The SE Asian Gateway: History and Tectonics of the Australia-Asia collision (Geological Society Special Publication 355)

The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) RICK LAW (USA) PHIL LEAT (UK) NICK ROBINS (UK) RANDELL STEPHENSON (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) GONZALO VEIGA (ARGENTINA ) MAARTEN DE WIT (SOUTH AFRICA )

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk. It is recommended that reference to all or part of this book should be made in one of the following ways: HALL , R., COTTAM , M. A. & WILSON , M. E. J. (eds) 2011. The SE Asian Gateway: History and Tectonics of the Australia–Asia Collision. Geological Society, London, Special Publications, 355. WIDIYANTORO , S., PESICEK , J. D. & THURBER , C. H. 2011. Subducting slab structure below the eastern Sunda arc inferred from non-linear seismic tomographic imaging. In: HALL , R., COTTAM , M. A. & WILSON , M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 139–155.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 355

The SE Asian Gateway: History and Tectonics of the Australia – Asia Collision EDITED BY

R. HALL Royal Holloway University of London, UK

M. A. COTTAM Royal Holloway University of London, UK

and M. E. J. WILSON Curtin University, Australia

2011 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 10 000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ 44 (0)20 7434 9944; Fax þ 44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ 44 (0)1225 445046, Fax þ 44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2011. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/11/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-329-5 Typeset by Techset Composition Ltd, Salisbury, UK Printed by MPG Books Ltd, Bodmin, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ 1 800-972-9892 Fax þ 1 802-864-7626 E-mail: [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ 1 918-584-2555 Fax þ 1 918-560-2652 E-mail: [email protected] Website: http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ 91 11 2327-9113/2326-4180 Fax þ 91 11 2326-0538 E-mail: [email protected]

Contents HALL, R., COTTAM, M. A. & WILSON, M. E. J. The SE Asian gateway: history and tectonics of the Australia –Asia collision

1

METCALFE, I. Palaeozoic –Mesozoic history of SE Asia

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CLEMENTS, B., BURGESS, P. M., HALL, R. & COTTAM, M. A. Subsidence and uplift by slab-related mantle dynamics: a driving mechanism for the Late Cretaceous and Cenozoic evolution of continental SE Asia?

37

GRANATH, J. W., CHRIST, J. M., EMMET, P. A. & DINKELMAN, M. G. Pre-Cenozoic sedimentary section and structure as reflected in the JavaSPANTM crustal-scale PSDM seismic survey, and its implications regarding the basement terranes in the East Java Sea

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HALL, R. Australia –SE Asia collision: plate tectonics and crustal flow

75

KOPP, H. The Java convergent margin: structure, seismogenesis and subduction processes

111

WIDIYANTORO, S., PESICEK, J. D. & THURBER, C. H. Subducting slab structure below the eastern Sunda arc inferred from non-linear seismic tomographic imaging

139

WATKINSON, I. M. Ductile flow in the metamorphic rocks of central Sulawesi

157

COTTAM, M. A., HALL, R., FORSTER, M. A. & BOUDAGHER-FADEL, M. Basement character and basin formation in Gorontalo Bay, Sulawesi, Indonesia: new observations from the Togian Islands

177

WATKINSON, I. M., HALL, R. & FERDIAN, F. Tectonic re-interpretation of the Banggai-Sula –Molucca Sea margin, Indonesia

203

RIGG, J. W. D. & HALL, R. Structural and stratigraphic evolution of the Savu Basin, Indonesia

225

AUDLEY-CHARLES, M. G. Tectonic post-collision processes in Timor

241

TILLINGER, D. Physical oceanography of the present day Indonesian Throughflow

267

HOLBOURN, A., KUHNT, W. & XU, J. Indonesian Throughflow variability during the last 140 ka: the Timor Sea outflow

283

VON DER

HEYDT, A. S. & DIJKSTRA, H. A. The impact of ocean gateways on ENSO variability in the Miocene

305

MORLEY, R. J. & MORLEY, H. P. Neogene climate history of the Makassar Straits, Indonesia

319

LELONO, E. B. & MORLEY, R. J. Oligocene palynological succession from the East Java Sea

333

WILSON, M. E. J. SE Asian carbonates: tools for evaluating environmental and climatic change in equatorial tropics over the last 50 million years

347

Index

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The SE Asian gateway: history and tectonics of the Australia– Asia collision ROBERT HALL1*, MICHAEL A. COTTAM1 & MOYRA E. J. WILSON2 1

SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK 2

Department of Applied Geology, Curtin University, GPO Box U1987, Perth, WA 6845, Australia *Corresponding author (e-mail: [email protected])

The SE Asian gateway is the connection from the Pacific to the Indian Ocean and it has diminished from a wide ocean to a complex narrow passage with deep barriers (Gordon et al. 2003) as plate movements caused Australia to collide with SE Asia. It is one of several major ocean passages that existed during the Cenozoic but has received much less attention than others that opened, such as the Drake Passage, Tasman Gateway, Arctic Gateway or Bering Straits, or that closed, such as the Panama Gateway or Tethyan Gateway (e.g. von der Heydt & Dijkstra 2006; Lyle et al. 2007, 2008). It is not entirely clear why there has been this comparative neglect, but it may reflect the relative limited knowledge of the large and remote areas of Indonesia and the western Pacific, in particular their geological history, and the relatively small number of active researchers in this large region. Unlike the Panama Gateway and Tethyan Gateway the SE Asian gateway is still partly open and the ocean currents that flow between the Pacific and Indian Oceans have been the subject of much recent work by oceanographers (e.g. Gordon 2005). We now know that the Indonesian Throughflow, the name given to the waters that pass through the only remaining low latitude oceanic passage on the Earth, plays an important role in Indo-Pacific and global thermohaline flow (Gordon 1986; Godfrey 1996), and it is therefore probable that the gateway is important for global climate (Schneider 1998). It is also known that today the region around the SE Asian gateway contains the maximum global diversity for many marine (Tomascik et al. 1997) and terrestrial organisms (Whitten et al. 1999a, b). It is not known when and why this diversity originated, if there is a connection between biotic diversity and oceanography, what is the role of the throughflow in the modern climate system, and how the restriction and almost complete closure of the passage between the Pacific and Indian Oceans may be linked to the history of climate change. However, all of these are likely

consequences of, or related to, the closure of the wide ocean that separated Australia and SE Asia at the beginning of the Cenozoic. The gateway closure was caused by the tectonic changes accompanying the northward movement of Australia as it converged with Asia. Collision between Australia and SE Asia began in the earliest Miocene. The gateway was fully open before 25 Ma and significantly restricted by 5 Ma but understanding its history requires detailed reconstruction of an area of great geological complexity (Hall 2002; Kuhnt et al. 2004). Biogeographers have given the name Wallacea to the area bounded by the Wallace Line in the west, marking an eastern limit of truly Asian faunas and floras, and Lydekker’s Line in the east, which is the western boundary of Australasian faunas and floras, and Wallacea is at the centre of the SE Asian Gateway. Wallacea includes the islands of Sulawesi and the Banda Arc and is marked by high numbers of endemic species, complex distribution patterns, and unusual variations in species richness (Whitmore 1987). The biogeographical complexity reflects the significant changes in distribution of land and sea during the Neogene which in turn reflects the complex geological history of Wallacea, largely driven by subduction, and the rapid changes that have occurred – for example, some of the largest islands in the Banda Arc, such as Seram and Timor, have emerged from the sea only in the last 3 million years (Hall 2001). The Neogene history of the gateway records a complex history of rapid changes in tectonics, topography and land/sea distributions (Hall 1998). As the deepwater connection closed, mountains rose, there was an increase in land and shelf areas, but new deep basins also formed. There were numerous changes that accompanied the closure. High mountains rose first in Borneo and later elsewhere in Indonesia. Rainfall and erosion rates must have changed. Changes in geologically-controlled passages would have influenced oceanic circulation. There was a change from warm South Pacific to

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 1–6. DOI: 10.1144/SP355.1 0305-8719/11/$15.00 # The Geological Society of London 2011.

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colder North Pacific waters passing through the gateway. There was a change from drier to wetter climate. This tropical gateway is likely just as important as the opening and closure of other oceanic passages and the rise of Tibet for the global climate system. The physical changes influenced biogeography of the SE Asian and West Pacific region. There were major changes in carbonate depositional systems, including increased development of coral reefs (Wilson 2008) and the present-day global centre of biodiversity in some way reflects the interplay of geology, oceanography and climate (Wilson & Rosen 1998; Renema et al. 2008). The connections between geology and biodiversity were the subject of a Geological Society of London conference held at Royal Holloway University of London in September 2009. The meeting aimed to bring together a range of scientists from a variety of disciplines in Earth and life sciences to better understand the geological history of the gateway, the causes and timing of its closure, and their effects. This Special Publication includes

papers by predominantly physical science contributors to the meeting and a second volume will contain papers mainly by life scientists (Gower et al. 2011). The papers in this volume (Fig. 1) have been arranged to first explain and discuss the Palaeozoic and Mesozoic geological development of the region, and then its Cenozoic history, which provide the background to understand the present Indonesian Throughflow, oceanographic changes since the Neogene, and finally some aspects of the climate history.

Pre-Cenozoic geological history One of the important factors that influenced the late Cenozoic history of the SE Asian gateway was the complex structure of the basement acquired since the late Palaeozoic. Metcalfe reviews the fragmentation of Gondwana and assembly of Gondwana fragments in SE Asia, accompanied by the closure of Tethyan oceans. He highlights a number of significant recent changes in our understanding

Fig. 1. Numbered boxes show the areas discussed in the papers in this volume. 1, Metcalfe; 2, Clements et al.; 3, Granath et al.; 4, Hall; 5, Kopp; 6, Widyantoro et al.; 7, Watkinson; 8, Cottam et al.; 9, Watkinson et al.; 10, Rigg & Hall; 11, Audley-Charles; 12, Tillinger; 13, Holbourn et al.; 14, van der Heydt & Dijkstra; 15, Morley & Morley; 16, Lelono & Morley; 17, Wilson. The background image is a digital elevation model of SE Asia based on satellite gravity-derived bathymetry combined with SRTM (Shuttle Radar Topographic Mission) topography (Sandwell & Smith 2009).

INTRODUCTION

of SE Asian basement structure, notably recognition of the importance of the Sukhothai arc and associated terranes that separated Sibumasu and Indochina in the Permian and Triassic, the addition of the West Sumatra and West Burma blocks in the PermoTriassic, and the identification of Borneo, Java and West Sulawesi as the Argoland blocks rifted from NW Australia in the Late Jurassic. The collision of these Australian fragments was complete by the mid Cretaceous and subduction ceased around Sundaland until the Eocene. Clements et al. propose that a widespread regional unconformity was a dynamic topographic response to termination of subduction and is the reason for the almost complete absence of Upper Cretaceous and Paleocene rocks throughout most of Sundaland. Sedimentation began again in the Eocene when subduction resumed. Based on new regional deep seismic surveys Granath et al. show that beneath the unconformity in the Java Sea is an unexpected and thick sedimentary section which is probably Precambrian to Triassic in age and was deposited when this basement block was still part of the Australian margin.

Cenozoic subduction During the Eocene, subduction beneath Indonesia began as Australia moved north, gradually closing the deep passage linking the Pacific and Indian Oceans. In the Early Miocene the leading edge of the Australian continent began to collide with the SE Asian margin in East Indonesia. Hall interprets the development of the Neogene collision to have been strongly influenced by the shape of the Australian continental margin, due to Jurassic rifting, and the presence of an oceanic embayment, leading to subduction rollback into the embayment. However, young deformation is also attributed to a component of lower crustal flow which has enhanced the effects of sediment loading and driven uplift and exhumation of mountains in northern Borneo and Sulawesi. Subduction has been the most important tectonic driving force of change in Indonesia but its consequences are very varied. Kopp reviews subduction along the Java margin and shows how the deep structure of the margin varies from west to east. Features of the margin, its seismogenic character and seismic hazards reflect many factors such as sediment supply, relief of the subducting slab, and geometry of the plate interface. The still deeper structure of the subduction zone can be imaged from P wave seismic tomography and Widiyantoro et al. show how this reveals a complex geometry with a hole in the subducting slab beneath East Java, and a possible tear beneath eastern Indonesia. This provides valuable information which can be related to the surface expression of subduction, such as deformation in

3

the volcanic arc and variation in history and type of volcanic activity.

Sulawesi and Neogene tectonics Sulawesi is situated close to the centre of Wallacea and includes parts of the pre-Neogene Sundaland margin and Australian crust that was added in the Cretaceous and the Neogene. West Sulawesi is cut by the Palu-Koro Fault which is an active strike-slip fault with spectacular surface expression, important seismicity and clear evidence of young deformation. It represents a potentially major hazard but little is known of its history. Watkinson shows that deformed rocks close to the fault reveal a complex structure and ductile deformation which must pre-date Pliocene to present-day movement on it. Just east of the Palu-Koro Fault is the wide Gorontalo Bay, also known as Tomini Bay, which is one of the enigmatic inter-arm bays that give Sulawesi its unusual K-shape. Seismic and multibeam data have recently been acquired from the bay and provide almost the first information about the submarine parts of it, but there are a number of small islands including Una-Una volcano that erupted destructively in 1983, and the Togian group, that reveal some of its history. The Togian Islands have been reported to include igneous rocks that have been interpreted as part of the Cretaceous – Eocene East Arm ophiolite, or as younger volcanic rocks of uncertain tectonic setting. Cottam et al. report new observations from the Togian Islands that contribute to understanding the origin of Gorontalo Bay, its basement and the volcanic history of the area. They show that there is an old volcanic basement, probably Palaeogene, that the southern part of the bay was close to sea level during the Middle Miocene when shallow marine carbonates were deposited, and there were explosive eruptions from a nearby volcano and marine deposition of reworked volcanic ash in the Pliocene. However, alluvial fan deposits show that there were Pliocene connections to the East Arm from which the islands are now separated by a deep marine basin implying rapid and large subsidence of parts of Gorontalo Bay. The eastern end of the East Arm includes the collisional contact of the ophiolite with one of the microcontinental fragments of eastern Indonesia: the Banggai-Sula block. Banggai-Sula and other microcontinental blocks have long been interpreted as sliced from the Bird’s Head of New Guinea and carried east in the left-lateral Sorong Fault Zone. Watkinson et al. cast doubt on this interpretation based on new offshore multibeam and seismic data by showing that faults that can be traced offshore from the East Arm are dextral, not sinistral, that splays of the Sorong Fault do not exist where they previously

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have been interpreted, and that through-going thrusts shown on many maps are not connected and have different causes. All of these studies of Sulawesi indicate that previous models for tectonic development of this region require substantial re-evaluation.

Banda Arc tectonics The Banda Arc, and especially Timor, is the source of many ideas about arc –continent collision and is also notable for many controversies, such as the origin of the arc, the nature of the crust within the arc and the age of collision. North and west of Timor is the Savu Basin which has a strange triangular shape, widening west towards Sumba which is situated in an anomalous fore-arc position north of the Java Trench where Indian ocean crust is being subducted, and narrowing to the east, north of Timor where arc –continent collision began in the Pliocene. The Savu Basin is situated immediately north of the position of the change from oceanic subduction to arc –continent collision and new seismic data are discussed by Rigg & Hall that help to understand this tectonic transition. The Savu Basin is interpreted to be underlain by Australian continental crust incorporated in the SE Asia margin in the Cretaceous. Seismic sequences offshore can be correlated with stratigraphy onshore and indicate rapid subsidence in the Middle Miocene associated with subduction rollback into the Banda embayment. Subduction of part of the Australian continental margin led to uplift of Sumba and began deformation of former deepwater deposits that are now tilted and slumping northwards into the basin as the former trench became blocked. Audley-Charles discusses the effects that followed the Pliocene collision of the volcanic Banda Arc with the Australian margin as the trench was eliminated. Different parts of the Australian margin sequence were detached at major decollements and stacked up beneath the leading edge of the fore-arc represented by the highest nappes of the Banda allochthon. Contraction in this deformed collision complex caused the distance between the former volcanic arc and the Australian crust to be reduced to as little as 25 km.

The Indonesian Throughflow The Indonesian Throughflow is the last remaining equatorial ocean gateway, allowing heat transfer as water flows from the Pacific into the Indian Ocean. Today, it is regarded as a major component of the modern thermohaline circulation, influencing global climate on short and long timescales (Gordon et al. 2003; Kuhnt et al. 2004). Tillinger describes the causes of the Indonesian Throughflow, the

controls on shallow and deep flow, and its variations in different passages. Fluctuations in the West Pacific Warm Pool are related to variability in the Indonesian Throughflow which acts as a control on inter-annual climate variation such as the El Nin˜o-Southern Oscillation (ENSO) and the SE Asian monsoon. Short term modelling of the effects of restricting the throughflow (e.g. Schneider 1998) suggest that it is likely to affect sea surface temperatures, position of ocean warm pools, land temperatures, rainfall, and wind stresses. The longer term history of the Indonesian Throughflow is of great interest because of the links to global climate but is largely unknown and has been little studied. Holbourn et al. use d18O, d13C and Mg/Ca analyses of benthic and planktonic foraminifera to estimate variations in sea surface temperature, salinity and water mixing over the last 140 ka. The changes are correlated with glacial and interglacial periods and imply links between the Pacific and Indian Oceans via different passages between the oceans as sea level changed, as well as slowing of global thermohaline circulation during glacial intervals. The first restriction of the Indonesian Gateway, and termination of deepwater flow, from the Early Miocene appears to have coincided with major perturbations in the global climate system including rapid warming in the Late Oligocene followed by a brief glaciation pulse and associated significant positive carbon isotopic excursion in the earliest Miocene (Zachos et al. 2001). Climate and geological records suggest that ENSO variability may have existed on Earth as far back as the Eocene. The longterm development of the Indonesian Throughflow has been controlled by the geological history of the region but up to now there have been only a few studies of it (e.g. Kuhnt et al. 2004). However, the effects of other gateways have been modelled and von der Heydt & Dijkstra discuss such studies. They also analyse the effect of increased levels of atmospheric greenhouse gases and open tropical gateways on ENSO variability using fully coupled climate model simulations. Their modelling suggests that greenhouse gas variations have only small effects on ENSO variability but changes in oceanic gateways may cause a stronger and less frequent ENSO. A deeper and more open Indonesian Passage would not prevent a Western Pacific Warm Pool from developing, but could cause the warm pool to move into the Indian Ocean.

