kerguelen Plateau. It may also aim to rediscover the

composite microcontinent in the Indian Ocean.

2. It is necessary to establish a better control of Elan Bank crustal structure integrating seismic

interpretation, forward gravity modeling and velocity analysis. This will require improved

processing of existing data or even may be additional seismic data acquisition in the Elan Bank.

The additional data in the northern margin of the Elan Bank will be needed to interpret the

hyper-extended crustal architecture and to perform seismic paring with conjugate East India

seismic profiles to understand the evolution of microcontinents in a better way.

3. The role of breakup localization and strike-slip faulting in microcontinent releasing mechanism,

presented in this study should be further investigated. Additional examples from different


285

microcontinents associated with sheared margins around the world can be a key to understand

this process better. The results of the future study can be important to understand the strain

evolution and thermal evolution of passive margins associated with microcontinents.


286

References

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References

Acton, G. D., 1999. Apparent polar wander of India since the Cretaceous with implications for regional tectonics and true polar wander. In: Radhakrishna, T. et al. (Eds.), The Indian subcontinent and Gondwana: a palaeomagnetic and rock magnetic perspective. Geological Society of India Memoir, 44, pp. 129-175.

Aitchison, J. C., Ali, J. R. and Davis, A. M., 2007. When and where did India and Asia collide? Journal of Geophysical Research, 112, B05423, doi:10.1029/2006JB004706.

Ajayakumar, P., Kurian, P. J., Rajendran, S., Radhakrishna, M., Nambiar, C.G., Mahadevan, T.M., Heterogeneity in crustal structure across the Southern Granulite Terrain (SGT): Inferences from an analysis of gravity and magnetic fields in the Periyar plateau and adjoining areas, Gondwana Research, 10, 18–28, doi:10.1016/j.gr.2005.11.011

Alvey, A., Gaina, C., Kusznir, N.J., and Torsvik, T.H., 2008. Integrated crustal thickness mapping and plate reconstructions for the high Arctic, Earth and Planetary Science Letters, 274, 3-4, 310-321

Anders, M.H. and Schlische, R.W., 1994. Overlapping faults, intrabasin highs, and the growth of normal faults. Journal of Geology, 102, 2, 165-179

Baksi, A. K., 1995. Petrogenesis and timing of volcanism in the Rajmahal flood basalt province, northeastern India. Chem. Geol., 121: 73-90.

Bastia, R., 2007. Geologic settings and petroleum systems of India’s East coast basins: concepts and application. Technology Publications, Dehradun, India.

Bastia, R., Radhakrishna, M., Das, S., Kale, A.S., and Catuneanu, O., 2010. Delineation of the 85_E ridge and its structure in the Mahanadi Offshore Basin, Eastern Continental Margin of India (ECMI), from seismic reflection imaging, Marine and Petroleum Geology, 27, 1841-1848, doi:10.1016/j.marpetgeo.2010.08.003

Bénard, F., Callot, J.P., Vially, R., Schmitz, J., Roest, W., Patriat, M., Loubrieu, B., and The Extra Plac Team, 2010. The Kerguelen plateau: Records from a long-living/composite microcontinent, Marine and Petroleum Geology, 27, 3, 633-649

Bhattacharya, S., 1996. Eastern Ghats granulites terrain of India: an overview. Journal of Southeast Asian Earth Sciences, 14, 3-4, 165-174, doi:10.1016/S0743-9547(96)00055-4

Bhattacharya, S., 2001. Archaean High-T Decompression and Partial Melting in the Eastern Ghats Belt, India: Correlation with the Antarctic Napier Complex. Gondwana Research, 4, 4, 575-576, doi: 10.1016/S0743-9547(96)00055-4

Bird, D. E., 2001, Shear margins: continent – ocean transform and fracture zone boundaries: The Leading Edge, vol. 20, no. 2, pp. 150-159.

Bird, D. E., 2009. Offshore East India, Two Dimensional Gravity and Magnetic Models. Bird Geophysical, Internal Report for Reliance Industries Ltd., unpublished

Bird, D. E., Burke, K., Hall, S. A., and Casey, J. F., 2011, Tectonic evolution of the Gulf of Mexico basin: in, Buster, N. A., and Holmes, C. W. (editors), Tunnell, J. W., Felder, D. L., and Earle, S. A. (series editors), The Gulf of Mexico origin, waters, and biota. Volume 3, geology: Texas A&M University Press, pp. 3-16.

References

288

Biswal, T. K. and Sinha, S., 2004. Fold-thrust-belt structure of the Proterozoic Eastern Ghats Mobile Belt: a proposed correlation between India and Antarctica in Gondwana. Gondwana Research, 7, 1, 43-56.

Biswas, S.K., 2003. Regional tectonic framework of the Pranhita–Godavari basin, India, Journal of Asian Earth Sciences. 21, 1– 9

Black, L. P., Harley, S. L., Sun, S. S. and McCulloch, M. T., 1987. The Rayner Complex of East Antarctica: complex isotopic systematics within a Proterozoic mobile belt. Journal of Metamorphic Geology, 5: 1-26

Blakely, R. J., 1995, Potential Theory in Gravity and Magnetic Applications: Cambridge Univ. Press, 441

Borissova, I., Coffin, M. F., Charvis, P. and Operto, S., 2003. Structure and development of a micro-continent: Elan Bank in the southern Indian Ocean. Geochemistry,Geophysics,Geosystems. 4(9), 1071, doi:10.1029/2003GC000535

Borissova, I., Moore, A., Sayers, J., Parums, R., Coffin, M.F., and Symonds P.A., 2002. Geological framework of the Kerguelen Plateau and adjacent ocean basins, Geoscience Australia Record 2002/05

Brozena, J.M., and White, R.S., 1990. Ridge jumps and propagations in South Atlantic Ocean, Nature, 348, 6297, 149-152

Brun, J.P., and Beslier, M.O., 1996. Mantle exhumation at passive margins. Earth and Planetary Science Letters, 142, 1-2, 161-173 doi: 10.1016/0012-821X (96)00080-5

Buck, R.W., 1986. Small-scale convection induced by passive rifting: the cause for uplift of rift shoulders Earth and Planetary Science Letters, 77, 3–4, 362-372

Bullard, E.C., Everett, J.E. and Smith, A.G., 1965. The fit of the continents around the Atlantic, Royal Society of London Transactions, series A., 258, 41-51

Catuneanu, O., et al., 2009. Towards the standardization of sequence stratigraphy, Earth-Science Reviews, 92, 1–2, 1–33

Catuneanu, O., Wopfner, H., Eriksson, P.G., Cairncross B., Rubidge, B.S., Smith, R.M.H., Hancox, P.J.,2005. The Karoo basins of south-central Africa, Journal of African Earth Sciences, 43, 211–253

Cawood, P.A., Kröner, A. and Pisarevsky, S., 2006. Precambrian plate tectonics: criteria and evidence, GSA Today, 16, 4–11.

Chakraborty, C., Mandal, N., and Ghosh, S.K., 2003. Kinematics of the Gondwana basins of peninsular India, Tectonophysics, 377, 299– 324

Chamberlin, R. M., 2000. Partitioning of Dextral Slip in an Incipient Transverse Shear Zone of Neogene Age, Northwestern Albuquerque Basin, Rio Grande Rift, New Mexico: in Cole J.C. (ed.), U.S. Geological Survey Middle Rio Grande Basin Study— Proceedings of the Forth Annual Workshop, Albuquerque, New Mexico,U.S. Geological Survey Open-File Report 00-488, p. 51.

Charvis P., Operto S., Lesne O., and Royer J.-Y., 1997. Velocity structure of the Kerguelen volcanic province from wide-angle seismic data: petrological implications, Eos, Trans. Amer. Geophys. Union, 78: 711

Charvis, P., and Operto, S., 1999. Structure of the Cretaceous Kerguelen volcanic province (southern Indian Ocean) from wide-angle seismic data, Journal of Geodynamics, 28, 51-71

References

289

Charvis, P., Recq, M., Operto, S., and Brefort, D., 1995. Deep structure of the northern Kerguelen Plateau and hot spot related activity. Geophysical Journal International, 122, 899-924

Chaudhuri, A., Rao, M. V., Dobriyal, J. P., Saha G. C., Chidambaram, L., Mehta, A. K., Ramana, L. V., and Murthy, K. S., 2010. Prospectivity of Cauvery Basin in deep syn-rift sequences, SE India., Search and Discovery Article #10232

Chetty, T.R.K., Vijay, P., Narayana, B.L., and Giridhar G.V., 2003. Structure of the Nagavali Shear Zone, Eastern Ghats Mobile Belt, India: Correlation in the East Gondwana Reconstruction, Gondwana Research, 6, 2, 215-229.