Climate history Cane & Molnar (2001) suggested that Pliocene plate tectonic changes, including the northward movement of New Guinea, caused a change from warm saline South Pacific Water to colder North

INTRODUCTION

Pacific water passing through the SE Asian Gateway. They proposed that cooler surface water in the Indian Ocean resulted in increased aridity over eastern Africa. Decreased heat transport out of the tropics may have also stimulated global cooling, resulting in the formation of ice sheets. The climatic changes related to oceanographic and atmospheric changes in the gateway can be assessed using fossils. Morley & Morley provide such an assessment based on palynological studies of cores from petroleum exploration wells in the Makassar Straits, which is now the main passage for the Indonesian Throughflow. Their results provide a record of the vegetation and climate change for the last 30 ka and indicate there were rain forests in Borneo in contrast to extensive grasslands, suggesting a distinctly seasonal climate, in south Sulawesi and the Java Sea during the last glacial maximum. They argue that the equatorial climate has been everwet since the Middle Miocene, but at subequatorial latitudes seasonal climates were established from the Late Pliocene. Lelono & Morley use palynomorph assemblages from marine cores to determine Oligocene climate change in the East Java Sea area. They propose that the Early Oligocene had an everwet climate that favoured rain forest, there was a more seasonal climate in the early part of the Late Oligocene marking reduction in rain forest and increase in grasslands, and a return to rain forest with a superwet climate in the latest Late Oligocene. Some taxa suggest dispersal into Sundaland via the Ninetyeast Ridge in the Oligocene, earlier than previously thought, rather than from Australia during the SE Asia– Australia collision. SE Asia contains the most diverse shallow marine biota on Earth, and a large proportion of this diversity is associated with coral reefs and associated habitats. Carbonates are also a particularly valuable source of environmental and climatic information for the Cenozoic but studies of carbonates of this type have barely begun (Fulthorpe & Schlanger 1989; Wilson 2008). Wilson reviews the information that carbonates are beginning to provide in SE Asia, from annual to million year timescales. Terrestrial runoff, nutrient upwelling, tectonics, volcanism and human activity are major influences on the modern carbonate systems. Quaternary and Pleistocene deposits reveal significant localized tectonic uplift and coeval subsidence, and allow quantification of factors such as interglacial to glacial temperature changes (up to 3–6 8C), ENSO fluctuations (+2 8C extending back at least 130 ka), meltwater pulses associated with ice sheet breakup, and movement of the Intertropical Convergence Zone (ITCZ), where winds of the northern and southern hemispheres meet. During the Cenozoic major changes in oceanography,

5

plate tectonics, climate change and perhaps fluctuating atmospheric CO2 influenced significant changes in carbonate producers and the types of platforms that were constructed. Marine biodiversity reaches a global maximum in the region. Fossils and molecular data suggest that this diversity dates from at least the Early Miocene (Wilson & Rosen 1998; Meyer 2003; Duda & Kohn 2005; Read et al. 2006; Renema et al. 2008) suggesting a possible link to restriction of the SE Asian gateway. These biotic and oceanographic shifts may reflect environmental and tectonicallydriven changes in the distribution and nature of shallow seas.

Summary The collection of papers in this volume written by a variety of Earth and physical scientists reveal significant new data on the processes and timing of large-scale plate tectonic changes in SE Asia, re-evaluate the geological development of specific areas and show the need for significant revisions to previous models. They show the importance of the Indonesian Throughflow and its impact on interannual and longer-term regional and global climate and make a start on unravelling the history of environmental and climatic change of the region from the biota and the rocks. These studies have implications for past distributions of land and sea, terrestrial and marine environments, as well as oceanography and climatology. We hope these contributions to understanding the region’s geological, oceanographic and climatic history will aid the cross-fertilization of ideas with life scientists investigating the enigmatic biology and biodiversity of SE Asia because a symbiotic relationship of life and Earth scientists is essential for a real understanding of this fascinating region. The SAGE meeting was organised jointly by the SE Asia Research Group at Royal Holloway University of London, and the Natural History Museum, London. It was sponsored by the Geological Society of London Tectonic Studies Group and Petroleum Group, the Linnean Society, Malacological Society and Systematics Association. We are grateful to the consortium of oil companies who support the SE Asia Research Group: ENI, ExxonMobil, INPEX, Marathon and Murphy; and SEAPEX (SE Asia Petroleum Exploration Society and the Indonesian Petroleum Association for their help and contributions. We thank Niko Indonesia for their generous support for the meeting and publication.

References Cane, M. & Molnar, P. 2001. Closing of the Indonesian Seaway as a precursor to east African aridification around 3 –4 million years ago. Nature, 411, 157–162.

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Duda, Jr. T. F. & Kohn, A. J. 2005. Species-level phylogeography and evolutionary history of the hyperdiverse marine gastropod genus Conus. Molecular Phylogenetics and Evolution, 34, 257 –272. Fulthorpe, C. S. & Schlanger, S. O. 1989. Paleooceanographic and tectonic settings of early Miocene reefs and associated carbonates of offshore southeast Asia. Bulletin of the American Association of Petroleum Geologists, 73, 729–756. Godfrey, J. S. 1996. The effect of the Indonesian throughflow on ocean circulation and heat exchange with the atmosphere: a review. Journal of Geophysical Research, 101, 12217–12238. Gordon, A. L. 1986. Interocean exchange of thermocline water. Journal of Geophysical Research, 91C, 5037–5046. Gordon, A. L. 2005. Oceanography of the Indonesian seas and their throughflow. Oceanography 18, 14–27. Gordon, A. L., Giulivi, C. F. & Ilahude, A. G. 2003. Deep topographic barriers within the Indonesian seas. Deep-Sea Research, 50, 2205–2228. Gower, D. J., Richardson, J. E., Rosen, B. R., Ru¨ber, L. & Williams, S. T. (eds) 2011. Biotic Evolution and Environmental Change in Southeast Asia. Cambridge University Press, UK, in press. Hall, R. 1998. The plate tectonics of Cenozoic SE Asia and the distribution of land and sea. In: Hall, R. & Holloway, J. D. (eds) Biogeography and Geological Evolution of SE Asia. Backhuys Publishers, Leiden, The Netherlands, 99– 131. Hall, R. 2001. Cenozoic reconstructions of SE Asia and the SW Pacific: changing patterns of land and sea. In: Metcalfe, I., Smith, J. M. B., Morwood, M. & Davidson, I. D. (eds) Faunal and Floral Migrations and Evolution in SE Asia–Australasia. A. A. Balkema (Swets & Zeitlinger Publishers), Lisse, 35–56. Hall, R. 2002. Cenozoic geological and plate tectonic evolution of SE Asia and the SW Pacific: computerbased reconstructions, model and animations. Journal of Asian Earth Sciences, 20, 353– 434. Kuhnt, W., Holbourn, A., Hall, R., Zuvela, M. & Kase, R. 2004. Neogene History of the Indonesian Throughflow. In: Clift, P., Wang, P., Kuhnt, W. & Hayes, D. E. (eds) Continent– Ocean Interactions within East Asian Marginal Seas. American Geophysical Union, Geophysical Monograph Series 149, 299– 320. Lyle, M., Gibbs, S., Moore, T. C. & Rea, D. K. 2007. Late Oligocene initiation of the Antarctic Circumpolar Current: evidence from the South Pacific. Geology, 35, 691–694. Lyle, M., Barron, J. et al. 2008. Pacific Ocean and Cenozoic evolution of climate. Reviews of Geophysics, 46, RG2002, doi: 10.1029/2005RG000190.

Meyer, C. P. 2003. Molecular systematics of cowries (Gastropoda: Cypraeidae) and diversification patterns in the tropics. Biological Journal of the Linnean Society, 79, 401–459. Read, C. I., Bellwood, D. R. & van Herwerden, L. 2006. Ancient origins of Indo-Pacific coral reef fish biodiversity: a case study of the leopard wrasses (Labridae: Macropharyngodon). Molecular Phylogenetics and Evolution, 38, 808–819. Renema, W., Bellwood, D. R. et al. 2008. Hopping hotspots: global shifts in marine biodiversity. Science, 321, 654– 657. Sandwell, D. T. & Smith, W. H. F. 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. Schneider, N. 1998. The Indonesian Throughflow and the global climate system. Journal of Climate, 11, 676–689. Tomascik, T., Mah, A. J., Nontji, A. & Moosa, M. K. 1997. The Ecology of the Indonesian Seas. The Ecology of Indonesia Series, Periplus Editions, Oxford University Press, UK. von der Heydt, A. & Dijkstra, H. A. 2006. Effect of ocean gateways on the global ocean circulation in the Late Oligocene and Early Miocene. Paleoceanography, 21, doi: 10.1029/2005PA001149. Whitmore, T. C. (ed.) 1987. Biogeographical Evolution of the Malay Archipelago, Clarendon Press, Oxford. Whitten, T., Whitten, J., Goettsch, C., Supriatna, J. & Mittermeier, R. A. 1999a. Sundaland. In: Mittermeier, R. A., Gil, P. R. & Goettsch-Mittermeier, C. (eds) Biodiversity Hotspots of the World. Cemex, Prado Norte, Mexico. Whitten, T., Whitten, J., Goettsch, C., Supriatna, J. & Mittermeier, R. A. 1999b. Wallacea. In: Mittermeier, R. A., Gil, P. R. & Goettsch-Mittermeier, C. (eds) Biodiversity Hotspots of the World. Cemex, Prado Norte, Mexico. Wilson, M. E. J. 2008. Global and regional influences on equatorial shallow marine carbonates during the Cenozoic. Palaeogeography, Palaeoclimatology, Palaeoecology, 265, 262–274. Wilson, M. E. J. & Rosen, B. R. 1998. Implications of paucity of corals in the Paleogene of SE Asia: plate tectonics or Centre of Origin? In: Hall, R. & Holloway, J. D. (eds) Biogeography and Geological Evolution of SE Asia. Backhuys Publishers, Leiden, The Netherlands, 165–195. Zachos, J., Pagani, M., Sloan, L., Thomas, E. & Billups, K. 2001. Trends, rhythms, and aberrations in global climate 65 Ma to Present. Science, 292, 686–693.

Palaeozoic– Mesozoic history of SE Asia IAN METCALFE1,2 1

Earth Sciences, Earth Studies Building C02, School of Environmental and Rural Science, University of New England, Armidale NSW 2351, Australia

2

National Key Centre for Geochemical Evolution and Metallogeny of Continents (GEMOC), Department of Earth and Planetary Sciences, Macquarie University, NSW 2109, Australia (e-mail: [email protected]) Abstract: SE Asia comprises a collage of Gondwana-derived continental blocks assembled by the closure of multiple Tethyan and back-arc ocean basins now represented by suture zones. Two major biogeographical boundaries, the Late Palaeozoic Gondwana– Cathaysia divide and the Cenozoic-Recent Australia– Asia divide (Wallace Line) are present. Palaeozoic and Mesozoic evolution involved the rifting and separation of three collages of continental terranes from eastern Gondwana and the opening and closure of three successive ocean basins, the PalaeoTethys (Devonian– Triassic), Meso-Tethys (Permian–Cretaceous) and Ceno-Tethys (Late Triassic–Cenozoic). This led to the opening and closing of ocean gateways and provision of shallow-marine and terrestrial land bridges and stepping-stones for biotic migration. The SE Asia core (Sundaland) comprises a western Sibumasu block, an eastern Indochina–East Malaya block, and the Sukhothai Island Arc terrane between. The Jinghong, Nan-Uttaradit and Sra Kaeo sutures represent the Sukhothai closed back-arc basin. The Palaeo-Tethys is represented by the Changning-Menglian, Chiang Mai/Inthanon and Bentong-Raub suture zones. The West Sumatra and West Burma blocks were accreted to the Sundaland core in the Late Permian– Early Triassic. SW Borneo and/or East Java– West Sulawesi are now identified as the missing ‘Argoland’ that separated from NW Australia in the Jurassic and accreted to SE Sundaland in the Cretaceous.

SE Asia is located at the zone of convergence between the ESE moving Eurasia Plate, the NE moving Indian and Australian Plates and the ENE moving Philippine Plate (Fig. 1). SE Asia and adjoining regions comprise a complex collage of continental blocks, volcanic arcs, and suture zones that represent the closed remnants of ocean basins (including back-arc basins). The continental blocks of the region were derived from the margin of eastern Gondwana as three successive continental strips or collages of continental blocks that separated in the Devonian, Early Permian and Triassic– Jurassic and which then assembled during the Late Palaeozoic to Cenozoic to form present day East and SE Asia (Metcalfe 2005). Global, regional and local Palaeozoic –Mesozoic tectonic evolution resulted in changes to continent– ocean configurations, dramatic changes in relief both on land and in the seas, and changes in palaeo-ocean currents, including the opening and closing of oceanic gateways. The significant effect on ocean circulation caused by ocean gateway closure/ opening is well documented (e.g. Von der Heydt & Dijkstra 2006, 2008). The changes in continent– ocean, land–sea, relief, and ocean current patterns

are fundamental factors leading to both global and regional climate-change and to important changes in biogeographical patterns. Changes in biogeographical barriers and bridges caused by geological evolution and consequent climate-change have also influenced the course of migration, dispersal, isolation and evolution of biota, both globally and in SE Asia. This paper provides an overview of the tectonic framework, and Palaeozoic and Mesozoic geological evolution and palaeogeography of SE Asia and adjacent regions as a background to and to underpin studies of the Indonesian Throughflow Gateway and the distribution and evolution of biota in the region.

Geological and tectonic framework of SE Asia and adjacent regions Mainland East and SE Asia comprises a giant ‘jigsaw puzzle’ of continental blocks, volcanic arc terranes, suture zones (remnants/sites of destroyed ocean basins) and accreted continental crust (Figs 2 & 3).

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 7–35. DOI: 10.1144/SP355.2 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Topography and main active faults in East Asia and location of SE Asia at the zone of convergence of the Eurasian, Philippine and Indian– Australian plates. Large arrows represent absolute (International Terrestrial Reference Frame 2000, Altamimi et al. 2000) motions of plates (After Simons et al. 2007).

Continental blocks of SE Asia The principal continental blocks located in mainland SE Asia (Fig. 2) have been identified and established over the last two decades (e.g. Metcalfe 1984, 1986, 1988, 1990, 1996a, 1998, 2002, 2006) and include the South China block, the Indochina–East Malaya block(s), the Sibumasu block, West Burma block and SW Borneo block (Fig. 3). More recently, the West Sumatra block has been established outboard of Sibumasu in SW Sumatra (Barber & Crow 2003, 2009; Barber et al. 2005) and a volcanic arc terrane is now identified, sandwiched between Sibumasu and Indochina–East Malaya (Sone & Metcalfe 2008). A series of smaller continental blocks are identified in eastern (maritime) SE Asia and these were accreted to the mainland core of SE Asia in the Mesozoic – Cenozoic. The continental terranes of SE Asia and adjacent regions are here categorized into six types based on their specific origins, times of rifting and

separation from Gondwana, and amalgamation/ accretion to form SE Asia. These are discussed below and the suture zones between them are described separately.

Continental blocks derived from Gondwana in the Devonian The South China, Indochina and East Malaya blocks (Figs 2 & 3) are interpreted to have formed part of the India–Australian margin of Gondwana in the Early Palaeozoic and to have rifted and separated from Gondwana by the opening of the PalaeoTethys ocean in the Early Devonian (Metcalfe 1984, 1988, 1990, 1996a, b, 1998, 2002, 2005, 2006). The West Sumatra block (originally proposed by Hutchison 1994; Barber & Crow 2003) and possibly the West Burma block (originally called the ‘Mount Victoria Land Block’ by Mitchell 1986, 1989) are now also interpreted to have originally formed part of this collage of terranes

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

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Tethys 22 Meratus Proto South China Sea 23 Boyan

Fig. 2. Distribution of principal continental blocks, arc terranes and sutures of eastern Asia. WB, West Burma; SWB, SW Borneo block; S, Semitau block; L, Lhasa block; QT, Qiangtang block; QS, Qamdo-Simao block; SI, Simao block; SG, Songpan Ganzi accretionary complex; KL, Kunlun block; QD, Qaidam block; AL, Ala Shan block; LT, Linchang arc terrane; CT, Chanthaburi arc terrane.

(which also included North China and Tarim) that separated from Gondwana in the Devonian (Barber et al. 2005; Metcalfe 2005, 2009a; Barber & Crow 2009). For more detailed description of these blocks and assessment of the evidence for Gondwana origin see Metcalfe (1988, 1996a, 2006). The Late Palaeozoic faunas and floras of these continental blocks are warm-water, equatorial Tethyan/Cathaysian Province biotas that contrast starkly with temporally coeval cold-water and coldclimate Gondwana biotas (Metcalfe 2005). This indicates that these terranes had already separated from Gondwana by Carboniferous times and migrated northwards to more equatorial palaeolatitudes. This is supported by palaeomagnetic data (Zhao et al. 1996; Li & Powell 2001; Li et al. 2004; see Fig. 4). A newly described Early Permian flora (Comia flora) from the West Sumatra block (Booi et al. 2009) indicates relationships with North China and possibly Angara in the Early Permian suggesting a continental migration zone running from the North China Block (via South China and Indochina/East Malaya to the West Sumatra– West Myamar terrane.

Arc terranes derived from South China/ Indochina in the Carboniferous– Permian The Nan-Uttaradit suture, formerly considered to represent the main Palaeo-Tethys ocean, and to mark the boundary between the Sibumasu and Indochina blocks in Thailand, is now regarded as representing a closed back-arc basin (Wu et al. 1995; Ueno 1999; Ueno & Hisada 1999, 2001; Wang et al. 2000). This recognition, and correlation of this suture with the Sra Kaeo suture in southern Thailand and the Jinghong Suture in southern China led Sone & Metcalfe (2008) to propose the Sukhothai Arc System derived from the margin of South China –Indochina– East Malaya by back-arc spreading in the Late Carboniferous –Early Permian. The Sukhothai Arc System is represented by the Lincang block in SW China, the Sukhothai block in central Thailand and the Chanthaburi block in SE Thailand –Cambodia (Fig. 3). The western boundary of the arc is delineated by the Inthanon (Chiang Mai) Suture in NW Thailand, the cryptic Chanthaburi Suture in southern Thailand and the Bentong-Raub Suture in Peninsular Malaysia (Fig. 3; Sone & Metcalfe 2008). The arc is here interpreted to have

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Terranes derived from South China/Indochina in the Cretaceous-Tertiary

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Terranes derived from New Guinea region in the Cenozoic

Terranes derived from Cathaysialand in the Carboniferous-Permian

Extended South ChinaIndochina Continental Crust

Continental crust accreted to Sundaland in the Cretaceous-Cenozoic

Fig. 3. Distribution of continental blocks, fragments and terranes, and principal sutures of SE Asia. Numbered microcontinental blocks: 1, East Java; 2, Bawean; 3, Paternoster; 4, Mangkalihat; 5, West Sulawesi; 6, Semitau; 7, Luconia; 8, Kelabit-Longbowan; 9, Spratly Islands-Dangerous Ground; 10, Reed Bank; 11, North Palawan; 12, Paracel Islands; 13, Macclesfield Bank; 14, East Sulawesi; 15, Bangai-Sula; 16, Buton; 17, Obi-Bacan; 18, Buru-Seram; 19, West Irian Jaya. LT, Lincang Terrane; CT, Chanthaburi Terrane; C-M, Changning-Menglian Suture; C.-Mai – Inthanon, Chiang Mai – Inthanon Suture; Nan-Utt., Nan-Uttaradit Suture.

a thin continental basement that formed the margin of the South China –Indochina–East Malaya superterrane. It separated from Indochina by back-arc spreading in the Early –Middle Permian and was then accreted to Indochina by back-arc collapse in the Late Permian (Fig. 5). Extension of this arc terrane into the Malay Peninsula is equivocal and the previously recognized east Malaya block may form this continuation, but a more likely extension is beneath the Central belt of the Malay Peninsula (Fig. 3) that forms a gravity high (Ryall 1982).

Continental blocks derived from Gondwana in the Early Permian Palaeozoic faunas and floras and the presence of Late Palaeozoic cool/cold water shallow-marine environments, coupled with the presence of Upper Carboniferous –Lower Permian glacial-marine diamictites and palaeomagnetic data suggests that the Sibumasu block, together with the Baoshan, Tengchong, Qiangtang and Lhasa blocks of western China and Tibet remained attached to NE

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

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Fig. 4. Palaeolatitude v. Time for some principal SE Asian continental blocks (After Li et al. 2004). Note northwards migration of South China, Sibumasu and Lhasa from southern to northern latitudes in the Late Silurian– Early Devonian, Permian, and Jurassic– Cretaceous respectively.

Gondwana until Early Permian times. At the end of the Sakmarian stage of the Early Permian the elongate Cimmerian continental strip (Sengo¨r 1984) separated from eastern Gondwana. The eastern portion of this Cimmerian continent includes the Baoshan and possibly the Tengchong blocks of Yunnan, China (Jin 1994; Wopfner 1996), and the Sibumasu block (Metcalfe 1984). These eastern Cimmerian blocks are characterized by Late Palaeozoic Gondwana faunas and floras and by Late Carboniferous–Early Permian glacialmarine diamictites that are interbedded with other marine clastics and turbidites that fill rift grabens (Jin 1994; Wopfner 1996; Wang et al. 2001; Ridd 2009). Metcalfe (1988, 1990) included the Qiangtang and Lhasa blocks as part of the separating eastern Cimmerian continent, but recognized the later docking of the Lhasa block to Eurasia in the Late Jurassic –Early Cretaceous. Metcalfe (1999 & subsequent papers) retained the Lhasa block on the margin of Gondwana until the Late Triassic, a scenario supported by Golonka et al. (2006). Other authors (e.g. Baud et al. 1993; Dercourt et al. 1993) have maintained an Early Permian separation of Lhasa as part of the ‘Mega-Lhasa’ Block. A Triassic –Jurassic separation is still advocated here as proposed and discussed by Metcalfe (1996a). By the end of the Sakmarian, however, the Sibumasu, Baoshan and Tengchong blocks were separating from Gondwana as the eastern part of the Cimmerian continental strip of Sengo¨r (1979, 1984). Post Sakmarian strata of these blocks no longer contain any Gondwana biota but instead show progressive changes in biotic provinces from peri-Gondwana (Indoralian) to Cimmerian to Cathaysian provinces as this continental strip migrated northwards into lower latitudes and warmer climates during the Permian. Recent confusing usage of the term ‘Shan –Thai’ and incorrect correlation with Sibumasu requires

some clarification here. The term Sibumasu (Metcalfe 1984) replaced previous terms used for the elongate Gondwana-derived block in SE Asia characterized by Late Palaeozoic Gondwana biotas and Late Carboniferous –Early Permian glacialmarine diamictites. The acronym SIBUMASU explicitly included ‘SI’ for Sino and Siam, ‘BU’ for Burma, ‘MA’ for Malaya and ‘SU’ for Sumatra where unequivocal Early Permian glacial-marine diamictites are known. ‘Shan–Thai’, ‘Sinoburmalaya’ and ‘West Malaya’ were considered unusable, principally because they did not include the Sumatran element of the block. Recent usage of the term ‘Shan– Thai’ has become so diverse as to become confusing at best and meaningless at worst. Sibumasu and ‘Shan– Thai’ have unfortunately been used interchangeably by many recent authors. In addition, recent interpretations of the Late Palaeozoic Gondwana –Cathaysian biogeographical divide in mainland SE Asia have led to erroneous placements of the eastern margin of Sibumasu and misidentification of the location of the PalaeoTethyan suture zone by some authors. A discussion of these issues is contained in Metcalfe (2009a, b) and will not be repeated here. The tectonic framework for the Sundaland region recently proposed by Ferrari et al. (2008) is not supported here. The use of the term ‘Shan–Thai’ by Ferrari et al. (2008) for a Cathaysian continental block, which in fact includes both continental crustal and suture zone elements, and which bears very little resemblance or relationship to the Gondwanan ‘Shan–Thai’ block of Bunopas (1982) – see Figure 6, is here rejected (see Metcalfe 2009b for details).

Continental blocks derived from Gondwana in the Jurassic Rift-related basin formation and development of unconformities on the NW Australian margin,

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Fig. 5. Cartoon showing the tectonic evolution of Sundaland (Thailand– Malay Peninsula) and evolution of the Sukhothai Arc System during Late Carboniferous – Early Jurassic times (after Ueno & Hisada 1999; Metcalfe 2002; Sone & Metcalfe 2008).

sediment source and palaeocurrent data from Timor, and offshore ocean floor magnetic anomaly data suggest that a piece or pieces of continental crust rifted and separated from Australian Gondwana in the Jurassic. The rifting microcontinents were identified as South Tibet, Burma, Malaya, SW Borneo and Sumatra by Audley-Charles (1988). Veevers et al. (1991) did not identify the continental block that separated from the Argo abyssal plain region in the Jurassic but named this ‘Argo Land’ (subsequently ‘Argoland’). Metcalfe (1990) identified the continental block that must have separated from the Argo abyssal plain as the ‘Mount Victoria Land’ block of Mitchell (1989) located in western Burma. Hard evidence supporting this was lacking because the age and nature of the schist basement of this terrane was not known and no rocks older than Triassic were known. Metcalfe (1996a, b) re-named the block ‘West Burma Block’ to avoid confusion with Mount Victoria

Land in Antarctica. Mitchell (1993) re-interpreted the block as part of an island arc formed by SW directed subduction that was then accreted on to mainland Asia. This interpretation is recently re-proposed by Hall et al. (2009). Other authors have continued to identify ‘Argoland’ as West Burma (e.g. Jablonski & Saitta 2004; Heine & Mu¨ller 2005). The West Burma block is bounded to the east by the Mogok Metamorphic Belt that has recently been correlated with the Medial Sumatra Tectonic Zone (Barber & Crow 2009). Its western boundary is formed by a belt of ophiolites that includes the Mount Victoria metamorphics. The report of Middle Permian rocks from the West Burma block near Karmine with Cathaysian fusulinids similar to those of the West Sumatra block (Oo et al. 2002) suggests that the West Burma block has a Palaeozoic or older continental basement, and may have, together with the West Sumatra block, formed part of a Cathaysian terrane derived from the South

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA 13

Fig. 6. Comparison of the (a) Gondwanan Shan– Thai (ST) block of Bunopas (1982) and (b) Cathaysian ‘Shan–Thai’ block of Ferrari et al. (2008).

14

I. METCALFE

China –Indochina–East Malaya composite terrane and later disrupted by the opening of the Andaman Sea (Barber & Crow 2009). If this interpretation is correct, then the identity of ‘Argoland’ is yet to be established. Hall et al. (2008), Hall (2009) and Hall et al. (2009) have identified ‘Argo’ and ‘Banda’ blocks that separated from the Argo abyssal plain and Banda embayment, NW Australia respectively in the Jurassic. They identify the Argo block as comprising the East Java and West Sulawesi blocks and the Banda block as SW Borneo. A Jurassic Gondwana origin for SW Borneo was previously ruled out on the basis that Cathaysian faunas were known from the Carboniferous– Lower Permian Terbat Limestone on the Sarawak– Kalimantan border (Sanderson 1966; Metcalfe 1985) which were considered part of the SW Borneo basement (Metcalfe 1988). The recognition of a small continental block, the Semitau block, sandwiched between the Lupar and Boyan melanges in west Sarawak (Metcalfe 1990) de-coupled the Cathaysian fusuline-bearing Terbat limestones from the core of the SW Borneo block which then allows SW Borneo to becomes a candidate for the Australian Gondwana derived ‘Argoland’ or ‘Banda’ blocks. This would be supported by the occurrence of diamonds in headless placers (placer diamond deposits without any obvious local or regional diamond source) in Kalimantan (Bergman et al. 1988), SW Borneo (Fig. 7). Nitrogen-defect aggregation studies of these diamonds suggest a Gondwana mantle source (Taylor et al. 1990) consistent with SW Borneo having been derived from NW Australia in the Jurassic. Recent provenance studies (Smyth et al. 2007) have identified an Australian Gondwana-derived East Java terrane. The previously recognized Bawean Arch and Paternoster Platform preCenozoic continental blocks (Manur & Barraclough 1994) are also possibly of Australian Gondwana origin but hard data supporting this is at present lacking. Other small continental blocks postulated to have had their origin on the Mesozoic margin of Australian Gondwana include the West Sulawesi block (which has been linked with the East Java block) and the Mangkalihat block in NE Borneo. It is possible that these micro-continental blocks (numbered 1 –5 on Fig. 3) may in fact represent two disrupted terranes derived from NW Australia (Hall et al. 2009).

Continental blocks derived from South China/ Indochina in the Cretaceous – Cenozoic A number of small micro-continental blocks, the Semitau, Luconia, Kelabit-Longbawen, Spratley Islands-Dangerous Ground, Reed Bank, North Palawan, Paracel Islands and Macclesfield Bank

Fig. 7. Map of mainland SE Asia, showing the distribution of Late Carboniferous– Early Permian glacial-marine sedimentary rocks and major alluvial diamond deposits. Inset photo: Dropstone in glacial marine diamictite oriented vertical to bedding, Singa Formation, Langkawi Islands, Peninsular Malaysia.

(numbered 6–13 on Fig. 3) are interpreted to have originated on the South China –Indochina margin and been translated southwards during NW –SE extension of eastern Sundaland and opening and spreading of the South China Sea. This collage of small blocks may be the disrupted parts of one or two larger terranes. Clift et al. (2008) and Hall et al. (2009) have suggested that these small blocks represent a single large ‘Dangerous Grounds’ terrane that was accreted to Sundaland in the Early Cretaceous and then disrupted by rifting and spreading of the South China Sea. Yan et al. (2010) identify ‘Nasha’ (¼Spratley-Dangerous Ground) and ‘ReedNortheastern Palawan’ blocks in the southern South China Sea region and data from dredged Mesozoic tonalites indicate a probable Precambrian continental basement. The Spratley Islands comprise Cenozoic carbonate reefs constructed on Triassic – Cretaceous sedimentary cuesta basement (Hutchison & Vijayan 2010). Hutchison & Vijayan (2010) do not regard the Spratley-Dangerous Ground region as allochthonous.

Palaeozoic and Mesozoic suture zones of SE Asia The suture zones of SE Asia represent the sites of closed oceanic or back-arc basins and form the

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

boundaries of the continental and arc terranes of the region. The Palaeozoic and Mesozoic suture zones (Figs 2, 3 & 8) comprise the Changning-Menglian,

15

Inthanon, Chanthaburi (cryptic) and Bentong-Raub sutures that represent the destroyed Main PalaeoTethys ocean, the Ailaoshan, Song Ma, Shan

Fig. 8. Tectonic subdivision of Thailand and adjacent regions of Sundaland showing the principal suture zones. Ages of deep marine radiolarian cherts are shown in boxes. C-M S.Z., Changning-Menglian Suture Zone. Modified after Sone & Metcalfe (2008).

I. METCALFE

L

E

Sra Kaeo Suture

L

E

? Unconformity ?

Nan Suture

M

? Unconformity ?

Jinghong Suture (Yunnan)

Semanggol Fm (Successor Basin)

L

Bentong-Raub Suture Zone (Malay Peninsula)

E

Inthanon Suture Zone (Thailand)

TRIASSIC

M Changning-Menglian Suture Zone (Yunnan, China)

The Changning-Menglian suture zone of Yunnan Province, SW China and the Inthanon and Bentong-Raub sutures in Thailand and Peninsular Malaysia are here interpreted to represent the main Palaeo-Tethys Ocean. The Chanthaburi cryptic suture is inferred in southern Thailand and details of this hidden suture are poorly known due to younger cover strata. The Inthanon Suture in Thailand (Figs 2, 3 & 8) corresponds broadly to the Inthanon Zone of Ueno & Hisada (1999) and Ueno (2003) and to the Chiang Mai Suture of Metcalfe (2005) and Wakita & Metcalfe (2005). The suture zone includes radiolarian cherts and deep oceanic sediments that range in age from Middle Devonian to Middle Triassic (Fig. 9). Despite being largely cryptic, Late Devonian, Late Permian and Middle Triassic radiolarian cherts are known from the Chanthaburi Suture in south Thailand (see Sone & Metcalfe 2008 for details; and Fig. 8). Carboniferous–Permian shallow-marine limestones with Cathaysian faunas constructed on intra-oceanic volcanic edifices within the Inthanon Suture are interpreted as Palaeo-Tethyan sea mounts following Metcalfe (2005), Wakita & Metcalfe (2005), Feng et al. (2008) and Ueno et al. (2008). Ocean Plate Stratigraphy (OPS) can be observed in some single outcrop exposures or can be reconstructed from dating of clasts in melange (Wakita & Metcalfe 2005). One such example of OPS with a sequence ranging from pillow basalt up through radiolarian chert, interbedded radiolarian chert and pelagic limestones to deep sea argillites exposed in a single road cutting south of Chiang Mai, Thailand is shown in Figure 10. Recent studies of melange kinematics within the Inthanon Suture, northern Thailand confirm original northwards (present day eastwards) subduction of the Palaeo-Tethys during the Permian–Triassic (Hara et al. 2009).

Sukhothai Back-arc Basin Suture Zone

L

PERMIAN

Changning-Menglian, Inthanon, Chanthaburi and Bentong-Raub (Main Palaeo-Tethys Ocean) sutures

Palaeo-Tethys Suture Zone

CARBONIFEROUS

Boundary and Medial Sumatra sutures that represent other branches of the Palaeo-Tethys, the Jinghong, Nan-Uttaradit and Sra Kaeo sutures that represent the Sukhothai back-arc basin, the Meratus-Lok-Ulo Meso-Tethys suture and the Boyan Proto-South China Sea suture. Other postulated Palaeo-Tethys sutures include the ‘Song Da’ suture (zone) and the ‘Da Nang-Zeijiang’ suture (central Vietnam–South China), Tamky-Phuoc Son suture (South Vietnam), Poko suture (South Vietnam), and Dian-Qiong suture (Yunnan–Hainan). These are briefly discussed below.

DEVONIAN

16

M E

Fig. 9. Ages of oceanic sediments (mainly radiolarian cherts) reported from the main Palaeo-Tethys suture zone segments and from sutures that represent the closed Sukhothai back-arc basin. Grey shaded areas show maximum age ranges for the Palaeo-Tethys ocean and Sukhothai back-arc ocean. E, Early; M, Middle; L, Late. Modified after Sone & Metcalfe (2008).

The Bentong-Raub Suture Zone of the Malay Peninsula includes oceanic radiolarian cherts that range in age from Devonian to Upper Permian (Figs 9 & 11). Triassic cherts of the Semanggol Formation have been interpreted as forming in a successor basin developed on top of the accretionary complex (see Metcalfe 2000 for discussion). A slightly earlier (Early Triassic) closure of PalaeoTethys in the Malay Peninsula compared to a Late Triassic closure in Thailand is indicated.

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

17

Fig. 10. Ocean Plate Stratigraphy (OPS) of the Inthanon Suture Zone exposed in a road cutting south of Chiang Mai, northern Thailand.

Jinghong, Nan-Uttaradit and Sra Kaeo (Sukhothai back-arc) sutures The Jinghong, Nan-Uttaradit and Sra Kaeo Sutures represent the closed back-arc basin that opened in the Permian when the Sukhothai volcanic arc separated from the margin of South China – Indochina– East Malaya. Radiolarian cherts in these sutures are restricted in age from Lower to Upper Permian compared to the age-range for the main PalaeoTethys ocean of Devonian to Triassic (see Fig. 9 and discussion in Sone & Metcalfe 2008).

Ailaoshan, Song Ma and Dian-Qiong (Palaeo-Tethys Branch) sutures A broad NW–SE mobile belt forms the complex boundary zone between the South China and Indochina blocks in north Vietnam and southern China. This complex zone probably also includes microcontinental blocks and accreted volcanic arcs, has been difficult to interpret, and has led to a range of models for the timing and nature of collision between South China and Indochina. The Red River Fault, a major Cenozoic strike-slip fault has often been portrayed as the boundary between South China and Indochina but this is not regarded here as a suture zone (Fig. 8). The amount of displacement along the Red River Fault and fault

reversal history is debated and linked to the India– Asia collision driven Cenozoic extrusion model for the region (e.g. Tapponnier et al. 1982, 1990). A number of narrow belts that contain oceanic rock assemblages (‘ophiolites’, ultramafics, ocean –floor basalts, deep-marine radiolarian cherts) are recognized in north Vietnam –SE China and interpreted as suture zones. The Ailaoshan suture zone has been interpreted as probably representing a back-arc basin between the Simao block and South China (Wang et al. 2000) or an Atlantic type ocean basin (Jian et al. 2009), see Figure 8. Ophiolitic assemblages of the suture include meta-peridotite, gabbro, diabase and basalt capped by radiolarian-bearing siliceous rocks. Plagiogranite (Shuanggou ophiolite) has been dated as early Carboniferous (Jian et al. 1998a, b) and the radiolarian siliceous rocks are Lower Carboniferous in age (Wang et al. 2000; Yumul et al. 2008). The suture was correlated with the Jinshajiang suture to the north and the Nan-Uttaradit suture to the south by Wang et al. (2000). Sone & Metcalfe (2008) however, correlate the Nan-Uttaradit back-arc suture with the Jinghong suture (Fig. 8). The Song Ma Suture zone includes massive serpentinite, altered gabbro and chromitite. The gabbros have a MORB-like affinity and chromianspinels suggest serpentinites represent remnants of the Palaeo-Tethys oceanic lithosphere (Trung

18

I. METCALFE

THAILAND L. Carb. (Tourn)

102E

L. Perm., U. Perm., M. Trias.

5N

t aul

Zo

L. Carb. (Visean)

Raub

N TER EAS LT BE ne

ir F L eb

L. Perm., U. Perm

LT L BE

N

Jengka

U. Dev. L. Perm.

U. Dev. (Fam)

Bentong L. Perm. L. Carb. (Tourn)

L. Carb. (Visean)

Su ture

SIBUMASU TERRANE (PART)

Raub

?L. Carb.

N TER WES LT BE

Cameron Kuala Highlands Kangsar

EAST MALAYA (INDOCHINA) TERRANE (PART)

TRA

CEN

5N

Bentong

M. Trias.

150 km

?

Alor Star

Gunong Semanggol

100

50

0

Malacca

Sheared Diamictite

Bahau U. Dev., L. Carb.

Suture Zone Rocks Muar Semanggol Formation Radiolarian locality with age Sukhothai Arc extension? 102E

?