Choudhuri, M., Guha, D., Dutta, A., Sinha, S., and Sinha, N., 2010. Spatio temporal variations and kinematics of shale mobility in the Krishna-Godavari basin, India, in: Wood, L., (ed), Shale tectonics, AAPG Memoir, 93, 91 – 109. doi: 10.1306/13231310M933420

Choudhuri, M., Nemčok, M., Stuart, C., Welker, C., Sinha, S.T. and Bird, D., 2013, 85° E Ridge, India - constraints on its development and architecture, Jour. Geol. Soc. Ind., accepted, JGSI-D-12-00285.

Christensen, N.I. and Mooney, W.D., 1995. Seismic velocity structure and composition of the continental crust. Journal of Geophysical Research, 100, 9761-9788

Clark, S. P., Jr. 1966. Handbook of Physical Constants. Geological Society of America, Memoirs, 97

Clarke, G. L., 1988. Structural constraint on the Proterozoic reworking of Archaean crust in the Rayner Complex, MacRobertson and east Kemp Land, East Antarctica. Precambrian Research, 40/41: 137-156.

Coffin, M.F., Symonds, P., Ramsay, D., Bernardel, G. Gladczenko, T., 1997. A deep seismic transect across the Kerguelen Plateau. EOS. Transaction of the AGU 1997 Fall meeting. 78, 46, suppl. page 711–712

Coffin, M. F., Pringle, M. S., Duncan, R. A., Gladczenko, T. P., Storey, M., Muller, R. D., and Gahagan L. A., 2002. Kerguelen hotspot magma output since 130Ma, Journal of Petrology, 43, 1121-1139

Coffin, M. F., Frey, F. A., Wallace, P. J., et al., 2000. Proc. ODP, Init. Repts.,183: Ocean Drilling Program, Texas A&M University, College Station, TX

Coffin, M. F., Davies, H. L., and Haxby, W. F., 1986. Structure of the Kerguelen Plateau province from Seasat altimetry and seismic reflection data, Nature, 324, 134-136

Coffin, M. F., and Eldholm, O., 1992. Volcanism and continental break-up: a global compilation of large igneous provinces, In: Storey, B.C., Alabaster, T., and Pankhurst, R.J. (eds) Magmatism and the Causes of Continental Breakup, Geological Society, London, Special Publications, 68, 17–30. doi:10.1144/GSL.SP.1992.068.01.02.

Collier, J.S., Sansom, V., Ishizuka, O., Taylor, R.N., Minshull, T.A., and Whitmarsh, R.B., 2008. Age of Seychelles–India break-up, Earth and Planetary Science Letters, 272, 264–277, doi:10.1016/j.epsl.2008.04.045

Collins, A. S., and Pisarevsky, S. A., 2005. Amalgamating eastern Gondwana: The evolution of the Circum-Indian Orogens, Earth-Science Reviews, 71, 3–4, 229–270, doi:10.1016/j.earscirev.2005.02.004

Colwell, J.B., Stagg, H. M. J., Direen, N. G., Bernardel, G., and Borissova, I. 2006, The Structure of the Continental Margin off Wilkes Land and Terre Adélie Coast, East Antarctica, in: Fütterer, D.K., Dr. Damaske, D., Kleinschmidt, G. Miller, H., and Tessensohn, F., (eds), Antarctica: Contributions to Global Earth Sciences, 327-340, ISBN: 978-3-540-30673-3

References

290

Commission for the Geological Map of the World, 1990. International Geological Map of Africa 6: 6 of 6.

Condie, C., 2005. Earth as an evolving planetary system. Elsevier Academic Press, Amsterdam, 17-25, 265-313

Conrad, C.P., Gurnis, M., 2003. Seismic topography, surface uplift, and the breakup of Gondwanaland: integrating mantle convection backward in time, Geochem. Geophys. Geosyst, 4, 1031.

Cooper, A., Stagg, H., and Geist, E., 1991, Seismic stratigraphy and structure of Prydz Bay, Antarctica: Implications from Leg 119 drilling, in: Barron, J., Anderson, J., Baldauf, J.G., and Larsen, B., (eds), Proceedings of the Ocean Drilling Program, Scientific results, 119: College Station, Texas, 5–26.

Crosby, A., White, N., Edwards, G., Shillington, D.J., 2008. Evolution of the Newfoundland–Iberia conjugate rifted margins, Earth and Planetary Science Letters, 273, 214–226, doi: 10.1016/j.epsl.2008.06.039

Curray, J.R. and Munasinghe, T., 1991. Origin of the Rajmahal Traps and the 85°E Ridge preliminary reconstructions of the trace of the Crozet hotspot Geology, 19,1237-1240.

David, J. S., Srihari, J., Basha, S. M. J., Balachandrudu, V., Narahari, S. T. and Augustine, P. F., 1998. Craton-mobile belt relation along the western margin of EGMB in: A. P. Eastern India: an evidence for thrusted contact. Sem. Abst. Vol. Intl. Sem. On Precambrian Crust in Eastern and Central India, BBSR, pp. 216-218.

Davis, M., and Kusznir, N.J., 2004, Depth-dependent lithospheric stretching at rifted continental margins: in Karner, G.D., ed., Proceedings of National Science Foundation Rifted Margins Theoretical Institute: New York, Columbia University Press, pp. 92-136

Dehler, S. A., and Welford, J. K., 2013. Variations in rifting style and structure of the Scotian margin, Atlantic Canada, from 3D gravity inversion, Geological Society, London, Special Publications, 369, 289-300, doi: 10.1144/SP369.11

Delvaux, D., 2001. Tectonic and palaeostress evolution of the Tanganyika-Rukwa-Malawi rift segment, East African Rift System. In: P.A. Ziegler, W. Cavazza and A.H.F. Robertson and S. Crasquin-Soleau (eds.), Peri-Tethys Memoir 6: PeriTethyan Rift/Wrench Basins and Passive Margins. Mém. Mus. Natn. Hist.nat., Paris, 186 : 545-567.

Desa, M., Ramana, M.V., and Ramprasad, T., 2006. Seafloor spreading magnetic anomalies south off Sri Lanka, Marine Geology, 229, 3–4, 227–240, http://dx.doi.org/10.1016/j.margeo.2006.03.006

Driscoll, N.W. and Karner, G.D., 1998. Lower crustal extension across the northern Carnarvon Basin, Australia; evidence for an eastward dipping detachment. Journal of Geophysical Research, 103(B3): 4975-4991.

Drury, S. A., Harris, N. B. W., Holt, R. W., Reeves-Smith, G. J., and Wightman R. T., 1984. Precambrian tectonics and crustal evolution in South India. The Journal of Geology, 92, 1, 3–20

Dunbar, J. A., and D. S. Sawyer 1987. Implications of continental crust extension for plate reconstruction: An example from the Gulf of Mexico, Tectonics, 6(6), 739–755, doi:10.1029/TC006i006p00739.

Duncan, R.A., 2002. A time frame for construction of the Kerguelen plateau and Broken Ridge. Journal of Petrology 43 (7), 1109–1119

Dyment, J., 1998, Evolution of the Carlsberg Ridge between 60 and 45 Ma— Ridge propagation, spreading asymmetry, and the Deccan-Reunion hotspot: Journal of Geophysical Research, 103, 24 067–24 084

References

291

Erram, V. C., Rajaram, M. and Anand, S. P., 2005. Comparative magnetic study of the hydrocarbon-bearing Cauvery basin of East coast of India. www.cosis.net/abstracts/IAGA2005/00316/IAGA2005-A-00316.pdf

Ewart, A., Marsh, J. S., Milner, S. C., Duncan, A. R., Kamber, B. S., and Armstrong, R. A., 2004. Petrology and Geochemistry of Early Cretaceous Bimodal Continental Flood Volcanism of the NW Etendeka, Namibia, Part 1: Introduction, Mafic Lavas and Re-evaluation of Mantle Source Components, Journal of Petrology, 45, 1, 59-105, doi: 10.1093/petrology/egg083

Ferraccioli, F., Finn, C.A., Jordan, T.A., Bell, R. E., Anderson, L.M., and Damaske, D., 2011. East Antarctic rifting triggers uplift of the Gamburtsev Mountains, Nature, 479, 388–392, doi: 10.1038/nature10566

Fraser, S.I., Fraser, A.J., Lentini, M.R. and Gawthorpe, R.L. (2007) Return to rifts - the next wave: fresh insights into the petroleum geology of global rift basins. Petroleum Geoscience, 13, 99-104.

Frey, F.A., Weis, D., Borisova, A., Yu Xu, G., 2002. Involvement of continental crust in the formation of the Cretaceous Kerguelen Plateau; new perspectives from ODP Leg 120 sites. Journal of Petrology, 43 ,7, 1207–1239

Fuloria, R. C., Pandey, R. N., Bharali, B. R., and Mishra, J. K., 1992. Stratigraphy, structure and tectonics of Mahanadi offshore basin. in: Recent Geoscientific Studies in the Bay of Bengal and the Andaman Sea, Geological Survey of India, Special Publication, 29, 255-265.