Fig. 11. Map showing the distribution of the Palaeo-Tethys Bentong-Raub Suture Zone and Semanggol Formation rocks of the Malay Peninsula, ages of radiolarian cherts, and postulated possible extension of the Sukhothai Arc beneath the Central Belt. After Metcalfe (2000).

et al. 2006). Large scale folding, thrust and nappe formation in the Early–Mid Carboniferous, blanketing Middle Carboniferous strata and plant biogeography suggests that the suture is of Early Carboniferous age (Metcalfe 1998). However, geochronological data suggests an Early Triassic event (Lepvrier et al. 2008). It seems unlikely that this event is a collisional one in view of evidence of rifting in Vietnam and South China in the Permo-Triassic. Carter et al. (2001) and Carter & Clift (2008) suggest that there is little evidence to support Indosinian Triassic collision and mountain building in Indochina–South China and that the Early Triassic thermochronology event relates to the accretion of Sibumasu to Indochina. Other authors maintain that the Song Ma Suture provides evidence of continental collision such as granulite-facies metamorphism (Nakano et al. 2008). However, the studied granulites are an allochthonous boulder with no proven source. Deep-marine radiolarian-bearing sediments of Devonian to Triassic age along the southern border of the Nanpanjian Basin in South China have been known for some time and these were regarded as failed-rift sediments related to the Devonian

rifting and separation of South China from Gondwana (Zhao et al. 1996; Metcalfe 1998). Discovery of associated mid-ocean ridge basalts and identification of a central Hainan late Palaeozoic suture zone (Zhou et al. 1999) led Zhang et al. (2006), Zhang & Cai (2009) and Cai & Zhang (2009) to propose the Dian-Qiong suture in South China extending to Hainan Island (Figs 2, 3 & 8). This suture is interpreted to correlate with and be originally contiguous with the Song Ma suture now disrupted by Cenozoic strike-slip telescoping. This suture is interpreted to be of Triassic age and indicates separation of Indochina and South China in the Permian–Early Triassic. This seems to be at variance with biogeographical data and other evidence for an earlier suturing between Indochina and South China along the Song Ma suture zone (see above) and evidence of Permian–Triassic plume-related rifting in the region. Other possible suture zones recognized in the Indochina region include the ‘Song Da suture’ (N. Vietnam) and the ‘Da Nang-Zeijiang’ suture (central Vietnam–South China), Tamky-Phuoc Son suture (South Vietnam), Poko suture (South Vietnam). The Song Da ‘suture’ or zone is here regarded as a Permian–Triassic continental rift zone. The Song Da Rift includes Middle –Late Permian rift related basic volcanic rocks, including komatiite rocks, equivalent to those known in South China (Emenshian volcanic province) and related to mantle–plume activity, and Lower Triassic terrigenous facies sedimentary rocks (see Lepvrier et al. 2008 for discussion). The Da NangZeijiang suture, equivalent to the South Trung Song ophiolite Belt of Hutchison (1989) was proposed by Hoe & Rangin (1999). The nature and age of this ‘suture’ is unclear. Ultramafic and mafic rocks are associated with Cambrian –Silurian rocks and Precambrian blocks but evidence for these representing a subduction zone is unconvincing (Hoe & Rangin 1999). Tamky-Phuoc Son ‘suture’ (Tran 1979; Sengo¨r et al. 1988) includes serpentinized ultramafic rocks in association with mafic and felsic volcanic rocks mixed with continental formations to form a me´lange zone. The zone was regarded as a rift zone by Gatinsky & Hutchison (1987). The Poko ‘suture’ is a major shear zone that incorporates some peridotite, pyroxenite and layered gabbro associated with rhyolitic and andesitic rocks. This suture was regarded as a continental rift by Ferrari et al. (2008) relating to their ‘Orang Laut terranes’.

Shan Boundary and Medial Sumatra Tectonic Zones (Palaeo-Tethyan ‘sutures’) The Cathaysian West Sumatra and West Burma blocks are now positioned outboard (current west)

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

of the Gondwanan Sibumasu block. The mechanism of emplacement to their present relative locations to other continental blocks of the region is interpreted as strike-slip tectonics (Barber & Crow 2003; Wakita & Metcalfe 2005, 2009a; Barber & Crow 2009). The boundary between the Sibumasu block and the SW Sumatra block is the Medial Sumatra Tectonic Zone (Barber & Crow 2003) that represents a major transcurrent shear zone. There is no evidence to date of the remnants (ocean floor stratigraphy, melange, ophiolites) of the intervening branch of Palaeo-Tethys that must have existed. This zone appears to correlate with the Mogok Metamorphic Belt in Burma that forms the boundary between Sibumasu and West Burma and which is also interpreted as a major transcurrent shear zone (Barber & Crow 2009).

Meratus-Lok-Ulo Meso-Tethys suture The Jurassic –Cretaceous SW Borneo Meratus and central Java Lok-Ulo sutures represent the destroyed Meso-Tethys ocean that separated the East Java, Bawean and Paternoster blocks from SW Borneo/Sundaland. The Meratus suture complex comprises melange, siliceous shale, limestone, basalt, ultramafic rocks and schist. Radiolarian cherts range in age from Middle Jurassic to Early Cretaceous (Wakita et al. 1997, 1998). The LukUlo suture complex comprises similar lithologies. Reconstructed ocean plate stratigraphies represent the entire Cretaceous and include sea mount rock associations (Wakita & Metcalfe 2005).

Boyan Proto-South China Sea suture The Boyan suture is located between the small Semitau block and SW Borneo. Melange in the suture is of Late Cretaceous age formed by destruction of the Proto-South China Sea (Metcalfe 1999).

Palaeozoic – Mesozoic evolution and palaeogeography of SE Asia The Gondwana origins of all component continental blocks of SE Asia is now widely accepted. These continental blocks rifted and separated from NE Gondwana as three continental slivers or collages of terranes in the Early –Middle Devonian, Early Permian, and Late Triassic –Jurassic. Successive Tethyan ocean basins, the Palaeo-Tethys, MesoTethys and Ceno-Tethys opened between each separating sliver and Gondwana. The separated continental blocks migrated successively northwards to in some cases amalgamate, and then accrete to form the core of East and SE Eurasia. The timings

19

of rifting and separation of these continental blocks and ages of amalgamation and accretion in relation to the three successive Tethyan ocean basins are shown in Figure 12.

Early – Middle Palaeozoic evolution and palaeogeography Tectonostratigraphical, biogeographical, geochemical, provenance study, and palaeomagnetic data indicate that all the principal continental blocks of East and SE Asia formed a greater Indian– Australian Gondwana margin in the Early Palaeozoic (Fig. 13a). Metcalfe (1988, 1990, 1996a, 1999, 2006) has presented the evidence for such placement and this will not be repeated here. Cambro-Ordovician faunas on the North China, South China, Tarim, Indochina, Sibumasu, Qiangtang, and Lhasa blocks define an Asia –Australia province at this time (Fig. 14) and palaeomagnetic data is consistent with their placement on the India–Australian Gondwana margin. Other workers have invoked a similar reconstruction scenario (e.g. Fortey & Cocks 1998; Golonka 2000, Golonka et al. 2006; Hendricks et al. 2008). By Mid –Late Silurian times, Gondwana had rotated clockwise significantly but NE Gondwana remained in low northern palaeolatitudes (Fig. 13b). The Sundaland/Asian terranes remained in their previous relative positions, continuing to form a greater Gondwana margin. Again, biogeographical data indicates an Asia –Australian province particularly well illustrated by the distribution of the Retziella brachiopod fauna (Fig. 13b). In the Late Silurian, a rifting event occurred on the margin of Gondwana and an elongate continental sliver comprising the South China, Tarim, Indochina and North China blocks began to separate from Gondwana in the Early Devonian (Metcalfe 1996a, b). By Late –Early to Middle Devonian times, oceanic spreading between this continental sliver and Gondwana opened the Palaeo-Tethys ocean basin as evidenced by oceanic deep sea radiolarian cherts in the suture zone. By latest Devonian –earliest Carboniferous times the separating sliver had almost broken away from Gondwana but retained continental connection in the east explaining continued Devonian fish faunal connections (Metcalfe 2001). Clockwise rotation of the sliver away from Gondwana corresponds to documented anti-clockwise rotation of Gondwana in the Late Devonian (Metcalfe 2001). Interestingly, the distribution of the Chuiella brachiopod fauna (Chen & Shi 1999) in the shallow seas of South China and Tarim on the western extremity of the continental sliver is consistent with this scenario (Fig. 15).

20

I. METCALFE COLLISION (AMALGAMATION & ACCRETION)

Si b

M er i a on , K go az lia ak h n A st rcs an

Early Carboniferous (S. China-Indochina) Early Permian (Tarim-Eurasia) Late Triassic (S. China-N. China) Jurassic (N. China-Eurasia)

ific

ac

P o-

e

a , hin doc , Tarim , In a ina matr h Su hC ort est & N ya, W s uth So t Mala ma ethy eo-T Eas t Bur s Pala We

Eurasia SEPARATION Middle Devonian

la Pa

Lower to Upper Triassic (Sibumasu-Indochina/E. Malaya) Lower Triassic (W. Sumatra/W. Burma-Sibumasu) Middle-Late Triassic (Sibumasu/Qiangtang-Indochina/Simao)

Cimmerian Continent (Sibumasu, Qiangtang)

Early Permian

Latest Jurassic (Lhasa-Eurasia) Meso

Lha Ba sa, A w r We ean, golan st S Pat d ula erno (East we si, ster, S Java, Ma ngk W Bo alih rne o?, at)

-Tethy

Late Triassic Late Jurassic

s

C

Cretaceous (Argoland/SW Borneo-Sibumasu/ E. Malaya/W. Sumatra) Cretaceous (Woyla-Mawgyi Arc-W. Sumatra)

en

I nd

Gondwanaland

o-

Te

ia

ys

Paleocene (India-Incertus Ocean Island Arc) Eocene/Oligocene (India-Eurasia)

n

cea

an O

Indi

th

Fig. 12. Schematic diagram showing times of separation and subsequent collision of the three continental slivers/ collages of terranes that rifted from Gondwana and translated northwards by the opening and closing of three successive oceans, the Palaeo-Tethys, Meso-Tethys and Ceno-Tethys.

Late Palaeozoic evolution and palaeogeography In Late –Early Carboniferous (Visean) times the faunas and floras of North China, South China and Indochina–East Malaya no longer exhibited any Gondwana affinities and these blocks were located in equatorial to low northern palaeolatitudes (Fig. 16). South China and Indochina had probably amalgamated along the Song Ma suture zone. Late –Early Carboniferous floras of South China and Indochina–East Malaya are very similar suggesting continental connection between these blocks at this time (Laveine et al. 1999). The Visean biogeographical distribution of the shallowmarine conodont genus Mestognathus indicates that Laurentia and Gondwana were connected but isolated from other continental terranes and the distinctive shallow-marine conodont genus Montognathus links the Sibumasu block with eastern Australia at this time (Fig. 16). The Sibumasu block and the Argo/SW Borneo blocks remained on the NW Australian margin of Gondwana. The Tarim block, following its separation from Gondwana in the Devonian, collided with Siberia

in the Late Carboniferous to Early Permian and was welded to proto-Asia by the Middle Permian (Carroll et al. 1995). Gondwanan glaciation was at its maximum development in the Latest Carboniferous–Earliest Permian and ice sheets covered large parts of the super continent, including Australia. Ice rafted onto the shallow-marine continental shelf of Australian Gondwana and dumped glacial debris into marine sediments resulting in the glacial-marine diamictite bearing deposits on the Sibumasu block (Fig. 7). The Sibumasu block was already at this time in the process of rifting from Gondwana and as a result, glacial-marine strata filled rift grabens (Fig. 5). The Early Permian was also a time of high provinciality of global floras and faunas and the Sibumasu block floras were typical Gondwanan Glossopteris floras at this time. Early Permian floras of the North China, South China and Indochina– East Malaya blocks (located in isolated intra-Tethyan positions) are typical Cathaysian warm-climate Gigantopteris floras (Fig. 17). Conodont faunal provinciality was also marked in the early Permian with a distinct southern hemisphere high-latitude periGondwana cool-water province characterized by

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

21

(a)

CAMBRO-ORDOVICIAN (TREMADOC - 485 Ma)

30N AUSTRALIA

NC

PALAEO-PACIFIC LAURENTIA SIBERIA

S B SW

KAZAKHSTAN

D

LAN

NA WA

BALTICA

T

D ON

G

I

0

Aporthophyla tianjingshanensis Brachyhipposiderus spp. Peelerophon oehlerti Spanodonta spp. Serratognathus Tasmanognathus Songtaoia spp. Land Aurilobodus Shallow Sea Asaphopsoides Deep Sea Koraipsis

0 QI

L

GI

ANTARCTICA

SC INDIA 30S

30S AFRICA

(b)

MID-LATE SILURIAN (420 Ma)

30N

PALAEO-PACIFIC KAZAKHSTAN

Shallow clastics

TARIM ND ALA

HIN

A

N WA

BRACHIOPODS

IND

G

N TA

Retziella Fauna

G

N

IA

Q

Subduction Zone Land

U

B

SIB

Palaeo Equator

Marginal marine

NIC ELF H CA OL D/S F V SLAN I

Nericodus and Tuberocostadontus

EL

SOUTH CHINA

CONODONTS

SH

OC

ND

GO

UM

LAURENTIA

NEW GUINEA

SW

BALTICA

NORTH CHINA

AS

SIBERIA

AUSTRALIA

A

AS

LH

WC

GREATER INDIA

Shallow Sea

INDIA

ANTARCTICA

Deep Sea

Fig. 13. Reconstructions of eastern Gondwanaland for (a) Cambro–Ordovician (Tremadoc) and (b) Mid - Late Silurian showing the postulated positions of the East and SE Asian terranes, distribution of land and sea, and shallow-marine fossils that illustrate Asia– Australia connections at these times. NC, North China; SC, South China; T, Tarim; I, Indochina/East Malaya/West Sumatra; QI, Qiangtang; L, Lhasa; S, Sibumasu; WC, Western Cimmerian Continent; GI, Greater India; SWB, Argoland/SW Borneo.

Indochina

South China

North China

Tarim

SW Borneo

22

Sibumasu 65 L

CRETACEOUS

Europe, China Yunnan, Kwangsi

M

Laurasia

Laurasia, Tibet Ryoseki Type

Tethyan

E

145 L

East Asian, Japan, Philippines

Ryoseki Type Laurasia

JURASSIC

M

Laurasia Tethyan

Tethyan

Tethyan

?

E

202 L

TRIASSIC 252.5

Eastern Tethyan

M

M

299

Eastern Tethyan

Pangea

South China, Indochina Sibumasu Province N. W. Australia Gondwanaland

Tethyan

Laurasia

?

Tethyan Eastern Tethys

Tethyan

Tethyan

Cathaysian

Angaran

Cathaysian

Cathaysian

L Arctic-Eurasian

E

Eastern Australia Arctic-Eurasian N. W. Australia

South China China China

Palaeo-Tethyan

Palaeo-Tethyan

L

Tarim

Eastern Gondwanaland

M

Eastern Gondwanaland S. China, Eastern Australia

E

Peri-Gondwanaland

South China

Eastern Australia

416

SILURIAN 444

PRI LUD WEN

LLY

488

Gastropods Bivalves

Nautiloids

Small forams

Australia

Indochina, East Gondwana

S. China

Sino-Australian Province

Trilobites Conodonts

Brachiopods Sino-Australian Province

Sino-Australian Province

South China Sino-Australian Province

Fusulines

Plants Vertebrates Corals

Stromatoporoids

L

ORDOVICIAN

Sino-Australian Province

East Gondwana Sibumasu, E. Australia South China, East Gondwana

?

Palaeo-Tethyan

Indochina

359

DEVONIAN

?

Cathaysian

Tethyan

I. METCALFE

E

CARBONIFEROUS

Tethyan Eastern Tethys

E L

PERMIAN

Japan Yunnan Laurasia

M

S. China (Pagoda Fm) S. China Australia, Tibet, N. China

E

S. China, Argentina

L M

N. W. Australia

Sibumasu Sibumasu

Terrestrial Vertebrates Sino-Australian Province

Australia, Tibet, N. China Sibumasu, Argentina

Major biotic provinces Sino-Australian Province Gondwana

N. W. Australia

Sino-Australian Province

Transitional

Cathaysian/Tethyan

CAMBRIAN E

542

Redlichiid Asia-Australian Realm N China, Australia

Redlichiid Asia-Australian Realm S China, Australia

Redlichiid Asia-Australian Realm

Fig. 14. Palaeozoic and Mesozoic faunal and floral provinces and affinities v. time for the principal East Asian continental blocks (After Metcalfe 2001).

Angaran

Laurasia

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

SIBERIA

KAZAKHSTAN

30N LAURENTIA

AND

NAL

WA ND GO

LATE DEVONIAN EARLY CARBONIFEROUS (TOURNAISIAN)

NC

T Palaeo Equator

I

SC

YS

PA L

A

E TH E O-T

SWB

QI WC

L

30S

Shallow Sea Deep Sea

S

AUSTRALIA

Chuiella Subduction Zone Land

23

AFRICA

INDIA

ANTARCTICA

Fig. 15. Reconstruction of eastern Gondwana in Late Devonian to Lower Carboniferous (Tournaisian) times showing the postulated positions of the East and SE Asian terranes. Also shown is the distribution of the endemic Tournaisian brachiopod genus Chuiella. NC, North China; SC, South China; T, Tarim; I, Indochina/East Malaya/West Sumatra/West Burma; QI, Qiangtang; L, Lhasa; S, Sibumasu; SWB, Argoland/SW Borneo; WC, Western Cimmerian Continent.

the genus Vjalovognathus, an equatorial warmwater Sweetognathus-dominated province and a northern hemisphere high-latitude cool-water Neostreptognathodus-dominated province (Fig. 18). Continental connection or close proximity of South China and Indochina in the Kungurian is indicated by the endemic occurrence of Pseudosweetognathus on these two blocks (Metcalfe & Sone 2008; Fig. 18). In the Asselian– Sakmarian, Sibumasu block faunas were peri-Gondwanan Indoralian Province faunas, but as Sibumasu separated and moved northwards during the Permian its faunal characteristics changed, first to endemic Sibumasu province faunas in the Middle Permian and then to Cathaysian Province faunas in the Late Permian (Shi & Archbold 1998; Ueno 2003). As Sibumasu was translated northwards during the Permian, the Palaeo-Tethys was subducted beneath northern Pangaea, North China and the amalgamated South China –Indochina–East Malaya terrane (Cathaysialand). Subduction beneath Cathaysialand resulted in the Sukhothai Arc on its margin which was then separated from Cathaysialand by back-arc spreading

to become an Island Arc in the Late Permian (Fig. 18). The resulting narrow back-arc basin collapsed at the end of the Permian to form the Jinghong, Nan-Uttaradit and Sra Kaeo sutures (Sone & Metcalfe 2008). Collision of the Sibumasu block with the Sukhothai Island Arc terranes and Cathaysialand closed the southeastern PalaeoTethys in the Late Permian–Early Triassic producing the Changning-Menglian, Inthanon and Bentong-Raub suture zones. A later timing (Late Triassic or even Jurassic) for this collision has been suggested by some authors based on interpretation of the Semanggol cherts and equivalents as Palaeo-Tethyan deposits (e.g. Sashida et al. 1995, 2000a, b; Kamata et al. 2002; Ueno et al. 2006; Ishida et al. 2006; Hirsch et al. 2006). The earlier timing is here supported following Metcalfe (2000) and Barber & Crow (2009). A younger (late Triassic) collision and suturing to the north along the Changning-Menglian suture in SW China is however considered possible (Liu et al. 1996). It is postulated here that the West Burma and West Sumatra blocks (initially as a single block) were translated westwards by transcurrent tectonics

24

I. METCALFE

EARLY CARBONIFEROUS - VISEAN (340 Ma) SIBERIA KAZAKHSTAN LAURENTIA

AND

L ANA DW

N

GO

Montognathus

Mestognathus KAZ North China

30

Tarim

LAURENTIA

South China

PA L

Indochina/East Malaya

AE

O-

0

TE

QS

TH

YS

WC

S QI L

AUSTRALIA Subduction Zone Land

30

AFRICA

INDIA

Shallow Sea Deep Sea

Fig. 16. Early Carboniferous (Visean) reconstruction showing postulated positions of East and SE Asian blocks. The biogeographic distributions of the conodont genera Mestognathus (Illustrated specimen is Mestognathus beckmanni from the Kanthan Limestone, Peninsular Malaysia) and Montognathus (Montognathus carinatus from Peninsular Malaysia illustrated) are also shown.

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

25

KAZAKSTAN NORTHEAST CHINA (COMPOSITE)

FLORAL PROVINCES

TARIM KL QT L INDIA

AL

QD QS

NORTH CHINA

SG

KT

Angaran

SOUTH CHINA WB

Cathaysian

SI I ND H OC I NA

MASU

SIBU

Gondwanan LOWER PERMIAN FLORAS American

T EAS YA LA MA

S

Euramerican SWB

(a)

WEST SUMATRA

30 T NORTH CHINA CATHAYSIALAND

L PA

WS

O-

WB INDOCHINA

TE YS

TH

NG

WC

EA

(b)

0

AE

PA EARLY PERMIAN (290 Ma)

SOUTH CHINA

QI

30 S SWB

L INDIA

AUSTRALIA

60

Fig. 17. Distribution of Lower Permian floral provinces plotted on (a) present-day geographical map, and (b) Early Permian palaeogeographic map. KT, Kurosegawa Terrane. Other abbreviations as for Figures 2 and 3.

to their current positions outboard of the Sibumasu terrane in the Early Triassic. This translation was coeval with and partly the result of the collision of Sibumasu and Cathaysialand, which occurred at the zone of convergence between the north moving Meso-Tethys and west moving Palaeo-Pacific plates. The South and North China blocks were in close proximity during the Permian. The timing of their collision and welding is an ongoing controversy with Mid-Palaeozoic, Late Palaeozoic and Late Triassic –Jurassic timings being proposed. Studies of low grade metamorphics in the Sulu belt (Zhou et al. 2008) and geochronological and structural data (e.g. Faure et al. 2003) indicate Permian

subduction of South China beneath North China. Identification of a Devonian –Triassic accretionary wedge that includes eclogites, and which formed a coeval volcano-plutonic arc that stretches from the Longmen Shan to Korea supports subduction beneath the Qinling –Sino-Korean plate and a Permian –Triassic collision (Hacker et al. 2004). A land connection between Indochina and Pangaea in the Late Permian is indicated by the confirmed presence of the Late Permian tetrapod vertebrate Dicynodon in Laos (Battail 2009). The most likely land connection was via South and North China rather than via the western Cimmerian continental strip that was largely submerged below sea level in the Permian (Fig. 18).

26

Fig. 18.

I. METCALFE

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

27

60

NC SG

EM WS WB

SO

-TE

TH

NG EA

Land

30

I S

WC QI ME

PA

Late Triassic (Rhaetian)

SC

Shallow Sea

L

YS

0

SWB 30

Deep Sea

Fig. 19. Palaeogeographic reconstructions of the Tethyan region for the Late Triassic (Rhaetian) showing relative positions of the East and SE Asian terranes and distribution of land and sea. NC, North China; SG, Songpan Ganzi; SC, South China; WC, Western Cimmerian Continent; QI, Qiangtang block; I, Indochina block; S, Sibumasu block; EM, East Malaya block; WS, West Sumatra block; WB, West Burma block; L, Lhasa block; SWB, Argoland/SW Borneo.

Mesozoic evolution and palaeogeography Collision and welding of the Sibumasu block to Indochina–East Malaya, begun in the latest Permian, continued in the Early–Middle Triassic and was completed by Late Triassic times (Fig. 19). Collision between South and North China began in the Permian and continued in the Triassic. Comparisons of apparent polar wander paths (APWPs) of these blocks indicates that collision between these blocks also continued into the Jurassic but was complete by the Late Jurassic. The time of rapid (18/Ma) relative angular velocity between the two plates (225 to 190 Ma) coincides with a peak in U–Pb and Ar–Ar dates obtained from metamorphic rocks in the Qingling-Dabie-Sulu suture (Gilder & Courtillot 1997). Thus, the initial consolidation of what is now Sundaland and mainland East and SE Asia took place in Late Triassic–Jurassic times. The Songpan Ganzi giant

suture knot represents Palaeo-Tethyan ocean crust trapped between the western Cimmerian continent, Cathaysialand, North China and Siberian Pangaea and covered by thick Triassic deposits eroded from adjacent collisional orogens. A third episode of rifting of the Indian– Australian margin of Gondwana was initiated in the Triassic and continued into the Jurassic/ Cretaceous (Fig. 20). The Lhasa block is here interpreted to have separated from Indian Gondwana in the Late Triassic (following Metcalfe 2002; Golonka et al. 2006; Golonka 2007) but other authors have advocated an earlier separation as part of the Cimmerian continent (e.g. Alle`gre et al. 1984; Dercourt et al. 1993, 2000). A Permian separation of Lhasa may be supported by Permian limestone blocks interpreted as possible seamount caps in the Indus-Yarlung suture zone (Shen et al. 2003) but this would require the unlikely

Fig. 18. Palaeogeographic reconstructions of the Tethyan region for (a) Early Early Permian (Asselian –Sakmarian), (b) Late Early Permian (Kungurian) and (c) Late Permian (Changhsingian) showing relative positions of the East and SE Asian terranes and distribution of land and sea. Also shown are the Early Permian occurrences of the Comia flora linking West Sumatra, North China and Pangaea; Asselian– Sakmarian ice sheet and peri-Gondwana glacial-marine diamictite localities; Late Early Permian biogeographical provinces and distribution of biogeographically important conodonts and the bipolar fusulinid Monodiexodina; and Late Permian tetrapod vertebrate Dicynodon localities on Indochina and Pangaea in the Late Permian. SC, South China; T, Tarim; I, Indochina; EM, East Malaya; WS, West Sumatra; NC, North China; SI, Simao; S, Sibumasu; WB, West Burma; QI, Qiangtang; L, Lhasa; SWB, SW Borneo; WC, Western Cimmerian Continent.

28

I. METCALFE

QI

QS SG

SC

Kohistan Arc

SI

QI

NP

Land

WSu

EM

S Lhasa

Deep Sea

0 Mawgyi Arc

Woyla Arc

WB

NP I

S

Shallow Sea

I

SA

SC SI

L

EM

Mawgyi Arc

WSu WB

0

Woyla Arc

MESO-TETHYS

MESO-TETHYS

SW Borneo

In

ce

rtu

Argoland

sI

SWneo r Bo

sla

nd

Argoland

30

PACIFIC OCEAN

SA

Ar c

CENO-TETHYS

30 GREATER INDIA

M16

M21

M16

M 7

N. GUINEA

INDIA

M21

M21 M21

Timor

Tanimbar

GREATER INDIA

M7

AUSTRALIA AUSTRALIA

60

60 ANTARCTICA

a

b

LATE JURASSIC (165 Ma)

EARLY CRETACEOUS (120 Ma)

L

L

Inc

SI

us

I

WB

Isla

nd

PS

I

EM

SWB

Lu

NP

M

WSu

P

0

EM SE

WB

WS

WSu

EJ B

GREATER INDIA

M Z PHILIPPINE SEA PLATE

WS EJ SWB P B 33

NINETY

CENO-TETHYS

INDIA

PA

NP PACIFIC OCEAN PLATE

Da

S

Arc

0

RB MB

PI

EAST

ert

30

PACIFIC OCEAN

SA

Inc

SC

Co ll ert ision us ? Arc

SC S

ES

Ba-Su O WIJ Bu B-S

N. GUINEA

30

N. GUINEA INDIA

33

INDIAN OCEAN

AUSTRALIA

0

33 M

INDIAN OCEAN

M

0

AUSTRALIA

60

60 ANTARCTICA

ANTARCTICA

c

LATE CRETACEOUS (80 Ma)

d

MIDDLE EOCENE (45 Ma)

Fig. 20. Palaeogeographic reconstructions for Eastern Tethys in (a) Late Jurassic, (b) Early Cretaceous (c) Late Cretaceous and (d) Middle Eocene showing distribution of continental blocks and fragments of SE Asia–Australasia and land and sea. SG, Songpan Ganzi accretionary complex; SC, South China; QS, Qamdo-Simao; SI, Simao; QI, Qiangtang; S, Sibumasu; SA, Sukhothai Arc; I, Indochina; EM, East Malaya; WSu, West Sumatra; L, Lhasa; WB, West Burma; SWB, SW Borneo; NP, North Palawan and other small continental fragments now forming part of the Philippines basement; M, Mangkalihat; WS, West Sulawesi; P, Paternoster; B, Bawean; PA, Incipient East Philippine arc; PS, Proto-South China Sea; Z, Zambales Ophiolite; ES, East Sulawesi; O, Obi-Bacan; Ba-Su, Banggai-Sula; Bu, Buton; WIJ, West Irian Jaya. M numbers represent Indian Ocean magnetic anomalies.

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

longitudinal splitting of the Cimmerian continent during its northwards movement and the opening of a new ocean basin between Lhasa and Qiangtang. A possible slab pull mechanism has been advocated by Stampfli & Borel (2002) but is here considered unlikely. A Late Triassic separation advocated here is supported by information on oceanic cherts from the Yarlung-Zangbo suture (Matsuoka et al. 2002) and recent palaeomagnetic data (Otofuji et al. 2007). A collage of small continental blocks then rifted and separated progressively westwards from the NW Australian margin in the Late Jurassic –Early

(d)

29

Cretaceous (Fig. 20). These included the Argoland block that separated by opening of the Argo Abyssal Plain and SW Borneo (referred to as the ‘Banda’ block by Hall et al. 2009) from the Banda Embayment region. These were previously identified as West Burma, and other small continental blocks in the Sumatra and Borneo region (Metcalfe 1990; Jablonski & Saitta 2004; Heine & Mu¨ller 2005). Argoland is now tentatively identified as the East Java, Bawean, Paternoster, Mangkalihat, and West Sulawesi blocks (numbered 2 –5 on Fig. 3) and the Banda block as SW Borneo, following Hall et al. (2009).

(h) SIBERIA KAZAKHSTAN

Late Triassic (Rhaetian)

LAURENTIA

PA

Early Carboniferous (Visean)

NG EA

D

LAN

NA WA

ND

GO

(c)

(g) SIBERIA

Late Permian (Changhsingian)

LAURENTIA

PA NG

Late Devonian - Early Carboniferous

KAZAKHSTAN

EA

D

LAN

ANA DW

N GO

(f)

(b) Late Silurian

SIBERIA KAZAKHSTAN

Early Permian (Kungurian) PA

BALTICA

NG

LAURENTIA

EA

ND

ALA

N WA

ND

GO

(a)

(e) Cambro-Ordovician (Tremadoc) Early Permian (Asselian-Sakmarian) PA

KAZAKHSTAN

SIBERIA

NG

LAURENTIA

EA

ND

ALA

AN DW

BALTICA

N GO

Continent

Ocean

Open Ocean Gateway

Closed Ocean Gateway

Fig. 21. Changing continent –ocean configurations in the Palaeozoic– Early Mesozoic and evolving ocean gateways.

30

I. METCALFE

SW Borneo and Argoland were translated northwards during the Cretaceous and by Late Cretaceous times had accreted to SE Sundaland. The Incertus Island Arc developed within the Ceno-Tethys during the Cretaceous (Aitchison et al. 2007; Hall et al. 2009) and collided with northwards moving India at c. 55 Ma. By Middle Eocene times (45 Ma), India (with accreted Incertus Arc segment) was probably in its initial collision with Eurasia (Fig. 20). The 45 Ma timing is temporally coincident with large-scale regional and global plate reorganizations at this time (Hall et al. 2009). A younger ‘hard’ collision between India and Eurasia at c. 35 Ma has however been recently proposed by Ali & Aitchison (2007, 2008) but challenged by Yin (2010) who maintains an early c. 60 Ma initial collision.

Palaeozoic – Mesozoic ocean gateway evolution Changes in global continent–ocean configurations during the Palaeozoic –Mesozoic (Fig. 21) have led to both opening and closure of oceanic gateways that undoubtedly had significant effects on both global and SE Asian ocean currents, circulation and upwelling, and climate. This in turn led to

changing biogeographical patterns and biotic provinces. In the early to Middle Palaeozoic, Gondwana was separated from other dispersed major continental blocks with significant ocean gateways between Gondwana and between Laurentia, Baltica, Siberia and Kazakhstan (Fig. 21a–c). Changes in the positions of blocks during the Cambrian and Devonian led to closure of the gateway between Laurentia and Siberia in the Silurian but a major ocean gateway between Gondwana and other blocks was maintained. In the Devonian, North China, South China, Tarim and Indochina/ East Malaya/West Sumatra/West Burma rifted and separated from NE Gondwana opening the Palaeo-Tethys. This eventually led to the opening of a gateway between Gondwana and SE Asian blocks in the Early Carboniferous, and with clockwise rotation of Gondwana, the closure of the major gateway and separating ocean between Laurentia and north African west Gondwana initiated the final formation of Pangaea. Northwards migration of SE Asian continental blocks in the Permian– Triassic resulted in the opening and closure of oceanic gateways between the Panthalassa in the east and the largely closed Tethys in the west (Fig. 21e–h). The changing gateways in the Tethyan region during the Permian – Triassic must have resulted in changing ocean

Spathian PANTHALASSA

60 40 20 TETHYS

0

?

–20 –40

50

100

150

200

250

300

PANGEA

350 PANGEA

Smithian PANTHALASSA

60

?

40 20 TETHYS

0 –20 –40

50 PANGEA

100

150

200

250

pantropic dispersion

300

350 PANGEA

weak dispersion intense dispersion

Fig. 22. Ammonoid dispersion for the Olenekian (Smithian and Spathian). After Brayard et al. (2009).

PALAEOZOIC–MESOZOIC HISTORY OF SE ASIA

currents and regions of deep ocean upwelling affecting climate patterns and dispersal of biota (e.g. ammonite dispersal of Brayard et al. 2009; Fig. 22).

31

but in particular with Robert Hall, Anthony Barber, Mike Crow and Masatoshi Sone. Anthony Barber and Mike Crow are also thanked for their helpful reviews of the paper.

Conclusions The Palaeozoic–Mesozoic evolution of SE Asia involved the rifting and separation of three collages of continental terranes from eastern Gondwana and the successive opening and closure of three ocean basins, the Palaeo-Tethys, Meso-Tethys and Ceno-Tethys. The Palaeo-Tethys is represented in SE Asia by the Inthanon (Chiang Mai), Chanthaburi (cryptic) and Bentong-Raub suture zones. The Sukhothai Island Arc System, including the Linchang, Sukhothai and Chanthaburi terranes is identified between the Sibumasu and Indochina– East Malaya terranes in mainland SE Asia. It was constructed on the margin of Indochina–East Malaya and separated by back-arc spreading in the Permian. The Jinghong, Nan-Uttaradit and Sra Kaeo sutures represent the closed back-arc ocean basin. The West Sumatra and West Burma blocks rifted and separated from Gondwana, along with Indochina and East Malaya in the Devonian and formed a composite terrane ‘Cathaysialand’ with South China in the Permian. In the Late Permian–Early Triassic, West Sumatra and West Burma were translated westwards to their positions outboard of Sibumasu by strike-slip translation at the zone of convergence between the Meso-Tethys and Palaeo-Pacific plates. The continental micro-blocks that rifted and separated from Gondwana in the Jurassic are here identified as East Java, Bawean, Paternoster, West Sulawesi, Mangkalihat and SW Borneo. The East Java, Bawean, Paternoster, West Sulawesi and Mangkalihat blocks comprise Argoland, derived from the Exmouth Plateau region of western Australia. SW Borneo is identified as the ‘Banda block’ derived from the Banda embayment region of western Australia. Argoland and SW Borneo were accreted to SE Sundaland in the Late Cretaceous. Changing continent–ocean configurations during the Palaeozoic and Mesozoic resulted in the closure and opening of ocean gateways that resulted in changes in climatic and ocean current patterns, ocean upwelling and changing patterns of biotic dispersion. I would like to thank facilities provided by the Earth Sciences Division, School of Environmental and Rural Science, University of New England, and the National Key Centre for Geochemical Evolution and Metallogeny of Continents (GEMOC), Macquarie University. This paper has benefited from discussions with many colleagues

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Subsidence and uplift by slab-related mantle dynamics: a driving mechanism for the Late Cretaceous and Cenozoic evolution of continental SE Asia? BENJAMIN CLEMENTS1,2*, PETER M. BURGESS1, ROBERT HALL1 & MICHAEL A. COTTAM1 1

SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK 2

Statoil ASA, Forusbeen 50, N-4035 Stavanger, Norway *Corresponding author (e-mail: [email protected])

Abstract: Continental SE Asia is the site of an extensive Cretaceous– Paleocene regional unconformity that extends from Indochina to Java, covering an area of c. 5 600 000 km2. The unconformity has previously been related to microcontinental collision at the Java margin that halted subduction of Tethyan oceanic lithosphere in the Late Cretaceous. However, given the disparity in size between the accreted continental fragments and area of the unconformity, together with lack of evidence for requisite crustal shortening and thickening, the unconformity is unlikely to have resulted from collisional tectonics alone. Instead, mapping of the spatial extent of the mid–Late Cretaceous subduction zone and the Cretaceous– Paleocene unconformity suggests that the unconformity could be a consequence of subduction-driven mantle processes. Cessation of subduction, descent of a northward dipping slab into the mantle, and consequent uplift and denudation of a sediment-filled Late Jurassic and Early Cretaceous dynamic topographic low help explain the extent and timing of the unconformity. Sediments started to accumulate above the unconformity from the Middle Eocene when subduction recommenced beneath Sundaland.