Gaina, C., Műller, R.D., Brown, B., and Ishihara, T., 2007. Breakup and early sea floor spreading between India and Antarctica, Geophysical Journal International, 170, 1, 151–169, doi: 10.1111/j.1365-246X.2007.03450.x

Gaina, C., Müller, R. D., Brown, B., and Ishihara, T., 2003. Microcontinent formation around Australia, In: Hillis, R. and Müller, R. D, (eds) The Evolution and Dynamics of the Australian Plate, Special Publication, Geological Society of America, 22, 399– 410

Gaina, C., Műller, R.D., Brown, B., Ishihara, T., 2002. Microcontinent formation around Australia. Geological Society of Australia, Special Publication 22, 399–410

Galloway, W.E., 1989. Genetic stratigraphic sequences in basin analysis, I. Architecture and genesis of flooding-surface bounded depositional units. American Association of Petroleum Geologists, Bulletin 73, 125–142

Gawthorpe, R.L. et al., 2003. Normal fault growth, displacement localisation and the evolution of normal fault populations; the Hammam Faraun fault block, Suez Rift, Egypt. Journal of Structural Geology, 25(6): 883-895

Geoffroy, L., 2005. Volcanic passive margins, Comptes Rendus Geosciences, 337, 16, 1395-1408. doi:10.1016/j.crte.2005.10.006

Geological Survey of India, 1993. Geological Map of India, Scale 1:5000000

Gernigon, L., Gaina, C., Olesen, P., Ball, P.J., Péron-Pinvidic, G., and Yamasaki, T., 2012. The Norway Basin revisited: From continental breakup to spreading ridge extinction, Marine and Petroleum Geology, 35, 1-19, doi:10.1016/j.marpetgeo.2012.02.015

Gernigon, L., Gaina, C., Péron-Pinvidic, G., and Olesen, O., 2010. Spreading evolution of the Norway Basin and implication for the evolution of the Møre rifted margin and its intermediate conjugate system (the Jan Mayen microcontinent), Central & North Atlantic Conjugate Margins Conference, Lisbon, V. 125 – 129, http://metododirecto.pt/CM2010

References

292

Ghatak, A., and Basu, A. R., 2011. Vestiges of the Kerguelen plume in the Sylhet Traps, northeastern India, Earth and Planetary Science Letters, 308, 1–2, 52-64

Gibbons, A.D., Whittaker, J. M., and Müller, R.D., 2013. The breakup of East Gondwana: Assimilating constraints from Cretaceous ocean basins around India into a best-fit tectonic model, Journal of Geophysical Research: Solid Earth, 118, 808–822, doi:10.1002/jgrb.50079, 2013

Gladczenko, T., Coffin, M.F., 2001. Kerguelen Plateau crustal structure and basin formation from seismic and gravity data. Journal of Geophysical Research, B, Solid Earth and Planets 106 (8), 16,583–16,601

Goff, J.A., and Cochran, J.R., 1996. The Bauer scarp ridge jump: a complex tectonic sequence revealed in satellite altimetry, Earth and Planetary Science Letters, 141, 1-4, 21-33

Goodliffe, A.M, and Taylor, B., 2007. The boundary between continental rifting and sea-floor spreading in the Woodlark Basin, Papua New Guinea, in: Karner, G.D, Manatschal, G., and Pinheiro, L.M (eds), Imaging, Mapping and Modeling of Continental Lithospheric Extension and Breakup, Geological Society, London, Special Publications, 282, 217-238, doi:10.1144/SP282.11

Gopalakrishnan, K., 1998. Extensions of Eastern Ghats mobile belt, India; a geological enigma. Proceedings of workshop on the Eastern Ghats mobile belt, Special Publication Series – Geological Survey of India, 44, 22-38.

Gordon, R. G., DeMets, C. and Argus, D. F., 1990. Kinematic constraints on distributed lithospheric deformation in the equatorial Indian Ocean from present plate motion between the Australian and Indian plates. Tectonics, 9(3): 409-422.

Griffiths, R.W., and Campbell, I.H., 1990. Stirring and structure in mantle starting plumes. Earth and Planetary Science Letters, 99, 66-78

Grikurov G. E., and Mikhalskii, E. V., 2002. Tectonic structure and evolution of East Antarctica in the light of knowledge about supercontinents, Russian Journal of Earth Sciences, 4, 4, 247–257,

Gurnis, M., Mitrovica, J.X., Ritsema, J., Van Heijst, H., 2000. Constraining mantle density structure using geological evidence of surface uplift rates: the case of the African Superplume, Geochem. Geophys. Geosyst, 1, doi:10.1029/ 1999GC000035

Halpin, J. A., Gerakiteys, C. L. , Clarke, G. L. , Belousova, E. A. and Griffin, W. L. , 2005. In-situ U–Pb geochronology and Hf isotope analyses of the Rayner Complex, east Antarctica, Contributions to Mineralogy and Petrology, 148, 6 , 689-706, doi: 10.1007/s00410-004-0627-6

Harley, S.I., and Kelly, N.M., 2007. Ancient Antarctica: the Archean of the east Antarctic shield, in: Van Kranendonk, M.J., Smithies, R.H., and Benne, V.C (eds) , Earth Oldest Rock, Developments in Precambrian Geology (Vol. 15), 149-186, doi: 10.1016/S0166-2635(07)15032-5

Harley, S.L. & Hensen, B.J., 1990. Archaean and Proterozoic high-grade terrenes of East Antarctica (40–80 E); a case study of diversity in granulite facies metamorphism. in: Ashworth, J.R. & Brown, M.C. (Eds), High-Temperature Metamorphism and Crustal Anatexis, Unwin Hyman, London, 320–370,

Harley, S.L., and Motoyoshi, Y., 2000. Al zoning in orthopyroxene in a sapphirine quartzite: evidence for (greater than) 1120° C metamorphism in the Napier Complex, Antarctica, and implications for the entropy of sapphirine. Contributions to Mineralogy and Petrology, 138, 293-307.

References

293

Harrowfield, M., Holdgate, G. R., Wilson, C. J.L., and McLoughlin, S., 2005. Tectonic significance of the Lambert graben, East Antarctica: Reconstructing the Gondwanan rift, Geology, 33,197-200, doi: 10.1130/G21081.1

Hassler, D. R., and Shimizu, N., 1998. Osmium isotopic evidence for ancient subcontinental lithosphere mantle beneath the Kerguelen Islands, Southern Indian Ocean, Science 280: 416-421

Hill, I.R., 1991. Starting plumes and continental break-up, Earth and Planetary Science Letters, 104, 398-416

Hill, R. I., Campbell, I. H., Davies, G. F., and Griffiths, R.W., 1992. Mantle plumes and continental tectonics, Science, 256, 186–193

Hopper, J. R., Funck, T., and Tucholke B. E., 2007. Structure of the Flemish Cap margin, Newfoundland: insights into mantle and crustal processes during continental breakup, in: Karner, G.D, Manatschal, G., and Pinheiro, L.M (eds), Imaging, Mapping and Modeling of Continental Lithospheric Extension and Breakup, Geological Society, London, Special Publications, 282, 47-61, doi:10.1144/SP282.3

Huismans, R.S. and Beaumont, C., 2011. Depth-dependent extension, two-stage breakup and cratonic underplating at rifted margins, Nature, 473,. doi:10.1038/nature09988

Huismans, R.S., and Beaumont C., 2007. Roles of Lithospheric strain softening and heterogeneity determining the geometry of rifts and continental margins, In: Karner, G.D., Manatschal, G. and Pinherio, L.M., (eds), Imaging, Mapping and Modeling Continental Lithospheric Extension and Break Up, Geological Soceity of London, Special Publication, 282, 111-138. doi: 10.1144/sp282.6F

Huismans, R.S., and Beaumont, C., 2005. Effect of Lithospheric Stratification on Extensional Styles and Rift Basin Geometry, 25th Annual Bob F. Perkins Research Conference, Petroleum Systems of Divergent Continental Margin Basins, Houston, pp. 12-55

Ingle, S., Weis, D., and Frey, F.A., 2002. Indian continental crust recovered from Elan Bank, Kerguelen plateau (ODP Leg 183,Site 1137)., Journal of Petrology, 43, 1241-1257

Jurine, D., Jaupart, C., Brandeis, G., and Tackley, P. J., 2005. Penetration of mantle plumes through depleted lithosphere, Journal of Geophysical Research, 110 (B10104)

Karner, G. D., 2000. Rifts of the Campos and Santos basins, southeastern Brazil; distribution and timing. In: M.R. Mello and B.J. Katz (eds), Petroleum systems of South Atlantic margins. AAPG Memoir, 73, 301-315

Karner, G. D., 2008. Depth-dependent Extension and Mantle Exhumation: An Extreme Passive Margin End-member or a New Paradigm? Central Atlantic Conjugate Margin Conference, Halifax, Canada, Extended abstract, 10-16,

Karner, G.D., and Driscoll, N.W., 1999. Style, timing, and distribution of tectonic deformation across the Exmouth Plateau, northwest Australia, determined from stratal architecture and quantitative basin modeling. in: Mac Niocaill, C., and P.D. Ryan (Eds)Continental Tectonics, Special. Publication, Geological Society of London, 164,271-311

Kearey, P., Klepeis, K.A, and Vine, F.J., 2009. Global Tectonics, 3rd Ed, Wiley-Blackwell, 94-97,

Kelmen, P.B. and Holbrook, W.S., 1995. Origin of thick, high-velocity igneous crust along the U.S. East Coast Margin, Journal of Geophysical Research, 100, 10077–10094.