Throughout continental SE Asia (referred to here as Sundaland; Fig. 1) there are almost no Upper Cretaceous and Paleocene strata, suggesting that much of the region was elevated during this time. Cenozoic rocks rest on older rocks with a profound unconformity. Rocks beneath the unconformity are considered basement and comprise predominantly Cretaceous and older granites, Mesozoic sedimentary rocks, accreted ophiolitic and arc rocks, and pre-Mesozoic metamorphic rocks. Sedimentary rocks above the unconformity are Eocene and younger (Fig. 1) and include siliciclastic, volcanogenic and carbonate lithologies that were deposited across the region in extensional halfgraben basins, and at the Sundaland continental margins. These deposits are initially commonly terrestrial and their depositional ages are often poorly constrained. The unconformity has previously been interpreted as the result of a poorly defined tectonothermal event that occurred throughout Indochina and the Malay Peninsula (e.g. Kra¨henbuhl 1991; Ahrendt et al. 1993; Dunning et al. 1995; Upton 1999) or as a consequence of continental collision at the Sundaland margin (e.g. Hall & Morley 2004; Smyth et al. 2007; Hall 2009). However,

there has been no attempt to estimate the extent of, or describe, the unconformity itself, or to propose a driving mechanism capable of producing uplift over such a large area. In this paper we demonstrate that the area covered by the unconformity is in excess of 5 600 000 km2 (c. 2000 by c. 2800 km – greater than the area of the Western United States; Fig. 2 or ten times the size of mainland France) and it extends from Indochina to SE Borneo and East Java (Fig. 1), and that micro-continental collisions in the Late Cretaceous were coincidental with the onset of regional uplift. However, from an assessment of the spatial extent of accreted continental crust, the corresponding extent of the unconformity (Fig. 1), and regional exhumation trends, we suggest that uplift of a Late Jurassic to Early –mid Cretaceous dynamic topographic low (DTL) also contributed to the formation of the unconformity. Uplift was triggered by termination of subduction, slab detachment and the resulting dynamic rebound across the region. Furthermore, we support the suggestion (Hall 2009) that the initiation of Cenozoic basin development may have been related to resumption of subduction at c. 45 Ma.

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 37– 51. DOI: 10.1144/SP355.3 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. (a) Our interpreted extent of the unconformity, the core of Sundaland as defined by Hamilton (1979), and the approximate positions of accreted ophiolitic and arc-type rocks and continental fragments added to the margin in the Late Cretaceous. (b) Sedimentary basins and their ‘ages’ in the region that is affected by the unconformity. Highlighted basins are the Malay, Nam Con Son, West Natuna, Central Sumatra, South Sumatra, Kutei, Barito and SE Java basins (Polochan et al. 1991; Gwang et al. 2001; Doust & Noble 2008; Smyth et al. 2008) – simplified lithostratigraphies are shown in Figure 6.

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Fig. 2. Indonesia, Malaysia and Indochina compared to the USA. The SE Asia Regional Unconformity is 5 600 000 km2 (c. 2000 by c. 2800 km) – greater than the area of the Western United States. The box is 608 from west to east and 308 from north to south (Indonesia and Malaysia comparison to USA modified from Hall & Smyth 2008).

Expected styles of uplift: collisionalv. mantle-driven Geologists have long recognized that convergence of continental lithospheric blocks or plates is accommodated by crustal thickening and regional uplift (orogenesis). This occurs over relatively short distances (tens to hundreds of km; Murrell 1986) perpendicular to the collision suture but may extend great distances along strike, forming an orogenic belt. During continental collision, deformation starts at the margin of the indenting plate with continued convergence leading to development of a fold-thrust belt that propagates outwards from the collision suture. Rapid exhumation of lower crustal material is common, marking significant rock uplift and related denudation, and the resulting unconformity often cuts deep into basement rocks with dimensions reflecting the pattern of crustal deformation (i.e. narrow across, and elongate along strike). The size of the indentor, rate of convergence and pre-existing structural trends are all important factors that may modify styles of deformation (e.g. Murrell 1986; Ellis 1996; Willingshofer & Sokoutis 2009), uplift and the extent of the resulting unconformity. Even major orogens such as the European Alps and New Zealand’s Southern Alps show topographic profiles that rarely exceed several hundred kilometres in width (Fig. 3) (Koons 1995). Deformation belts in smaller orogens involving less significant collisions are often much narrower. Where contractional deformation is observed away from continental margins these are often sites of older stretching and crustal thinning. Here, intra-plate stresses drive inversion that is initially linked to the reactivation of individual,

pre-existing (extensional) faults, with further compression also leading to fold-thrust belt development and orogenesis (e.g. the European Pyrenean orogen; Munoz 1992). Intra-plate stresses can also be manifest as vertical crustal movements that result from lithospheric folding (e.g. Lambeck et al. 1984; Cloetingh et al. 1999, 2006; Horva´th et al. 2006). Typically, theoretical studies suggest that coupled and decoupled behaviour of continental lithospheres generate mono- and biharmonic modes of folding respectively (Gerbault et al. 1999; Faccenda et al. 2009) and the spacing between regularly distributed folds can be expected to be 4–8 times the thickness of the competent layer (brittle crust) (Martinod & Davy 1994). However, examples of irregular lithosphere folding (e.g. Cloetingh et al. 1999), as observed in the Pannonian –Carpathian region, may be up to 25– 40 times (350–400 km wavelength) the thickness of the brittle crust (Dombra´di et al. 2010). In such instances of lithospheric folding the associated unconformity could be expected to be broad and with minimal incision,

Fig. 3. Topographic profiles from (a) Southern Alps, New Zealand and (b) the Western European Alps (modified from Koons 1995) illustrating that even major orogens rarely exceed several hundred kilometres in width.

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but not to extend beyond the fold crest. Importantly, physical models (Dombra´di et al. 2010) predict that such fold wavelengths are reduced dramatically when crustal heterogeneities are present, particularly if the lithosphere is hot and weak. Lithospheric flexure, driven by crustal loading/ subduction processes, may also contribute locally to uplift, although the viscoelastic nature of continental lithosphere dictates that such uplift cannot be sustained for (geologically) long periods of time (e.g. c. 104 to 106 years or more). Such estimates are dependent upon mantle viscosity and the flexural rigidity of the lithosphere. Processes in the viscous mantle (buoyancy forces/mass anomalies) have also been shown to drive uplift and subsidence of continental and oceanic lithosphere (e.g. Gurnis 1990, 1992, 1993; Lithgow-Bertelloni & Gurnis 1997). Mass anomalies in the mantle transmit stress to the base of the lithosphere via viscous flow and create dynamic topography (e.g. Gurnis 1993; Moucha et al. 2008). For example, viscous mantle flow associated with subduction of a cold, dense slab causes subsidence creating a dynamic topographic low (DTL; Fig. 10b). The DTL can extend thousands of kilometres from the subduction zone and have an amplitude of several hundred metres to more than 1 km depending on the dip and age of the slab (e.g. Burgess & Moresi 1999; Husson 2006; Steinberger 2007). When the slab detaches and sinks into the mantle the viscous forces maintaining the DTL are reduced or removed, and the area is uplifted (Fig. 10d). Any sedimentary rocks deposited in the DTL will tend to be eroded, generating an unconformity, the extent of which is similar to the original DTL (e.g. Gurnis et al. 1996; Burgess et al. 1997). Uplift is further accentuated by denudation and the resultant isostatic rebound, with a damped positive feedback driving further uplift.

Sundaland: continental SE Asia Much of Sundaland (Fig. 1) is allochthonous and it is a composite region of continental fragments (terranes), volcanic arcs and oceanic accretionary complexes that successively rifted and separated from the margin of eastern Gondwana at various times during the Palaeozoic and Mesozoic (e.g. Metcalfe 1996) and were added to a growing Eurasia. All of these terranes are interpreted to have been derived directly or indirectly from Gondwana (e.g. Sengor 1979; Audley-Charles 1983; Metcalfe 1988) based mainly on comparative studies of the stratigraphy, palaeontology and palaeomagnetism. The continental core of Sundaland comprises the Indochina–East Malaya Block and the Sibumasu Block, both of which separated from Gondwana in the Palaeozoic

and amalgamated with the South and North China blocks in the Triassic. Three further blocks were subsequently added to the core of Sundaland; the SW Borneo Block (Hall 2009; Hall et al. 2009) followed by the East Java– West Sulawesi Block (Smyth et al. 2007; Hall 2009) (Fig. 1) were both derived from Gondwana. The Dangerous Grounds Block (Fig. 1) was probably derived from the South China margin (Hall et al. 2009). Sundaland includes the landmasses of Borneo, Java, Sumatra and the Thai– Malay Peninsula and extends northwards into Indochina (Fig. 1) and is characterized by very little seismicity and volcanism in the interior, away from the active margins. The region has experienced terrestrial to shallow marine conditions for most of the Cenozoic. The area that lies between the major landmasses is referred to as the Sunda Shelf (Fig. 1) and is typically flat and extensively shallow with water depths rarely exceeding 200 m (Hall 2009). These features have led to the common misconception that Sundaland has been a stable area during the Cenozoic (see discussions in Hall 2002, 2009; Hall & Morley 2004) often being referred to as a shield or craton (Ben-Avraham & Emery 1973; Gobbett & Hutchison 1973; Tjia 1992, 1996) or plate (e.g. Davies 1984; Cole & Crittenden 1997; Replumaz & Tapponnier 2003). The apparent stability of the Sunda Shelf (e.g. Geyh et al. 1979; Tjia 1992, 1996; Hanebuth et al. 2000) has resulted in data from the region being used in global eustatic sea level curves (e.g. Haq et al. 1987; Fleming et al. 1998; Bird et al. 2007). Sundaland lithosphere however differs markedly from other regions (e.g. African, Australian, Baltic Canadian shields) of stability (e.g. Hall & Morley 2004; Hyndman et al. 2005; Currie & Hyndman 2006; Hall 2009), and exhibits high heat flow (Doust & Sumner 2007; Hall 2009) and low seismic velocities in the lithosphere and asthenosphere (e.g. Widiyantoro & van der Hilst 1997; Bijwaard et al. 1998; Ritsema & van Heijst 2000). These observations indicate that the lithosphere is thin and weak in the region (Hall & Morley 2004; Hyndman et al. 2005). These characteristics are a consequence of prolonged subduction (Hyndman et al. 2005) and are typical of other back-arc mobile belts such as the North American Cordillera and parts of the NW Pacific (Hyndman et al. 2005).

Evidence for Cretaceous subduction and collision in SE Asia Plutonic and volcanic rocks There are abundant plutonic and volcanic rocks of Jurassic and Early–mid Cretaceous age exposed in

SUBSIDENCE AND UPLIFT IN SE ASIA

Sumatra, SE Borneo, Vietnam, and along the eastern China margin that are generally accepted to be subduction related. These typically occur inboard from the zone of subduction complexes (where present – see below) and, in many places, demonstrate that there was subduction beneath the Sundaland – Eurasian margin prior to collisions in the early Late Cretaceous. In Sumatra there are abundant I-type plutons (Late Jurassic and Early Cretaceous) that are exposed along the entire active margin (McCourt et al. 1996) and which formed above a northeastward dipping subduction system (beneath Sundaland). Associated with these are volcanic rocks such as Lower Cretaceous andesites exposed in the Omblin Basin (e.g. Koning & Aulia 1985) as well as other examples from within the Sumatra Fault Zone (e.g. Rosidi et al. 1976). In SE Borneo there are andesitic lavas, tuffs and volcanic breccias that are assigned entirely to the Haruyan Formation by Wakita et al. (1998) or placed within the Alino Group by Sikumbang & Heryanto (1994) and Yuwono et al. (1988) that are interpreted to represent a volcanic arc suite. These lithologies are approximately Late Aptian to Cenomanian in age (115 –93.5 Ma) (Yuwono et al. 1988; Wakita et al. 1998). Widespread granitic magmatism in mainland eastern China during the Late Jurassic and Early Cretaceous is generally accepted to be subduction related. Jahn et al. (1976) suggest that a Cretaceous (120 –90 Ma) thermal event along the SE China margin was related to westward-directed Pacific subduction. Subduction-related magmatism had ceased in Southern China by 80 Ma (Li & Li 2007). Zhou et al. (2008) used geophysical data to propose that a Jurassic to Early Cretaceous subduction complex can be traced south from Taiwan along the present northern margin of the South China Sea and was displaced to Palawan by opening of the South China Sea. Cretaceous granites are also reported from Vietnam (Nguyen et al. 2004) with youngest ages of 88 Ma. There was probably collision along parts of the eastern margin of Sundaland during the Late Cretaceous. Hall et al. (2009) suggest that the Dangerous Grounds Block became part of the Asia margin at about 90 Ma having rifted from the China margin (Fig. 4). There is little evidence for subduction after 80 Ma along the South China margin.

Subduction complexes Cretaceous subduction complexes including ophiolitic and arc-type rocks are exposed along the west coast of Sumatra, in Java and in SE Borneo, and are products of prolonged subduction beneath Sundaland that continued until the early Late Cretaceous

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Fig. 4. Plate tectonic reconstruction of the region at c. 100 Ma (modified from Hall et al. 2009). Note that most of the region is surrounded by subduction zones and there is impending collision at both the southern and northeastern margins of Sundaland. The Sundaland region was in a dynamic topographic low (DTL) prior to collision. WA, Woyla Arc – exposed onshore Sumatra as the Woyla Group (Nappe) (e.g. Barber & Crow 2005); IA, Incertus Arc (after Hall et al. 2009) which is tentatively correlated with the Mawgyi Nappe of western Burma (Barber & Crow 2009).

(Fig. 1). In Sumatra, the Woyla Group includes ophiolitic rocks, pelagic and volcaniclastic sedimentary rocks (Fig. 5f), and basaltic-andesitic volcanic rocks, interpreted as a Late Jurassic–Early Cretaceous intra-oceanic arc (Barber & Crow 2005). The timing of collision with Sumatra is estimated at 98 –92 Ma (M. J. Crow, pers. comm. 2008) based on overthrust Aptian –Albian fringing reef carbonates (Fig. 5e) and associated metamorphism of rocks of mid Cretaceous age (Barber & Crow 2009). In Java, similar subduction-related lithologies comprise pillow basalts (Fig. 5a), cherts (Fig. 5b), limestones, schists and metasedimentary rocks, interpreted as arc and ophiolitic terranes (e.g. Parkinson et al. 1998; Wakita 2000). Ultrahigh pressure metamorphic rocks at Karangsambung, East Java, such as jadeite–quartz– glaucophane bearing rocks and eclogites (Fig. 5d), are diagnostic of subduction metamorphism (Miyazaki et al. 1998). Radiolarian biostratigraphy (Wakita & Munasri 1994) and K –Ar dates on muscovite from quartz-mica schist (124 –110 Ma; Miyazaki et al. 1998; Parkinson et al. 1998) yield Cretaceous

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Fig. 5. (a) Pillow basalts and (b) deformed radiolarian cherts from Karangsambung, Central Java, that were obducted during collision of the East Java– West Sulawesi microplate. (c) Sheared serpentinized peridotites from the Meratus Mountains accretionary belt, SE Borneo. (d) High temperature – very high pressure eclogite from Karangsambung, Central Java. This is of probable Cretaceous age and is related to the Meratus subduction system that existed prior to collision of the East Java– West Sulawesi microplate. (e) Limestone islets of the Woyla Group off the west coast of Sumatra, near Lhoknga, Aceh. These limestones have yielded Late Jurassic to Early Cretaceous fossils (Barber & Crow 2005, p. 41) and probably fringed volcanic cones of the Woyla oceanic arc (A. J. Barber, pers. comm., 2009). (f) Folded volcaniclastic sandstones of the Woyla Group, North Sumatra.

ages for subduction-related rocks. In the Meratus Mountains, SE Borneo, ultramafic rocks (Fig. 5c), basalt, chert, siliceous shale, melange and schist are interpreted to represent accreted arc and

oceanic-type crust (Parkinson et al. 1998; Wakita et al. 1998). Radiolarian biostratigraphy yields ages that range from early Middle Jurassic to early Late Cretaceous (Wakita et al. 1998).

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Timing of collision Collision of the East Java–West Sulawesi block (Hall 2009) was probably responsible for termination of subduction beneath Sundaland (Smyth et al. 2007). The collision must have been later than the youngest radiolarian ages associated with pillow basalts in Java and Borneo (early Late Cretaceous), and the fragment must have been in place before initiation of the present phase of

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subduction at c. 45 Ma (Hall 2002, 2009). New plate reconstructions (Hall et al. 2009), based on the evidence summarized above and the beginning of a widespread hiatus in magmatism along the margin, indicate that the fragment arrived between 92 and 80 Ma. Clements (2008) and Clements & Hall (2008) suggest an age of c. 80 –85 Ma based on U – Pb dating of zircons in Eocene fore-arc sandstones in West Java that record arc volcanism prior to collision and post-collisional magmatism.

Fig. 6. Regional lithostratigraphic chart. Basins have been chosen as representative of sedimentary fill across the region – basin locations are shown in Figure 1. The earliest sedimentary fill in many of the basins was deposited in a terrestrial setting and depositional ages are poorly constrained. Ascertaining more precise ages for the earliest stages of fill in these basins, particularly in SE Borneo, is the focus of current research. Note that ‘basement’ refers to the pre-Cenozoic section. Modified from Polochan et al. (1991), Gwang et al. (2001), Doust & Noble (2008) and Smyth et al. (2008).

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A Late Cretaceous to Eocene unconformity

Regional exhumation data

We draw attention to the size of the SE Asia Regional Unconformity and the SE Asia region in Figure 2 by comparing both to the United States. The extent of the unconformity is shown in Figure 1. Few sedimentary rocks of latest Cretaceous and Paleocene age (Fig. 6) are preserved within this area, except in NW Kalimantan and Sarawak where there are marginal marine and terrestrial clastic sediments (summarized by Hutchison 2005). Where observed on land – in places such as Sumatra, Java and Borneo – the unconformity is angular and overlies variously deformed rocks that are Cretaceous and older (Fig. 7 shows the unconformity in West Sumatra and Fig. 8a, b illustrates the nature of Eocene terrestrial sequences immediately above the unconformity). The unconformity has also been penetrated throughout the region by exploration drilling offshore (e.g. Hamilton 1979). In the next sections we discuss evidence for regional exhumation from thermochronological studies as well as the nature of rocks that were deposited in the DTL immediately prior to uplift in the early Late Cretaceous.

Thermochronological studies from across the region provide direct constraints on the widespread uplift, erosion and exhumation of basement rocks during the Late Cretaceous and Early Palaeogene. The paucity of such studies in the south (e.g. Java) reflects a lack of exposed pre-Cenozoic rocks, and an abundance of younger (Cenozoic) volcanic rocks that dominate the stratigraphy. The region that extends from the Shan Plateau of Myanmar and northern Thailand through Laos into the Lanping-Simao fold belt was elevated at the beginning of the Cenozoic (Hall & Morley 2004). This elevation is attributed to a ‘diffuse [and] poorly defined orogenic event’ by Upton (1999) and Hall & Morley (2004). Apatite Fission Track (AFT) studies in Thailand and Laos indicate slow cooling between 90 and 45 Ma (Racey et al. 1997; Upton 1999), and in NW Thailand Upton (1999) interprets ‘gentle’ cooling between 70 and 50 Ma, with modelled exhumation rates of 0.048– 0.083 km/Ma. Upton (1999) suggests that exhumation in NW Thailand was driven by minor (c. 600 + 200 m) ‘tectonic’ uplift that ‘started a cycle of erosional denudation driven by isostatic rebound’ sufficient to generate the observed levels of denudation. Along the western margin of the Khorat Plateau in eastern Thailand, 2.3–4.4 km of Jurassic–Cretaceous overburden is estimated to have been removed since c. 65 Ma, and 2– 6 km was removed during the Palaeogene in parts of western Thailand (Hall & Morley 2004 and references therein). Carter et al. (2000) interpret a common uninterrupted (slow) (c. 1–1.5 8C/Ma) cooling history in eastern Vietnam between 61 and 48 Ma based on AFT analyses with associated denudation of 1.4–2 km. K –Ar dates and ZFT ages from the Malay Peninsula also indicate a local increase in cooling rate in the Late Cretaceous (e.g. Kra¨henbuhl 1991).