References

294

Kendall, J.M., Stuart, G.W., Ebinger, C.J., Bastow, I.D. and Keir, D., 2005, Magma-assisted rifting in Ethiopia, Nature, 433, 146-148

Kent, R. W., Pringle, M. S., Muller, R. D., Saunders, A. D. and Ghose, N. C., 2002. 40Ar/39Ar geochronology of the Rajmahal basalts, India, and their relationship top the Kerguelen Plateau. Journal of Petrology, 43: 1141-1153.

Kieffer, B.A., Nicholas, T., Weis, D., 2002. A bimodal alkalic shield volcano on Skiff Bank; its place in the evolution of the Kerguelen Plateau. Journal of Petrology, 43 (7), 1259–1286

Kodaira, S., Mjelde, R., Gunnarsson, K., Shiobara, H., and Shimamura, H., 1998. Structure of the Jan Mayen microcontinent and implications for its evolution, Geophysical Journal International, 132(2), 383-400(18)

Kovach, V. P., Salnikova, B., Kotov, A. B., Yakovleva, S. Z. and Rao, A. T., 1997. Pan-African U-Pb zircon age from apatite-magnetite veins of Eastern Ghats Granulite Belt, India. Journal of Geological Society of India, 50: 421-424.

Krause, O., Dobmeier, C., Raith, M. M, and Mezger, K., 2001. Age of emplacement of massif-type anorthosites in the Eastern Ghats Belt, India: constraints from U–Pb zircon dating and structural studies, Precambrian Research, 109, 1–2, 25-38

Krishan, M.S., 1982. Geology of India and Burma. Sixth Ed, CBD Publishers & Distributors, India, 107-109.

Krishna, K.S. and Rao, D.G., 2000. Abandoned Paleocene spreading center in the northeastern Indian Ocean: evidence from magnetic and seismic reflection data, Marine Geology, 162, 2-4,215-224.

Krishna, K.S., 2003. Structure and evolution of the Afanasy Nikitin seamount, buried hills and 85 degrees E Ridge in the northeastern Indian Ocean, Earth Planet. Sciences Letter, 209, 3-4, 379-394

Krishna, K.S., Rao, D.G., Ramana, M.V., Subrahmanyam, V., Sarma, K.V.L.N.S., Pilipenko, A.I., Shcherbakov, V.S., Murthy, I.V.R., 1995. Tectonic model for the evolution of oceanic crust in the northeastern Indian Ocean from the Late Cretaceous to the Early Tertiary, Journal of Geophysical Research, 100, B10, 20011-20024.

Kumar, D., Mamallan, R. and Dwivedy, K. K., 1996. Carbonatite magmatism in northeast India. Journal of Southeastern Asian Earth Sciences, 13: 145-158

Kumar, P., Yuan, X., Ravikumar, M., Kind, R., Li, X., and Chadha, R. K., 2007. The rapid drift of Indian tectonic plate. Nature, 449, 894-897, doi: 10.1038/nature06214

Kusznir, N.J. 2009. South Australia - Antarctica Conjugate Rifted Margins: Mapping Crustal Thickness and Lithosphere Thinning Using Satellite Gravity Inversion-A Report from Geoscience Australia. Web: https://www.ga.gov.au/products/servlet/controller?event=GEOCAT_DETAILSandcatno=68655 (last accessed on 5th August,2010, 16:45)

Lakshminarayana, G., 2002. Evolution in Basin Fill Style during the Mesozoic Gondwana Continental Break-up in the Godavari Triple Junction, SE India, Gondwana Research, 5, 1, 227-244

Läufer, A.L. and Phillips G., 2007. Brittle deformation events in the Lambert glacier region (East Antarctica): insights into the tectonic control on the formation and evolution of the Lambert Graben, Terra Antarctica, 14, 1, 61-68

Lavier, L., and Manatschal, G., 2006. A mechanism to thin the continental lithosphere at magma-poor margins, Nature, 440, 324-328

References

295

Lawver, L. A., L. M. Gahagan, and Coffin M. F., 1992. The development of paleoseaways around Antarctica, in: Kennett, J. P. and Warkne, D. A (eds), The Antarctic Paleoenvironment: A Perspective on Global Change, Part One, Antarctica. Res. Ser., 56. 7–30, AGU, doi: 10.1029/AR056p0007.

Li, Z.X., and Powell, C.M., 2001. An outline of the paleogeographic evolution of the Australasian region since the beginning of Neoproterozoic, Earth Sciences Reviews, 53, 237-277

Lillie, R.J., 1999. Whole Earth Geophysics. An Introductory Textbook for Geologists and Geophysicists. Prentice Hall, Upper Saddle River, 361 pp.

Lisker, F. and Fachmann, S., 2001. Phanerozoic history of the Mahanadi region, India. Journal of Geophysical Research, 106 (B10): 22 027-22 050.

Lisker, F., Gibson, H. Wilson, C.J. and Läufe, A., 2007. Denudation and uplift of the Mawson Escarpment (eastern Lambert Graben, Antarctica) as indicated by apatite fission track data and geomorphological observation, in: Cooper, A.K. and C.R. Raymond et al. (eds) Antarctica: A Keystone in a Changing World – Online Proceedings of the 10th ISAES, USGS Open-File Report 2007-1047, Short Research Paper 105, 6 p.; doi:10.3133/of2007-1047.srp105

Lundin, E., and Doré, A. G., 2002. Mid-Cenozoic post-breakup deformation in the ‘passive’ margins bordering the Norwegian–Greenland Sea, Marine and Petroleum Geology, 19, 1, 79-93

MacNeil, A.J., Jones, B., 2006. Sequence stratigraphy of a Late Devonian ramp situated reef system in the Western Canada Sedimentary Basin: dynamic responses to sea-level change and regressive reef development. Sedimentology 53, 321–359

Mahalik, N. H., 1994. Geology of the contact between the Eastern Ghats Belt and North Orissa Craton, India. Journal of Geological Society of India, 44: 41-52.

Mahoney, J., Jones, W., Frey, F.A., Salters, V., Pyle, D., and Davies, H., 1995. Geochemical characteristics of lavas from Broken Ridge, the Naturaliste Plateau and Southermost Kerguelen Plateau: Early volcanism of the Kerguelen hotspot, Chemical Geology, 120, 315-345.

Manatschal, G. and Bernoulli, D. 1999. Architecture and tectonic evolution of nonvolcanic margins: Present-day Galicia and ancient Adria. Tectonics, 18, doi: 10.1029/1999TC900041

Manatschal, G. and Bernoulli, D., 1998. Rifting and early evolution of ancient ocean basins; the record of the Mesozoic Tethys and the Galicia-Newfoundland margins. Marine Geophysical Research, 20(4), 371-381.

Manatschal, G. and Karner, G. D., 2012. Inter-relationship between tectonic and magmatic processes during hyper-extension and break-up at rifted margins, International Geological Congress, Keynote address, Brisbane, Australia

Manatschal, G., 2004. New models for evolution of magma-poor rifted margins based on a review of data and concepts from West Iberia and the Alps. International Journal of Earth Sciences, 93(3), 432-466.

Manatschal, G., Müntener, O., Lavier, L.L, Minshull, T. A., and Péron-Pinvidic, G., 2007. Observations from the Alpine Tethys and Iberia–Newfoundland margins pertinent to the interpretation of continental breakup, In: Karner, G.D, Manatschal, G., and Pinheiro, L.M (eds), Imaging, Mapping and Modeling of Continental Lithospheric Extension and Breakup, Geological Society, London, Special Publications, 282,291-324, doi:10.1144/SP282.14

References

296

Mbede, E. I., and Dualeh, ,A. 1997. The coastal basins of Somalia, Kenya and Tanzania, in RC Selley ed., African Basins (Sedimentary Basins of the World), Elsevier Science, 211-233

McLoughlin, S., Drinnan, A.N., 1997. Revised stratigraphy of the Permian Bainmedart coal measures, northern Prince Charles Mountains, East Antarctica. Geological Magazine, 134, 335– 353

Meyers, J.B., Rosendahl, B.R., GroscheI-Becker, H., Austin (Jr), J. A., and Rona, P. A., 1996. Deep penetrating MCS imaging of the rift-to-drift transition, offshore Douala and North Gabon basins, West Africa. Marine and Petroleum Geology, 13, 7,791-835

Meyers, J.B., Rosendahl, B.R., Harrison, C.G.A. and Ding, Z.-D., 1998. Deep-imaging seismic and gravity results from the offshore Cameroon volcanic line, and speculation of African hotlines. Tectonophysics, 284(1-2): 31-63.

Mezger, K., and Cosca, M.A., 1999. The thermal history of the Eastern Ghats Belt (India) as revealed by U–Pb and 40Ar/39Ar dating of metamorphic and magmatic minerals: implications for the SWEAT correlation, Precambrian Research, 94, 3–4, 251-271

Minakov, A.N., Podladchikov, Y.Y., Faleide, J.I., Huismans, R.S., 2013. Rifting assisted by shear heating and formation of the Lomonosov Ridge, Earth and Planetary Science Letters, 373, 31–40, http://dx.doi.org/10.1016/j.epsl.2013.04.042

Minshull, T.A., 2009. Geophysical characterisation of the ocean–continent transition at magma-poor rifted margins, Comptes Rendus Geoscience, 341,5, 382–393, http://dx.doi.org/10.1016/j.crte.2008.09.003

Mishra, D.C., Chandra Sekhar, D.V., Venkata Raju, D.C., and Vijaya Kumar, V., 1999. Crustal structure based on gravity magnetic modelling constrained from seismic studies under Lambert rift, Antarctica and Godavari and Mahanadi rifts, India and their interrelationship: Earth and Planetary Science Letters, 172, 287–300

Mitchum Jr., R.M., and Vail, P.R., 1977. Seismic stratigraphy and global changes of sea-level, part 7: stratigraphic interpretation of seismic reflection patterns in depositional sequences,in: Payton, C.E. (Ed.), Seismic Stratigraphy — Applications to Hydrocarbon Exploration. AAPG Memoir-26. American Association of Petroleum Geologists, 135–144.

Mittelstaedt, E., Ito, G. and van Hunen, J. 2011. Repeat ridge jumps associated with plume-ridge interaction, melt transport, and ridge migration, Journal of Geophysical Research, 116, B01102, doi: 10.1029/2010JB007504

Mittelstaedt, E., Ito, G., 2005. Plume–ridge interaction, lithospheric stresses, and the origin of near-ridge volcanic lineaments, Geochem. Geophys. Geosys. 6 (6)

Mittelstaedt, E., Ito, G., and Behn, M.D., 2008. Mid-ocean ridge jumps associated with hotspot magmatism, Earth and Planetary Science Letters, 266, 3-4, 256-270

Mohriak, W. U., and Leroy, S., 2013. Atlantic and Red Sea-Gulf of Aden conjugate margins evolution: insights from the South Atlantic, North Architecture of rifted continental margins and break-up, Geological Society, London, Special Publications, 369, 497-535. doi: 10.1144/SP369.17

Mooney, W.D., and Meissner, R., 1992. Multigenic origin of crustal reflectivity: A review of seismic reflection profiling of continental lower crust and Moho. in: Fountain, D.M., Arculus, R., and Kay, R.W. (Eds.) Continental Lower Crust. Elsevier, Amsterdam, 45-79

Morley, C.K., 1999, Influence of Preexisting Fabrics on Rift Structure, in C.K. Morley ed., Geoscience of Rift Systems— Evolution of East Africa: AAPG Studies in Geology, 44, 151–160.

References

297

Müller R. D., Gaina C., Roest W., and Hansen D. L., 2001. A recipe for microcontinent formation, Geology 29, 203–206

Müller R. D., Roest W. R., and Royer J-Y., 1998. Asymmetric seafloor spreading expresses ridge–plume interactions. Nature, 396, 455–459

Müller, R.D., Roest, W.R., Royer, J.-Y., Gahagan, L.M. and Sclater, J.G., 1997. Digital isochrons of the world's ocean floor. Journal of Geophysical Research, 102, B2, 3211-3214

Műller, R.D., Royer, J.Y., Lawver, L.A., 1993. Revised plate motions relative to the hotspots from combined Atlantic and Indian Ocean hotspot tracks. Geology, 21, 275-278.

Müller, R.D., Sdrolias, M., Gaina, C. and Roest, W.R., 2008, Age spreading rates and spreading asymmetry of the world's ocean crust, Geochemistry, Geophysics, Geosystems, 9, Q04006, doi:10.1029/2007GC001743

Munschy, M., Dyment, J., Boulanger, M.O., Boulanger, D., Tissot, J.D., Schlich, R., Rotstein, Y., and Coffin, M.F., 1992. Breakup and sea floor spreading between the Kerguelen Plateau-Labuan Basin and Broken Ridge-Diamantina Zone. In Wise, S.W., Jr., Schlich, R., et al., Proceedings ODP, Scientific Results, 120, 931-944.

Murty, K. V. S. and Ramakrishna, M., 1980. Structure and tectonics of Godavari-Krishna coastal sedimentary basins. Bulletin of Oil and Natural Gas Corporation, 17: 147-158.

Naqvi, S.M. and Rogers, J.J.W., 1987. Precambrian Geology of India. Oxford Monographs on Geology and Geophysics, Oxford University Press, Oxford, UK

Narasimha Chari, M. V., Sahu, J. N., Banerjee, B., Zutshi, P. L. and Chandra, K. 1995. Evolution of the Cauvery basin, India from subsidence modeling, Marine and Petroleum Geology, 12, 6, 667-675

Nemčok, M. and Rosendahl, B. R., 2006a. CameroonSpan interpretation report. EGI report No. 01-0059-5000-50501399 for GX Technology.

Nemčok, M. and Rosendahl, B. R., 2006b. GabonSpan interpretation report. EGI report No. 01-0059-5000-50501399 for GX Technology.

Nemčok, M., Sinha, S. T., Stuart, C. J. Welker, C., Choudhuri, M., Sharma, S. P., Misra, A. A., Sinha, N and Venkatraman, S., 2013a. East Indian margin evolution and crustal architecture: integration of deep reflection seismic interpretation and gravity modeling, in: Conjugate Divergent Margins (Eds: Mohriak, W.U., Danforth, Al; Post, P. J., Brown, D. E., Tari, G. T., Nemčok, M., and Sinha, S. T.) Geological Society, London, Special Publication, 369, 373–401doi: 10.1144/SP369.6

Nemčok, M., Stuart, C., Rosendahl, B. R., Welker, C., Smith, S., Sheya, C., Sinha, S.T., Choudhuri, M., Bird, D., Allen, R., Reeves, C., Sharma, S. P., Venkatraman, S. and N. Sinha, 2013b. Continental break-up mechanism; lessons from intermediate- and fast-extension settings, in: Conjugate Divergent Margins (Eds: Mohriak, W.U., Danforth, Al; Post, P. J., Brown, D. E., Tari, G. T., Nemčok, M., and Sinha, S.T.,) Geological Society, London, Special Publication. 369, 477–496http://dx.doi.org/10.1144/SP369.14

Nemčok, M., Henk, A., Allen, R., Sikora, P.J., and Stuart, C. 2013c. Continental break-up along strike-slip fault zones: observations from the Equatorial Atlantic, Geological Society, London, Special Publications, 369, 537-556, doi: 10.1144/SP369.8

References

298

Neogi, S. and Das, N., 1998. Lithotectonic domains and metamorphic history of the Boundary zone of the Eastern Ghats orogen and Bastar craton, Deobhog area, MadhyaPradesh, and its tectonic implications. Sem. Abst. Vol. Intl. Sem. Precambrian Crust in Eastern and Central India, BBSR, pp. 87-89.

Nicolaysen, K., Bowring, S., Frey, F., Weis, D., Ingle, S., Pringle, M., and Coffin, M. F., 2001. Provenance of Proterozoic garnet-biotite gneiss recovered from Elan Bank, Kerguelen Plateau, southern Indian Ocean, Geology, 29, 235-238.