Rocks beneath the unconformity

Fig. 7. The SE Asia Regional Unconformity from Silangkai, West Sumatra. Palaeogene sandstones unconformably overlie the Kambayau Granite (not dated but granites in the same area are Permian to Early Triassic) – the base of the Omblin Basin.

Continental red beds of the Upper Jurassic–Lower Cretaceous Khorat Group and lateral equivalents (Racey 2009) are exposed over large areas of eastern Thailand, Laos, Cambodia and parts of Vietnam, southern Thailand, and Peninsular Malaysia (Racey 2009; Fig. 9), and are relatively undeformed (Harbury et al. 1990). These are predominantly fluvial, alluvial and lacustrine facies with lagoonal sandstones, mudstones and limestones also present (Racey & Goodall 2009) (Fig. 8c, d). Causes of different styles of basin evolution for the Upper Jurassic–Lower Cretaceous red beds in the region are generally poorly understood (Racey 2009). The Khorat Group in NE Thailand is

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Fig. 8. (a) & (b) are Eocene siliciclastic sedimentary rocks that lie immediately above the SE Asia Regional Unconformity (a) Upper Eocene quartz-rich sandstones of the Bayah Formation, West Java – the lower sequence comprises overbank mudstones and crevasse-splay sandstones; the upper sequence is a massive channel sand body. (b) Sequence of conglomerates and conglomeratic sandstones of the basal member, Tanjung Formation, Barito Basin, SE Borneo. A small weathered granitic outcrop (basement) lies immediately to the right, out of shot, of the section. (c) & (d) are Mesozoic alluvial and fluvial red bed sedimentary rocks that lie beneath the unconformity over large parts of the region and were deposited in the DTL. (c) fluvial and alluvial pebbly sandstones from the Khorat Group, NE Thailand and (d) from the Tembling Group, Malay Peninsula. Note the similarity in lithological character between the Mesozoic and Cenozoic sequences.

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Cretaceous (Harbury et al. 1990). Upper Jurassic – Lower Cretaceous red bed sequences typically unconformably overlie older, more intensely deformed metasedimentary rocks (Gobbett & Hutchison 1973; Harbury et al. 1990). In the south between Java, Sumatra and Borneo, basement has been penetrated by offshore drilling. The most abundant lithologies are granitic rocks and low-grade metasedimentary rocks, with gneiss, quartz diorite, diorite, and mafic and silicic rocks occurring locally (Hamilton 1979). K –Ar ages from this area reported by Hamilton (1979) are predominantly Cretaceous. Lower Cretaceous unmetamorphosed limestones lie beneath the unconformity in one area north of West Java (Hamilton 1979).

Testing the hypothesis: an unconformity due to dynamic topography

Fig. 9. The extent of Late Jurassic – Early Cretaceous red beds in Indochina and the Thai –Malay Peninsula. These sequences show remarkable lateral continuity and there is little evidence of major faulting controlling deposition – all consistent with deposition in a DTL.

commonly interpreted to have been deposited in a thermal sag basin (e.g. Cooper et al. 1989) following Late Triassic extension and orogenic collapse of the Early Triassic Indosinian orogen. Others suggest a foreland basin setting (e.g. Lovatt-Smith et al. 1996; Racey 2009) that formed in front of a Jurassic orogenic belt that was situated to the north or NE. Racey (2009) draws attention to the lateral continuity of the Khorat Group stating that the deposits are of ‘broad lateral extent [and of] relatively uniform thickness’. Lovatt-Smith et al. (1996) note that on seismic data there is little evidence for faultcontrolled accommodation during deposition and that the Khorat Group formations have a mainly ‘layercake’ appearance. Gentle folding of Mesozoic strata in the Malay Peninsula has been interpreted as evidence for a phase of uplift in the mid –Late

As highlighted above, there is no evidence for extensive crustal shortening or orogenesis in the early Late Cretaceous across much of Sundaland. Throughout the region there are older (Palaeozoic?) rocks that are, in places, highly deformed and of medium to high metamorphic grade (e.g. the Malay Peninsula; Harbury et al. 1990). However, these rocks are commonly overlain by relatively undeformed and unmetamorphosed sedimentary rocks of Jurassic and Early Cretaceous age and must therefore represent older regional tectonic events. There is little evidence for the existence of belts of exhumed (high-grade metamorphic) rocks that might be expected had there been major Late Cretaceous –Paleocene orogenesis and subsequent uplift throughout the region. There is some evidence for crustal shortening at the Sundaland margins during the Late Cretaceous (e.g. the Meratus Mountains, SW Borneo; western Sumatra – discussed above) and this is interpreted as the result of collision, but this deformation did not extend into the Sundaland interior for any considerable distance. Significant deformation (thrusting) at the Sundaland margin is reported by Clements et al. (2009) to affect Neogene rocks in Java, but this deformation also cannot be traced very far northward away from the margin. We speculate that in the Late Jurassic and Early Cretaceous there was an extensive and broadly low-lying region dominated by fluvial and alluvial sedimentation (e.g. the Khorat Group and lateral equivalents; Racey 2009) with limestones developing at the continental margins (e.g. limestones in part of the NW Java Sea; Hamilton 1979) and perhaps other strata since removed. This setting is consistent with that expected for a DTL maintained

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by subduction through the Late Jurassic and Early Cretaceous and filled by Upper Jurassic and Lower Cretaceous shallow marine and terrestrial strata. Regional exhumation data are all consistent with gradual (slow?), long-wavelength uplift between c. 85 and 45 Ma. It is clear that previous studies consider this uplift enigmatic, concluding, for example, that the event was ‘poorly defined’ (Dunning et al. 1995) or ‘diffuse’ (Hall & Morley 2004). The estimated c. 600 + 200 m of ‘tectonic’ uplift required to positively feedback and drive further uplift and denudation reported by Racey et al. (1997) in western Thailand is comparable to that generated by dynamic topography above a subducting slab (e.g. Burgess & Moresi 1999). Sundaland crust was probably thin, hot and weak as a consequence of prolonged subduction beneath the region during the Late Jurassic and Early Cretaceous (see discussion by Hyndman et al. 2005) precluding major crustal up-warping. The unconformity cannot be a consequence of intra-plate stresses alone given its extent, and the extreme heterogeneous character of the Sundaland continental region (Hall 2011), which would only serve to dramatically reduce the wavelength of irregular lithospheric folding (Dombra´di et al. 2010). Importantly, the thickness of the crust is unlikely to influence dynamic topography, which is driven by mantle processes transmitting stresses to the base of the lithosphere. Of critical importance however is the ability to distinguish subsidence due to stretching from subsidence due to dynamic topography; something that has proven problematical in Cenozoic studies of the region (e.g. Wheeler & White 2000). Eustatic sea level fall from the early Late Cretaceous through Palaeogene could also have contributed to unconformity development. However, pre-Neogene eustatic history is essentially unknown, and eustatic models predicting Cretaceous to Palaeogene eustatic change differ significantly, even regarding long-term changes. For example, Miller et al. (2005) show a eustatic curve derived from backstripping of strata on the New Jersey coastal plain. Their curve features an approximately constant eustatic sea level from Cenomanian through Maastrichtian time, followed by a c. 30 m rise into the Early Eocene, and then a fall of about 60 m into the Oligocene. This differs substantially from the Haq et al. (1987) curve which shows a long term fall of c. 50 m from the Cenomanian to Maastrichtian, possibly because the Haq curve is based in part on stratigraphic studies of SE Asia, although the method and data used in derivation of the Haq curve has always been rather obscure. Further work is required to evaluate and understand the impact of these possible eustatic histories in SE Asia although clearly the impact of subduction-driven subsidence

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and uplift means that such regions are not as stable as has previously been assumed and their suitability for inclusion in studies that assess eustatic sea level may be problematical at best. At c. 85 Ma collisions terminated subduction beneath Sundaland at the Sumatra–Java– Borneo margin and possibly in northern Borneo at c. 80 Ma. The termination of subduction in a zone surrounding much of continental SE Asia had a profound effect on the region. The detachment of Tethyan oceanic lithosphere followed by its slow descent into the mantle drove regional uplift and reversed a Late Jurassic –Early Cretaceous DTL (Fig. 10). This process of slow mantle-driven uplift was accentuated by denudation and associated isostatic rebound further driving regional uplift.

Fig. 10. Schematic diagram illustrating how dynamic topography forms above subducting slabs due to stresses generated at the base of the lithosphere by the slabs negative mass anomaly in the mantle. (a) Predicted dynamic topography (Modified from Burgess & Moresi 1999) (note that the vertical scale from this profile is different to that of parts b, c and d). (b) Dynamic topography above a subducting slab at the Sundaland margin and, (c) above facing subduction systems for example, southern Sundaland at c. 100 Ma (as in Fig. 4). (d) Collision and detachment of oceanic lithosphere would drive regional uplift creating an unconformity similar to the Late Cretaceous– Paleocene SE Asia Regional Unconformity. Note that in places there may still be a DTL stratigraphic sequence preserved as is the case in SE Asia, particularly in the northern parts of Sundaland.

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We suggest that renewed deposition above the unconformity was in part related to the onset of subduction at c. 45 Ma (Hall 2009). Although the style of sedimentation during the Eocene and Oligocene (clearly localized in fault-controlled basins) was markedly different to that during the Jurassic and Early Cretaceous (laterally continuous and not fault controlled), the onset of sedimentary basin formation coincided with the renewal of subduction. The resumption of subduction at this time clearly impacted the Sundaland region but it also imposed some other tectonic mechanism on the area that may already have been hot and weak (Hall & Morley 2004; Hyndman et al. 2005). The apparent younging of Cenozoic basins northward (Fig. 1), although poorly constrained, seems to broadly coincide with the mapped extent of sedimentary rocks (Upper Jurassic and Lower Cretaceous Khorat Group and lateral equivalents) deposited in the DTL (Figs 1 & 9). These observations could be explained by a subduction-driven model, such as that presented here, in so far as the areas expected to be affected most by subduction-related dynamic subsidence and uplift are those in the south of the region, where opposing subduction systems were closest. Furthermore, is there a relationship between the extent of DTL sediments (or where all DTL sediments have been removed) and the location of Cenozoic basins in the region? Such questions are beyond the scope of this paper but may be the focus of future research concerning the impact of subduction on the region as well as better understanding the distribution of Cenozoic basins.

Conclusions The SE Asia Regional Unconformity is observed across Sundaland with an area of c. 5 600 000 km2 (greater than the area of the Western United States) and represents c. 40 Ma of missing time. Uplift was unlikely to have been driven solely by collision tectonics. Vertical displacement of Sundaland continental lithosphere by reversal of a slabrelated dynamic topographic low explains the spatial extent of the unconformity, the regional geology both above and below the unconformity, regional exhumation data and the duration of apparent uplift. Furthermore, we suggest that the onset of subduction at c. 45 Ma resulted in renewed subsidence and the development of the numerous petroliferous sedimentary basins that are now present throughout the region. This is the first attempt at assessing the extent of the SE Asia Regional Unconformity as well as providing a plausible explanation for its development; more detailed modelling may test these suggestions.

We are grateful to the consortium of oil companies who have supported the SE Asia Research Group for many years. We thank Anthony Barber, Duncan Witts, Andrew Racey and Inga Sevastjanova for permissions to use photographs in Figures 5e, f and 7; Figure 8b– d respectively. Clare White, Ian Watkinson, Cesar Witt and Anthony Barber are thanked for helpful comments on various versions of the manuscript. We also thank Christopher F. Elders and Jason R. Ali for their constructive and supportive reviews.

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Pre-Cenozoic sedimentary section and structure as reflected in the JavaSPANTM crustal-scale PSDM seismic survey, and its implications regarding the basement terranes in the East Java Sea J. W. GRANATH, J. M. CHRIST, P. A. EMMET & M. G. DINKELMAN* ION-GXT BasinSPAN Programs, 2105 City West Blvd, Suite 900, Houston, TX 77042, USA *Corresponding author (e-mail: [email protected]) Abstract: A new long-offset, long-record crustal-scale seismic survey of 9600 km called JavaSPAN was acquired in the Java Sea and Makassar Strait. The East Java Sea is underlain by continental basement with a prolonged multiphase history of deposition punctuated by extensional and compressional events. This East Java Terrane is a major component of SE Sundaland lying between the Meratus suture, the contemporary Java arc, and the west Sulawesi orogenic belt, but is poorly constrained on the north under the North Makassar Basin and in Kalimantan. A Precambrian to Permo-Triassic sedimentary section up to some 8.5 km in thickness overlies crystalline basement in a number of fault blocks and synformal structures below a strong angular unconformity. A thin overlap assemblage of Cretaceous to early Cenozoic sediments overlies that unconformity. Middle Eocene to Neogene clastic and carbonate rocks overlie another angular unconformity that marks the initiation of a well known history of Palaeogene extension, sag, and Neogene inversion. The East Java Terrane rifted from the Bonaparte-Arafura sector of northern Australia in the Jurassic and accreted onto a magmatic arc on the SW flank of what is now Kalimantan in the Cretaceous.

SE Asia records a history of late Palaeozoic to Mesozoic assembly of Gondwanan terranes and their subsequent modification by Cenozoic processes. The core of the collage in Indonesia has traditionally been called Sundaland (Hall & Morley 2004, and references therein), and its extent includes the islands of Sumatra, Java and Borneo and the shallow seas between them (Fig. 1). The southern margin of Sundaland runs from SW Java northeastwards along the approximate trend of the Karimunjawa Arch in the Java Sea and into SW Borneo (e.g. Hamilton 1979). Until the early 1990s, regions to the east and SE (Fig. 1) were usually regarded as a fringe of accretionary crust surrounding Sundaland, facing south and east in front of the Mesozoic granitic core of a long-lived magmatic arc that occupied most of present-day Kalimantan, an interpretation dating from Hamilton (1979). That fore-arc region spanned the Makassar Strait and extended into the East Java Sea and into west Sulawesi. There is now good evidence that there is continental crust underlying southeastern Java (Smyth et al. 2007), the East Java Sea (Manur & Barraclough 1994; this paper), the Makassar Strait (Granath et al. 2009) and much of West Sulawesi (Elburg et al. 2003; van Leeuwen et al. 2007). This region has been termed the East Java–West Sulawesi Block (EJ– WS Block) by Hall (2009) and Hall et al. (2009). This paper is focused on the southern parts of that block where a new deep crustal seismic data set called JavaSPAN has imaged the

pre-Cenozoic, and for purposes of simplicity in the text is termed the East Java Terrane (EJT). Younger deformation related to the opening of the North and South Makassar Basins has obscured the character of the crust to the north (Granath et al. 2009), so that whether or not that area belongs to the same terrane or is a separate fragment remains conjectural. The boundaries of the EJT are in general obscured by younger geology. The western boundary is marked by a zone of deformed ophiolitic and arc-related rocks with which the EJT collided, along the Meratus suture. The extent of EJT crust under the Meratus assemblage can be traced in three JavaSPAN lines (Fig. 2). Figure 3 is a part of line 4600 that shows contrasting vergence in the basement between the Meratus assemblage and the basement of the EJT. The southern edge of the EJT is buried by volcanic products of the Sunda Arc. The northern boundary is hidden by younger sedimentary rocks but is generally constrained by a belt of ophiolitic rocks in Sabah and by the Celebes Sea. The eastern edge of the EJT was the locus of Miocene accretion to Sundaland, and lies in western Sulawesi (Hall 2009). On the basis of JavaSPAN imaging Granath et al. (2009) placed the southeastern edge at the Salayar Islands where Palaeogene section of the EJT ends, and at the northern margin of the Flores Sea oceanic crust. The petroleum potential of the East Java Sea was established early in the exploration history of

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 53– 74. DOI: 10.1144/SP355.4 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Sketch map of central and eastern Indonesia, showing the East Java–West Sulawesi Block (Hall 2009) in grey and the East Java Terrane (EJT) within it. Sunda Trench and selected Neogene tectonic elements in eastern Indonesia shown in black (convergent), blue (extensional) and magenta (strike-slip).

Indonesia, with oil and gas discoveries in what has traditionally been called the ‘Tertiary.’ Attention was naturally focused on the shallow part of the section, neglecting what was thought to be economic basement. Seismic reflection techniques were consequently tuned to the Cenozoic strata. JavaSPAN has a much longer geological history imaged, both in the crystalline basement and in the Mesozoic and Palaeozoic sedimentary sections beneath the Eocene. This paper is focused on the results of improved seismic imaging of this southeastern corner of Sundaland, and for the first time analyses the character of the basement and preEocene sedimentary sections and their relationships to the more familiar Cenozoic strata and to the classic inversion structures of the East Java Sea.

Data and methodology The BasinSPAN program Since 2001 ION-GXT has acquired a number of regional 2-D seismic reconnaissance surveys (‘SPAN’ surveys) over areas of interest to the petroleum industry (ION 2009). The surveys match or surpass the parameters of conventional industrial surveys and consequently are fully compatible in resolution with both modern 2D and 3D industrial seismic imaging. Their lengths and locations are

designed to add to understanding of the tectonic history of the subject area by providing transects across basins and continental margins adjacent to open oceans, and their deep record lengths (to 40 km) are intended to provide the context within which to interpret the full crustal-scale basin architecture. Acquisition of JavaSPAN (Fig. 2) was completed in early 2008 in the Java Sea back-arc region from central Java to Tukang Besi in the south, and extending northward between Kalimantan and Sulawesi into the Makassar Straits. The regional aspects of the setting of this survey are discussed by Dinkelman et al. (2008) and the implications to the deep crust by Granath et al. (2009).

Acquisition and processing of SPAN surveys The JavaSPAN survey is typical of the more recent SPAN surveys. It is comprised of 9600 line/km of 2D data (Dinkelman et al. 2008). Acquisition parameters include a 25 m shot interval, 12.5 m group interval and maximum offset of 9000 m. The record length is 18 s and the data have been processed to pre-stack time (PSTM) images of 16 s and pre-stack depth (PSDM) images of 40 km record length. Processing for JavaSPAN was completed at the end of 2008 and benefitted from geological interpretation of interim products, such as brute stack and preliminarily migrated data, for the

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Fig. 2. Map of the East Java Sea region showing location of JavaSPAN survey profiles. Shaded area represents the region of crystalline basement added to Sundaland, the East Java–West Sulawesi Block (Hall 2009) the southern part of which is termed the East Java Terrane. The nature of the crust in the north is obscured by Cenozoic extension in the North and South Makassar Basins, where JavaSPAN seismic data have not distinguished older crystalline crustal features. More traditional rim of Sundaland shown in the red dashed line, after Hamilton (1979). Boundaries of EJT defined by JavaSPAN in magenta lines, line with teeth marking the leading edge of the overthrust Meratus assemblage. Red highlights on JavaSPAN lines show seismic lines with pre-Cenozoic structures, also figured in Emmet et al. (2009). Oil (green), gas (magenta), and dry exploration wells are shown in the background, from IHS database. Four wells discussed in text identified with green lines. Nature of pre-Eocene rocks in wells shown by colour-coded symbols, as noted in legend. Numbers beside well locations are K– Ar whole rock ages of bottom-hole rocks as in well reports and reported by Bransden & Matthews 1992.

purpose of providing velocity constraints to the deep section.