O’Brien, P.E., and Stagg, H.M.J., 2007. Tectonic elements of the continental margin of East Antarctica, 38–164°E in: Cooper, A.K. and C.R. Raymond et al. (eds) Antarctica: A Keystone in a Changing World – Online Proceedings of the 10th ISAES, USGS Open-File Report 2007-1047, Short Research Paper 085, 4 p.; doi:10.3133/of2007-1047.srp085

O’Brien, P.E., Cooper, A.K., Richter, C. et al., 2001, Initial Reports, Prydz Bay-Cooperation Sea, Antarctica: glacial history and paleoceanography. Proceedings Ocean Drilling Program, Initial Reports, 188 (CD-ROM), Texas A & M University, College Station Texas

Odegard, M. E. 2003. Geodynamic evolution of the Atlantic Ocean: constraints from potential field data. (Abstract.) In: 8th International Congress of the Brazilian Geophysical Society, Rio de Janeiro. http://www.grizgeo.com/meo/pdfs/odegard_1098.pdf, accessed on January 27, 2012

Odegard, M.E., Dickson, W. G., Rosendahl, B.R., and Weger, R. E., 2002. Proto-oceanic crust in the north and south Atlantic: types, characteristics, emplacement mechanisms, and its Influence on Deep and Ultra-Deep Water Exploration (extended abstract). AAPG Hedberg Conference, Stavanger, Norway

Operto, S., and Charvis, P., 1996. Deep structure of the southern kerguelen Plateau (southern Indian ocean) from wide-angle seismic data, Journal of Geophysical Research, 101, 25077-25103

Peacock, D. C. P., 2004. The post-Variscan development of the British Isles within a regional transfer zone influenced by orogenesis, Journal of Structural Geology, 26, 12, 2225-2231

Peacock, D.C.P. and Sanderson, D.J., 1991. Displacements, segment linkage and relay ramps in normal fault zones. Journal of Structural Geology, 13, 6, 721-733.

Peacock, D.C.P., 2002. Propagation, interaction and linkage in normal fault systems. Earth-Science Reviews, 58, 1-2, 121-142.

Péron-Pinvidic, G., Manatschal, G., Gernigon, L., and Gaina, G., 2010. The formation and evolution of crustal blocks at rifted margins: new insights from the interpretation of the Jan Mayen microcontinent, Central & North Atlantic Conjugate Margins Conference, Lisbon, V, 231-235, http://metododirecto.pt/CM2010

Péron-Pinvidic, G., Manatschal, G., Minshull, T.A., and Sawyer, D.S., 2007. Tectonosedimentary evolution of the deep Iberia-Newfoundland margins: Evidence for a complex breakup history, Tectonics, 26, TC2011, 1-19, doi:10.1029/2006TC001970,

Plummer, Ph. S., and Belle, E. R., 1995. Mesozoic tectono-stratigraphic evolution of the Seychelles microcontinent, Sedimentary Geology, 96, 1-2, 73-91, doi: 10.1016/0037-0738(94)00127-G

Prabhakar, K.N. and Zutshi, P.L. 1993. Evolution of southern part of Indian east coast basins. Jour. Geol. Soc. India, 41, 215-230

References

299

Radhakrishna, B. P. and Naqvi, S. M., 1986. Precambrian continental crust of India and its evolution. The Journal of Geology, 94, 2, 145–166

Radhakrishna, M., Chand, S. & Subrahmanyam, C. 2000. Gravity anomalies, sediment loading and lithospheric flexure associated with the Krishna–Godavari basin, eastern continental margin of India. Earth and Planetary Sciences Letters, 175, 223–232

Ramakrishanan, M. and Vaidyanadhan, R., 2010. Geology of India (Volume 1), Geological Society of India, ISBN: 978-81-85867-98-4.

Ramakrishnan, M., 2003. Craton-mobile belt relations in Southern Granulite Terrain; Tectonics of Southern Granulite Terrain; Kuppam-Palani geotransect. Geological Society of India Memoir, 50, 1-24.

Ramakrishnan, M., Nanda, J. K. and Augustine, P. F., 1998. Geological evolution of the Proterozoic Eastern Ghats Mobile Belt. Geological Survey of India Special Publications, 44, 1-21.

Ramana, M. V., Ramprasad, T., Desa, M., and Subrahmanyam, V., 2000. Integrated geophysical studies over the 85°E ridge - Evaluation and interpretation. Visakha Science Journal 4 (1), 45–56, http://drs.nio.org/drs/handle/2264/428

Ramana, M.V, Subrahmanyam, V., Chaubey, A. K. Ramprasad, T., Sarma, K.V.L.N.S., Krishna, K.S., Desa, M., Murty, G.P.S. and Subrahmanyam. C, 1997. Structure and origin of the 85°E Ridge, Journal of Geophysical Research., 102, B8,17,995-18012.

Ramana, M.V., Krishna, K.S., Ramprasad, T., Desa, M., Subrahmanyam, V., and Sarma, K.V.L.N.S., 2001. Structure and tectonic evolution of the northeastern Indian Ocean. in: SenGupta, R., and Desa, Ehrlich (Eds) The Indian Ocean: A perspective. Oxford & IBH, 2,731-816.

Ramana, M.V., Nair, R.R. and Sarma, K.V.L.N.S., 1994. Mesozoic anomalies in the Bay of Bengal. Earth Planetary Sciences Letters, 121, 469–475.

Rao, G. D., Krishna, K. S. & Sar, D., 1997. Crustal evolution and sedimentation history of the Bay of Bengal since the Cretaceous. Journal of Geophysical Research, 102, 17747–17768

Rao, G. N., 1993. Geology and hydrocarbon prospects of east coast sedimentary basins of India with special reference to Krishna Godavari basin. Journal of the Geological Society of India, 41, 444-454.

Rao, G. N., 2001. Sedimentation, stratigraphy, and petroleum potential of Krishna-Godavari basin, East Coast of India. AAPG Bulletin, 85, 1623-1643.

Rao, Y. S. N., 1980. Geology of the eastern margin of the Indian subcontinent. SEAPEX Proceedings, vol. V, pp. 161-178.

Rapolla, A., Cella, F., and Dorre, A. S., 1995. Gravity study of the crustal structures of Somalia along International Lithosphere Program geotransects, Journal of African Earth Sciences, 20, 3-4, 263-274

Ravnås, R. and Steel, R. J., 1998. Architecture of Marine Rift-Basin Successions, AAPG Bulletin, 82, 1, 110–146.

Ray, D. K., 1963. Tectonic map of India. Geological Society of India, Calcutta, scale 1:2 000 000.

Reeves, C. 2009. Re-examining the evidence from plate-tectonics for the initiation of Africa’s passive margins, Geological Society of Houston/Petroleum Exploration Society of Great Britain, London, extended abstract

References

300

Reeves, C., and Wit, M. D., 2000. Making ends meet in Gondwana: retracing the transforms of the Indian Ocean and reconnecting continental shear zones, Terra Nova, 12, 6, 272-280

Reeves, C., Karanja F. M., and MacLeod, I. N., 1986. Geophysical evidence for a failed Jurassic rift and triple junction in Kenya, Earth and Planetary Science Letters, 81, 299-311.

Reeves, C.V., de Wit, M.J., Sahu, B.K., 2004. Tight reassembly of Gondwana exposes Phanerozoic shears in Africa as global tectonic players. Gondwana Research 7, 7–19.

Reeves, C.V., Sahu, B.K., de Wit, M.J., 2002. A re-examination of the paleo-position of Africa's eastern neighbours in Gondwana. Journal of African Earth Sciences 34, 101–108.

Reeves. C,. 2008. A plate-tectonic framework for the evolution of the passive margins of India, Reliance internal report, unpublished

Reston, T. J., 2007. The formation of non-volcanic rifted margins by the progressive extension of the lithosphere: the example of the West Iberian margin, in: Karner, G.D, Manatschal, G., and Pinheiro, L.M (eds), Imaging, Mapping and Modeling of Continental Lithospheric Extension and Breakup, Geological Society, London, Special Publications, 282, 77-110, doi:10.1144/SP282.5

Reston, T.J., 1996. The S reflector west of Galicia; the seismic signature of a detachment fault. Geophysical Journal International, 127(1): 230-244

Rey, S. S., Eldholm, O., and Planke, S., 2003. Formation of the Jan Mayen Microcontinent, the Norwegian Sea, American Geophysical Union, Fall Meeting 2003, abstract #T31D-0872.

Rickers, K., Mezger, K., and Raith, M.M., 2001. Evolution of the Continental Crust in the Proterozoic Eastern Ghats Belt, India and new constraints for Rodinia reconstruction: implications from Sm–Nd, Rb–Sr and Pb–Pb isotopes, Precambrian Research, 112, 3–4, 183-210

Roest, W. R., Lundin, E. R., Torsvik, T. H., and Olesen, O., 2002. The Jan Mayen Microcontinent: Computers animations of the plate tectonic history, American Geophysical Union, Fall Meeting, abstract #T12D-1344

Rosenbaum, G., Weinberg, R. F. and Regenauer-Lieb, K., 2008. The geodynamics of lithospheric extension, Tectonophysics, 458, 1-4, 1-8. doi:10.1016/j.tecto.2008.07.016.

Rosendahl, B. R. and Groschel-Becker, H., 1999. Deep seismic structure of the continental margin in the Gulf of Guinea; a summary report. In: Cameron, N. R., Bate, R. H. and Clure, V. S. (Eds.), The oil and gas habitats of the South Atlantic. Geological Society of London Special Publications, 153, 75-83.

Rosendahl, B. R., Mohriak, W. U., Nemčok, M., Odegard, M. E., Turner, J. P. and Dickson, W. G., 2005. West African and Brazilian conjugate margins; crustal types, architecture, and plate configurations. In: Post, P. J., Rosen, N. C., Olson, D. L., Palmes, S. L., Lyons, K. T. and Newton, G. B. (Eds.), Petroleum systems of divergent continental margin basins. Conference 25th annual Gulf Coast section SEPM Foundation, Bob F. Perkins research conference, symposium on Petroleum systems of divergent continental margin basins, Houston, TX, United States, Dec. 4-7, 2005, Program and Abstracts - Society of Economic Paleontologists, Gulf Coast Section, 25, 13-14.

Rotstein, Y., Munschy, M., Bernardel, A., 2001. The Kerguelen Province revisited: additional constraints on the early development of the Southeast Indian Ocean. Marine Geophysical Researches 22, 81–100

References

301

Rotstein, Y., Schlich, R., Munschy, M., and Coffin, M.F., 1992. Structure and tectonic history of the southern Kerguelen Plateau (Indian Ocean) deduced from seismic reflection data. Tectonics, 11, 1332-1347

Royer, J.Y., and Chang, T., 1991. Evidence for relative motions between the Indian and Australian plates during the last 20 Myr from plate tectonic reconstructions: Implications for the deformation of the Indo-Australian plate, Journal of Geophysical Research, 96, 11779–11802

Sandwell, D. T., and W. H. F. Smith, 2009. Global marine gravity from retracked Geosat and ERS-1 altimetry: Ridge Segmentation versus spreading rate, Journal of Geophysical Research, 114, B01411, doi:10.1029/2008JB006008, 2009.

Sarkar, A., Datta, A. K., Poddar, B. C., Bhattacharyya, B. K., Kollapuri, V. K. And Sanwall, R., 1996. Geochronological studies of Mesozoic igneous rocks from eastern India. Journal of Southeastern Asian Earth Sciences, 13: 77-81.

Sastri, V.V., Venkatachala, B.S. and Narayanan, V., 1981. The evolution of the east coast of India, Palaeogeography. Palaeoclimatology and Palaeoecology, 36, 23-54

Schettino, A., 1998. Computer aided paleogeographic reconstructions, Computers and Geosciences, 24,3, 259-267

Schettino, A., and Scotese, C. R., 2005. Apparent polar wander paths for the major continents (200 Ma to the present day): a palaeomagnetic reference frame for global plate tectonic reconstructions Geophysical Journal International, 163, 727-759

Scotese, C. R., Gahagan, L.M. and Larson, R.L., 1988. Plate tectonic reconstructions of the Cretaceous and Cenozoic ocean basins, Tectonophysics, 155, 27-48

Scotese, C., 2003. Paleomap projects, http://www.scotese.com/Default.htm (as on 19/11/2011)Scotese, C. R., Gahagan, L. M., and Larson, R. L., 1988. Plate tectonic reconstructions of the Cretaceous and Cenozoic ocean basins, Tectonophysics, 155, 27-48

Scott, R.A., Ramsey, L. A., Jones, S. M., Sinclair, S., and Pickles, C.S., 2005. Development of the Jan Mayen microcontinent by linked propagation and retreat of spreading ridges, Onshore-Offshore Relationships on the North Atlantic Margin, Proceedings of the Norwegian Petroleum Society Conference, Norwegian Petroleum Society Special Publications, 12, 69-82, doi:10.1016/S0928-8937(05)80044-X

Searle, R.C., Cowie, P.A., Mitchell, N.C., Allerton, S. C., MacLeod, J., Escartin, J., Russell, S.M., Slootweg, P.A., and Tanaka, T., 1998. Fault structure and detailed evolution of a slow spreading ridge segment: the Mid-Atlantic Ridge at 29°N, Earth and Planetary Science Letters, 154 1-4, 167-183

Sempéré, J.C., Blondel, P., Briais, A., Fujiwara, T., Geli, L., Isezaki, N., Pariso, J.E., Parson, L., Patriat, P., and Rommevaux, C., 1995. The Mid-Atlantic Ridge between 29°N and 31°30′N in the last 10 Ma, Earth Planetary Sciences Letters, 130, 45–55

Sen, G. 2001. Generation of Deccan Trap magmas, Proc. Ind.Acad. Science (Earth and Planetary Sci.), 110,409-431

Sengupta, S., 1966. Geological and geophysical studies in western part of Bengal Basin, India. AAPG Bulletin, 50: 1001-1017

Seton, M., Müller, R.D., Zahirovic, S., Gaina, C., Torsvik, T.H., Shephard, G., Talsma, A., Gurnis, M., Turner, M., Maus, S., Chandler, M., 2012., Global continental and ocean basin reconstructions since 200 Ma Earth-Science Reviews, 113, 3-4, 212-270, doi:10.1016/j.earscirev.2012.03.002

References

302

Shah, J., Srivastava, D. C., Pandian, M.S., Sarkar, S., Choudhuri, M., and Subramanian, V. 2007, Mesoscale fractures as paleostress indicators: A case study from Cauvery basin. Journal of Geological Soceity of India, 70, 571-583.

Shertaon, J.W., Tingey, R.J., Balck, L.P., Offe, L.A., and Ellis, D.J., 1987. Geology of an unusual Precambrian high-grade metamorphic terrane –Enderby Land and westerm Kemp Land, Antarctica, Bulletin-223, Australian Government Publishing Service, 3-23

Sheth, H. C., 2006. The Deccan Beyond the Plume Hypothesis. MantlePlumes.org,

Sheth, H. C., Mahoney, J. J. and Chandrasekharam, D., 2004. Geochemical stratigraphy of Deccan flood basalts of the Bijasan Ghat section, Satpura Range, India. Journal of Asian Earth Sciences, 23, 127-139.

Sibuet, J.C., S. Srivastava, and G. Manatschal, 2007. Exhumed mantle forming transitional crust in the Newfoundland-Iberia rift and associated magnetic anomalies. Journal of Geophysical Research, 112, B06105, doi: 10.1029/ 2005JB003856

Simon, K., Huismans, R.S., and Beaumont, C., 2009. Dynamical modeling of lithospheric extension and small-scale convection: implications for magmatism during the formation of volcanic rifted margins, Geophysical Journal International, 176, 327–350, doi: 10.1111/j.1365-246X.2008.03891.x.

Singh, A. P., 2002. Impact of Deccan volcanism on deep crustal structure along western part of Indian mainland and adjoining Arabian Sea. Curr. Sci., 82: 316-325.

Sinha, S. T., Nemcok, M., Choudhuri, M., Misra, A. A., Sharma, S. P., Sinha, N. & Venkatraman, S., 2010. The crustal architecture and continental break up of East India Passive margin: an integrated study of deep reflection seismic interpretation and gravity modeling, AAPG Annual Convention & Exhibition, New Orleans, USA, CD-Rom. http://www.searchanddiscovery.com/documents/2010/40611sinha/ndx_sinha.pdf

Skogseid, J., 2001. Volcanic margins: geodynamic and exploration aspects, Marine and Petroleum Geology, 18, 4, 457-461. doi:10.1016/S0264-8172(00)00070-2

Skogseid, J., and Eldholm, O., 1987. Early Cenozoic crust at the Norwegian continental margin and the conjugate Jan Mayen Ridge, Journal of Geophysical Research., 92, 11471-11491

Smallwood, J. R. and White, R. S., 2002. Ridge-plume interaction in the North Atlantic and its influence on continental breakup and seafloor spreading, in: Jolley, D.W. and Bell, B.R. (eds.), The North Atlantic igneous province; stratigraphy, tectonic, volcanic and magmatic processes, Geological Society, London, Special Publications, 197, 15-37

Smith, W. H. F., and D. T. Sandwell, 1997. Global seafloor topography from satellite altimetry and ship depth soundings, Science, 277, 957-1962,

Song, T. and Cawood, P. A.,2000. Structural styles in the Perth Basin associated with the Mesozoic break-up of Greater India and Australia, Tectonophysics. 317, 1-2, 55-72

Srivastava, D.C., and Shah, J., 2007. Regional Geological Studies in East Coast of India, Geological and Structural analysis of Krishna-Godavari basin, Reliance Industries Internal report, unpublished.

Srivastava, S.P., Sibuet, J.-C., Cande, S., Roest, W.R., and Reid, I.D., 2000. Magnetic evidence for slow seafloor spreading during the formation of the Newfoundland and Iberian margins. Earth Planet. Sci. Lett., 182(1):61–76. doi:10.1016/S0012-821X(00)00231-4

References

303

Srivastava, V. K. and Chowhan, R. K. S., 1987. Study of focal mechanism solutions of events distributed in and around the India lithospheric plate. Oxford and IBH Publishing, New Delhi, pp. 139-156.

Stagg, H.M.J., 1985, The structure and origin of Prydz Bay and the Mac.Robertson Shelf, East Antarctica, Tectonophysics, 114, 315–340

Stagg, H.M.J., J. B. Colwell, N. G. Direen, P. E. O’Brien, B. J. Brown, G. Bernardel, I. Borissova, L. Carson, L., and D. B. Close, 2005. Geological framework of the continental margin in the region of the Australian Antarctic Territory, Geoscience Australia Record 2004/5

Steckler, M.S., and ten Brink, U.S., 1986, Lithospheric strength variations as a control on new plate boundaries: Examples from the northern Red Sea region, Earth and Planetary Science Letters, 79, 120–132

Stewart, J., Watts, A. B. and Bagguley, J. O., 2000. Three-dimensional subsidence analysis and gravity modeling of the continental margin offshore Namibia. Geophysical Journal International, 141, 3, 724-746

Storey, B.C., 1995. The role of mantle plumes in continental break up, case histories from Gondwanaland. Nature, 377, 301-308

Storey, M., Kent, R. W., Saunders, A. D., Salters, V. J., Hergt, J., Whitechurch, H., Sevigny, J. H., Thirlwall, M. F., Leat, P., Ghose, N. C., and Gifford, M., 1992. Lower cretaceous volcanic rocks on continental margins and their Relationship to the Kerguelen Plateau, in: Wise, S. W., Jr., Schlich, R., et al., (eds), Proceedings of the Ocean Drilling Program, Scientific Results, 120

Subramanyan, C. & Chand, S. 2006. Evolution of passive continental margins of India-a geophysical appraisal. Gondwana Research, 10, 167–178.

Sutra, E., Manatschal, G., Mohn, G., and Unternehr, P., 2013. Quantification and restoration of extensional deformation along the Western Iberia and Newfoundland rifted margins, Geochem. Geophys. Geosyst, 14, 8, doi: 10.1002/ggge.20135

Tanzania Petroleum Development Corporation Promotion Brochure. 2009., Tanzania Petroleum Development Corporation 1-50.

Tiwari, R. S. and Tripathi, A., 1995. Palynological assemblages and absolute age relationship of Intertrappean beds in the Rajmahal Basin, India. Cretaceous Research, 16: 53-72.

Torsvik, T.H., Mosar, J. and Eide, E.A., 2001. Cretaceous-Tertiary Geodynamics: A North Atlantic exercise, Geophysical Journal International, 146, 850–866Vine, F. J., 1966. Spreading of the Ocean Floor: New Evidence, Science, 154, 3755, 1405–1415

Toulokian, Y. S., Judd, W. R.&Roy, R. F. 1981. Physical Properties of Rocks and Minerals. McGraw-Hill, New York.

Truswell, E.M., Dettmann, M.E. and O’Brien, P.E., 1999, Mesozoic palynofloras from the Mac. Robertson shelf, East Antarctica: geological and phytogeographic implications, Antarctic Sci. 11, 237–252

Tucholke, B. E., and Sibuet, J. C., 2007. Leg 210 synthesis: tectonic, magmatic, and sedimentary evolution of the Newfoundland- Iberia rift, in: Tucholke, B. E., Sibuet, J.-C., and Klaus. A., (eds), Proceedings of the Ocean Drilling Program, Scientific Results, 210

Turcotte, D.L. and Emerman, S.H., 1983. Mechanisms of active and passive rifting. In: Morgan P. and B.H. Baker (Eds), Tectonophysics, 39-50.

References

304

Turner, B.R., 1991. Depositional environment and petrography of pre-glacial continental sediments from hole 740a, Prydz Bay, East Antarctica, Proceedings of the Ocean Drilling Program, Scientific Results, 119

Valdiya, K.S, 2010. The making of Inida, geodynamic evolution, McMillan Publishers India, ISBN: 978-0230-32833-4,

Van Wagoner, J.C., 1995. Overview of sequence stratigraphy of foreland basin deposits: terminology, summary of papers, and glossary of sequence stratigraphy, in: Van Wagoner, J.C., and Bertram, G.T. (eds.), Sequence Stratigraphy of Foreland Basin Deposits: Outcrop and Subsurface Examples from the Cretaceous of North America, Memoir-64, American Association of Petroleum Geologists, ix–xxi

Veevers, J. J. and Tewari, R. C., 1995. Gonwana master basin of Peninsular India between Tethys and rthe interior of the Gondwanaland province of Pangea. GSA Bulletin, 187, p. 1-72.

Veevers, J.J., 2009. Palinspastic (pre-rift and -drift) fit of India and conjugate Antarctica and geological connections across the suture, Gondwana Research, 16, 90–108, doi:10.1016/j.gr.2009.02.007

Verzhbitsky, E. V., 2003. Geothermal regime and genesis of the Ninety-East and Chagos-Laccadive ridges. Journal of Geodynamics, 35, 3, 289–302. doi:10.1016/S0264-3707(02)00068-6

Vine, F. J. and Matthews, D. H., 1963. Magnetic Anomalies over Oceanic Ridges, Nature 199, 4897, 947–949

Vine, F.J., 1966. Spreading of the ocean floor; new evidence. Science, 154,3755, 1,405-1,415.

Watts, A. B., 2001. Gravity anomalies, flexure and, crustal structure at the Mozambique rifted margin, Marine and Petroleum Geology, 18, 445-455,

Watts, A. B., Rodger, M., Peirce, C., Greenroyd, C., and Hobbs R. W., 2009. Seismic structure, gravity anomalies and flexure of the Amazon continental margin, NE Brazil, J. Geophys. Res., doi:10.1029/2008JB006259.

Weis, D., Ingle, S., Damasceno, D., Frey, F. A., Nicolaysen, K. and Barling, J. 2001. Origin of continental components in Indian Ocean basalts: evidence from Elan Bank (Kerguelen Plateau - ODP Leg 183, Site 1137), Geology, 29, 147-150

Weissel, J. and Karner, G. D., 1989. Flexural uplift of rift flanks due to mechanical unloading of the lithosphere during extension. Journal of Geophysical Research 94, doi: 10.1029/89JB01075.

White, R. and McKenzie, D., 1989. Magmatism at rift zones; the generation of volcanic continental margins and flood basalts. In: W.P. Leeman and J.G. Fitton (Editors), Journal of Geophysical Research, pp. 7685-7729

White, R.S., McKenzie, D. and O'Nions, R.K., 1992. Oceanic crustal thickness from seismic measurements and rare earth element inversions, Journal of Geophysical Research, 97(B13), 19,683-19,715

White, R.S., Spence,G.D. , Fowler, S.R. ,McKenzie, D.P., Westbrook, G.K., Bowen, A.N., 1987. Magmatism at rifted continental margins,Nature, 330, 439–444.

Whitmarsh, R. B., and Manatschal, G., 2012. Evolution of magma poor continental margins: from rifting to the onset of seafloor spreading, in: Roberts, G.D and Bally, A,W., (eds), Phanerozoic passive margins, cratonic basins and Gobal tectonic maps, V: 1c, Elsevier Publications, 303-317

Whitmarsh, R.B., G. Manatschal, and T. A. Minshull, 2001. Evolution of magma-poor continental margins from rifting to seafloor spreading. Nature, 413: 150-154.

References

305

Whittaker, J., Williams, S., Kusznir, N., Müller, R.D., 2010. Restoring the continent-ocean boundary: constraints from lithospheric stretching grids and tectonic reconstructions, ASEG 2010 - Sydney, Australia Abstract guidelines

Wilson, P. G., Turner, J. P. and Westbrook, G. K., 2003. Structural architecture of the ocean-continent boundary at an oblique transform margin through deep-imaging seismic interpretation and gravity modeling; Equatorial Guinea, West Africa. Tectonophysics, 374, 1-2, 19-40.

Wilson, S.D., 1990. Kinematics of overlapping rift propagation with cyclic rift failure, Earth and Planetary Science Letters, 96 (3-4), 384-392

Wise, S. W., Coxall, H. K., Wahnert, V., Inokuchi, H., Coffin, M. F., Frey, F.A., Wallace, P. J., and the ODP Leg 183 Scientific Party, 2000. The Kerguelen Plateau: new palaeontologic and paleomagnetic age constraints on growth history from ODP Leg 183 drilling, Proc. ODP, Initial Rep., 183, Ocean Drilling Program, Texas A&M University, College Station, TX

Wu, D. and Bruhn, R.L., 1994. Geometry and kinematics of active normal faults, South Oquirrh Mountains, Utah; implication for fault growth. Journal of Structural Geology, 16(8): 1061-1075

Wyer, P. and A. B. Watts, "Gravity anomalies and segmentation at the East Coast, USA continental margin", Geophys. J. International, 166, 1015-1038, 2006

Yoshida, M., 1995. Assembly of East Gondwananland during the Mesoproterozoic and its rejuvenation during the Pan-African period. In: Yoshida, M. and Santosh, M. (Eds.), India and Antarctica during Precambrian. Geological Society of India Memoir, 34, pp. 25-45.

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