Potential methods modelling Modelling of public domain gravity and magnetic data, as well as ship-board data gathered simultaneously with the seismic data, is routinely used

to constrain the seismic processing stream for the deep crustal structure imaging and interpretation; such an approach places limits on the velocities of the deeply buried sedimentary and non-sedimentary rocks that are critical to the PSDM processing workflow. Such modelling (an example of which is shown in Fig. 3b) effectively integrates the seismic imaging with potential fields data, seismic refraction

56 J. W. GRANATH ET AL. Fig. 3. A portion of JavaSPAN line 4600 (location shown in Fig. 1), across the Meratus suture. (a) Full 40 km PSDM displayed in Seismic Micro Technology’s (SMT) average energy mode (square of maximum amplitude pixelated over length of seismic wavelet), often used in rocks with poor seismic impedance. Top of basement at about 2 km, with faults in black, seismic energy trains in thin pink lines. Thrust bounding Meratus assemblage in hanging wall and EJT in footwall shown as heavy magenta line. Transition to transparent zone at burgundy coloured horizon at 25–28 km (temperature-dependent?), and Moho in green near base of section. (b) gravity model used for the processing of line 4600, red line shows the part in a of this figure. Green curves at top are match of modelled and observed gravity profile. Courtesy Bird Geophysical.

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data and regional understanding of plate tectonics, crustal deformation and basin evolution. These potential field studies are incorporated twice into the processing/interpretation workflow, first during PSDM processing to sharpen imaging of deep, intrabasement structures, and secondly during a crustalscale interpretation of basement and its sedimentary cover. The initial components of these 2D models are high-resolution gravity and magnetic data acquired along the SPAN lines, sedimentary and crustal horizons from the CRUST 2.0 global model (which is based on several thousand seismic refraction data stations (Mooney et al. 1998)), and other control such as wells and information from published sources. The first phase of modelling focuses on understanding the geometry of intracrystalline crustal horizons and Moho to assist seismic processors with an understanding of subbasement velocity. The second phase of gravity modelling is focused on supra-basement features and variations in lithology that have an effect on the velocity

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structure, such as salt nappes or carbonate buildups. Guiding principles in modelling are: (1) to hold densities constant for layers throughout the study area, (2) to proportion magnetic anomaly wavelengths to their source depths, and, most importantly (3) to tie models to seismic refraction data, which independently fix the depths of basement and Moho.

Interpretation paradigm The large areas involved in SPAN data sets and the long distances across the surveys normally require that emphasis be placed on regional unconformities as the important seismic horizons, which tend to be time-transgressive. We have found it useful to adopt a conceptual model for the tectonic framework of SE Asia (Fig. 4) that emphasizes the important unconformities and ties them to the origin, departure, transit and docking of terranes. This analysis highlights previously unappreciated aspects of the geology and provides insight into the age and stratigraphic character of the pre-Cenozoic strata

Fig. 4. Interpretation paradigm (schematic) of the JavaSPAN survey, emphasizing unconformities that bracket the history of terrane formation. An allochthonous terrane experiences a history involving (1) rifting that eventually results in (2) separation from its origin terrane carrying a departure unconformity that is equivalent to the breakup unconformity in the origin, (3) transit toward its destination, acquiring a unique drift section, (4) collision with the destination terrane, involving an arrival unconformity that cross-cuts any structures that result and (5) an overlap assemblage of sedimentary rocks that stitches the allochthon and the destination together. Additional unconformities related to events during any of the phases are possible, and may overshadow these two more regional surfaces.

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and basement. The geological history from the viewpoint of the allochthonous terrane can be broken into separate episodes bracketed by ‘departure’ and ‘arrival’ unconformities: (1)

(2) (3)

commonality with the ‘origin’ terrane below the departure unconformity, which from the point of view of the source terrane, is equivalent to the breakup unconformity; provincial in-transit geological features, stratigraphic and structural, alien to both the source and destination terranes; and an overlap assemblage of rocks and postcollisional geology above an ‘arrival’ unconformity, whose character is common with the ‘destination’ terrane and foreign to the origin terrane.

For the EJT, the departure and arrival unconformities bracket a thin Jurassic to Cretaceous transit section that lies above NW Australian continental crust and its supracrustal cover. The familiar Cenozoic oil and gas habitat of the East Java Sea lies above a thin late Cretaceous to Palaeogene overlap assemblage which in turn lies above the arrival unconformity.

Previous work on the East Java Terrane The concept of assembly of SE Asia through accretion of Gondwana terranes is as old as the application of plate tectonic principles to the region (Audley-Charles et al. 1974; Hamilton 1979; Audley-Charles 1983), and has been refined as the geology has become progressively better known (e.g. Gatinsky & Hutchison 1986; Audley-Charles 1988; Metcalfe 1988, 1996, 1998; Hutchison 1989, 2007; Wakita & Metcalfe 2005). The history of rifting of continental terranes from Australia and sea floor spreading is discussed in Veevers et al. (1991). Evidence for the addition of continental crust at the Eurasian margin and amalgamation of SE Asia comes mainly from comparative studies of stratigraphy, palaeontology and palaeomagnetism and has been discussed by a number of authors (e.g. Sengor 1979; Audley-Charles 1983; Pigram & Panggabean 1984; Metcalfe 1988, 1996, 1998; Struckmeyer et al. 1993) and animated by Hall (2002). It is only recently, however, that the EJ–WS Block was recognized as a discrete continental fragment or composite of fragments. It was first proposed by Bransden & Matthews (1992) as the ‘East Java microplate’ and covered more or less the same area to which we apply the term EJT here. Bransden & Matthews (1992) did not define the boundaries for their unit, but indicated that it comprised an area added to SE Asia along the Meratus Suture/Karimunjawa Arch, extending

eastward to include eastern Java Island and western Sulawesi.

Nature of the crust in the East Java Terrane The basement is now thought to be a composite of continental basement blocks or terranes of Precambrian age (e.g. Smyth et al. 2007) potentially rifted away from north or northwestern Australia in the Jurassic and accreted to Sundaland in the middle Cretaceous. The approximate extent of this assemblage of basement blocks is shown by the shading in Figure 1. The presence of continental crust had previously been suspected in the East Java Sea area, but not its full geographical extent. Rose & Hartono (1978) recognized a conglomeratic section overlying preTertiary basement in the Paternoster Platform in Kalimantan, a section that is in turn overlain by mixed clastic rocks and the carbonate facies of the Platform. Hutchison (1989) interpreted this relationship to indicate the presence of a continental fragment, and we postulate the section is an overlap assemblage lying above the arrival unconformity. Metcalfe (1998) derived this continental terrane in the Late Triassic to Late Jurassic from Gondwana, associating its source region with northern New Guinea on his reconstructions (Metcalfe 1996). To the south Smyth et al. (2007) obtained Archaean to Cambrian-aged zircons from Tertiary lavas in the volcanic arc of eastern Java, the best indication to date of an ancient basement within the terrane. They identified a possible East Java continental fragment extending from southeastern Java Island through the southwestern arm of Sulawesi into north-central Sulawesi, similar to the extent advocated here. Onshore and nearshore eastern Java Island Sribudiyani et al. (2003) mapped basement units within the composite terrane and proposed a tectonic evolution. Hall (2009) grouped these occurrences of continental geology as the East Java –West Sulawesi block, recognizing that the area could comprise several smaller continental fragments, and Hall et al. (2009) suggest that the fragment arrived at the Sundaland margin in the mid Cretaceous. The JavaSPAN seismic data are the first to show a single coherent terrane with a common preCenozoic history occupies at least the southern part of the East Java –West Sulawesi Block.

Stratigraphy of the East Java Terrane The mid-Eocene and younger stratigraphy of the EJT is well known as that part of the section has been a key hydrocarbon producer onshore since the late 1800s and offshore since the 1960s, and is well-imaged in conventional industrial seismic

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data (e.g. Caughey et al. 1995; Emmet 1996). Much of that production comes from the so-called ‘Sunda folds’ (Eubank & Makki 1981) comprised of Eocene extensional features inverted during the Miocene. These afford the important Cenozoic unconformities guiding the interpretation of JavaSPAN shown in Figure 5. In terms of the interpretation paradigm, the Cenozoic rift, sag, and inversion phases are post-arrival sedimentary sequences. Those three unconformities mark the early Miocene inversion (top-most, red Lower Miocene horizon in Figs 5– 7), the midOligocene sag (orange horizon), and the base Middle Eocene onset of rifting (bright green horizon) at the base of the Ngimbang Formation (Fig. 5). The Lower Miocene unconformity bevels the Palaeogene rift-expanded section and seems to record a collision in the Sulawesi orogenic belt in Miocene time (Granath et al. 2009, and references therein). These unconformities are common to the Java and Makassar Sea region. Pre-Eocene rocks have been intersected by some wells in the East Java Sea (Bransden & Matthews 1992; Manur & Barraclough 1994; Sribudiyani

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et al. 2003). A variety of rock types are represented in those wells (Fig. 2) including, on the one hand, accretionary prism assemblages in the NW (on trend with the Meratus fore-arc system (Wakita et al. 1998)), and on the other granitic and metamorphic basement, indurated sedimentary rocks of unconstrained age and origin, and flysch-like rocks that contain unambiguously Cretaceous microfauna elsewhere. Pre-rift supracrustal sedimentary rocks of Cenozoic age are unknown onshore in Java (Smyth et al. 2005) and poorly represented if present offshore in wells in the EJT, but this might because they are barren of fossils. Some authors have indicated rocks broadly designated as ‘pre-Ngimbang’ are of early Eocene and possibly Palaeocene age (e.g. Phillips et al. 1991). The Cretaceous has been established in wells (Phillips et al. 1991; Bransden & Matthews 1992) but the top of the Upper Cretaceous succession is probably variable in age. As described below, unconformities cut into the Upper Cretaceous to various levels even within oil and gas fields. Phillips et al. (1991) described the pre-Cenozoic structure and stratigraphy in the Madura-Kangean area.

Fig. 5. Tectonostratigraphic scheme of the EJT stratigraphy. Cenozoic after Bransden & Matthews (1992) and Phillips et al. (1991). Palaeozoic and Precambrian lithostratigraphy assumed by analogy with Goulburn Graben-Arafura Shelf region (Struckmeyer 2006). See text for further discussion. This is the colour scheme used for the other figures in the paper.

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Fig. 6.

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They mentioned pre-Cenomanian ages for Cretaceous clastic rocks in that particular area, and from the description of the structural relationships these would seem to be the same section of Upper Cretaceous mentioned by Bransden & Matthews (1992). The details of the Cretaceous stratigraphy (the top of which is marked by the dark green unconformity in Figure 5 and subsequent figures) are key to the interpretation of the JavaSPAN dataset. Bransden & Matthews (1992) recognized that a mud-dominated pre-Cenomanian (age by reason of microfossils) section in wells correlated to a seismic-stratigraphic section that subcrops an angular unconformity at the base of the Cenozoic section. They drew analogies to other occurrences of deep water Cretaceous rocks in Sulawesi (Hasan 1991), and noted that some wells penetrated red beds younger than the dated Upper Cretaceous. That Cretaceous section in turn lies above another unconformity below which the section is more highly deformed. They attributed these structural relations to the collision of an East Java microplate with Sunda and suggested that reactivation of Cretaceous thrust faults may localize the Cenozoic extension and inversion. We interpret the muddominated Cretaceous section as the transit or drift section between the unconformities ‘d ’ and ‘a’, which has been gently folded by the collision event, and the overlying red beds as the overlap section (Fig. 5). The underlying more highly deformed section lies within the migratory block and is inherited from the origin of the EJT. Phillips et al. (1991) dismiss the possibility of hydrocarbon source potential in the pre-Cenozoic of the Pagerungan Field. Unpublished well completion reports, however, for the West Kangean-2 exploratory well, document gas flow from Cenomanian-late Aptian sandstone reservoirs and hydrocarbon source potential within that section, suggesting the pre-Cenozoic section has economic petroleum potential.

Geology and tectonics as reflected in JavaSPAN Many of the Sunda folds are underpinned in the East Java Sea by older structures in which strata are

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preserved in faulted synclines 20– 50 km wide and 5 –10 km thick. Several were intersected by lines of the JavaSPAN survey (Emmet et al. 2009) the locations of which are shown as red line segments in Figure 2; two are discussed in detail here and are illustrated in Figures 6 and 7. Emmet et al. (2009) referred to these as sedimentary ‘keels’ because of their synformal shape and the fact that invariably they penetrate deeply into basement below a marked angular unconformity. In fact, several structural styles are involved, ranging from extensional fault blocks to inverted graben with folded stratigraphy. The pre-Eocene in the EJT comprises two sedimentary packages separated by the profound angular unconformity above the keels (Fig. 8). We first discuss the impressive structures below the unconformity within the pre-Cenozoic keels, and then the thin section above that records the departure from its origin, transit and arrival of the EJT at Sundaland.

Internal stratigraphy and structure of the pre-Cenozoic keels The pre-Cenozoic sedimentary keels are imaged in several JavaSPAN lines (Fig. 2). These structures contain what appears to be a common sedimentary section, as discussed below, and taken collectively are interpreted to record a common series of events. The structures are truncated by a strong angular unconformity (Surface ‘F ’ in Figs 5 –7) above which lies a thin but important pre-mid Eocene section. Line 4700 (Fig. 6) shows the section below the ‘F ’ unconformity to be folded into a synformal structure that has localized Palaeogene normal faulting and subsequent Neogene inversion. A similar structure is illustrated by Emmet et al. (2009) on Line 4750: adjusted for vertical exaggeration, apparent dip in both limbs of both structures is about 308. Lines 4700 and 4750 strike nearly perpendicular to the ‘pre-Tertiary tectono-stratigraphic’ trend of Bransden & Matthews (1992, fig. 7) and lie downstrike from each other. Hence both appear to be close to true profile views. Internally, they show repeated periods of earlier extension and inversion, the last

Fig. 6. JavaSPAN seismic line 4700, as located in Figure 2; NW to left. Profile located along a trend of strongly inverted Palaeogene half graben trending east–west and may project into the structure in the next seismic line to the east (Line 4750), where only the deeper horizons are imaged (Emmet et al. 2009, Fig. 6). (a) interpreted; conventional amplitude display in Landmark colour scheme showing the most complete section of pre-Cretaceous sedimentary rocks, labelled units A through E. Horizons, from deepest to shallowest; dark pink at top of crystalline basement; blue top of unit A, brown top of unit B, light blue top of unit C, yellow top of unit D and base of unit E; dark green reflector is top Cretaceous (which, as noted in text this lies at or above the arrival unconformity). Bright green base of expanded mid-Eocene section with Oligocene (orange) and Miocene (red) surfaces as in Figure 5. (b) uninterpreted. Boxes locate detailed parts of line shown in Figure 9a, c.

62 J. W. GRANATH ET AL. Fig. 7. JavaSPAN seismic line 4600 in the vicinity of Kangean Island. Location shown in Figure 2. NW is to the left. (a) interpreted; conventional amplitude display in Landmark colour scheme showing the location of key wells that help to date the pre-Eocene section. Line crosses a region-wide trend of strongly inverted Palaeogene half graben c. 10 km to the west of Kangean Island. Horizons as presented in Figure 5. (b) gravity model for the segment of line 4600 shown in a, with colour-coded density distribution: magenta 3.39 g/cc, orange 3.09, dark yellow 2.9, yellow 2.7, green 2.5, blue green 2.4, light blue 2.25 and dark blue 2.15. Locations of crossing seismic lines 3600 and 3700 (Fig. 1) shown in vertical lines. The dark green coloured density shows the locations of thickened sedimentary section, green lines above are modelled and observed gravity profiles. Courtesy Bird Geophysical. (c) uninterpreted. Boxes indicate location of Figures 12 and 13.

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Fig. 8. Model unconformity relationships derived from Figure 6 showing (in dashed lines and italics) the major departure (letter designation ‘d ’), arrival (letter designation ‘a ’), and profound angular (‘F ’, for ‘Fitzroy’) unconformities defined in Figure. 5. Thicknesses of units schematic only.

of which is responsible for the keel structure. Figure 7 shows a different style in a section west of Kangean Island. A very thick Eocene rift section is more mildly inverted during the Neogene than the structure in Figure 6, and a thinner sequence below the ‘F ’ unconformity is apparently controlled by normal faulting. Only Units A and a small part of B in the Kangean West-2 area are preserved here, with the ‘F ’ unconformity cutting more deeply into the underlying structure. Line 4700 (Fig. 6) contains the most complete stratigraphic section internal to any of the preCenozoic keels, and forms the model for our interpretation. The stratigraphy below unconformity ‘F ’ is controlled primarily by the seismic data as no wells are known to have penetrated the youngest section in the keels. Hence the stratigraphy is of unknown age and lithotype. We have divided the stratigraphy of the keel shown in Figure 6 into five seismo-stratigraphic packages, oldest to youngest A –E (Fig. 8), based on their seismic character and bedding relationships. Units A and B are also represented in line 4600 (Fig. 7). NSA-1F (Fig. 9a, b) on line 4700 is the only well to have penetrated below unconformity ‘F ’ on a JavaSPAN line, where it encountered a well-indurated section of fine-grained clastic rocks within unit B. No useful age control is available from that section. The five packages represent more than 8 km of section, with the top not preserved. Package A, at the base, is a syn-rift section of variable thickness with chaotic internal reflections some of which are

quite strong. It is superseded by Unit B, a poorly reflective section which the NSA-1F well penetrated. Based on well descriptions and seismic character, Unit B might be a terrestrial section possibly containing red beds. Unit C is characterized by moderate to high amplitude continuous reflectors suggesting a well-bedded section. It thickens from north to south and shows some evidence of cut and fill sedimentation. It may be marine. Unit D shows low-amplitude discontinuous reflectors resembling Unit B and may in part be terrestrial or mixed marine and marginal marine in character. Unit E (post-yellow) is characterized by moderate to high amplitude continuous reflectivity indicating a probably well-bedded marine section, possibly of relatively fine-grained rocks. Figure 10 is a line-length reconstruction of the five stratigraphic packages of the keel illustrated in Figure 6. The extension in Units A and B is evident, with a narrow rift during A widening during the deposition of B, suggesting a prolonged period of at least intermittent rift-related sedimentation. Unit C thickens to the NW but the variation does not appear to be fault controlled, suggesting some regional control on differential subsidence. Within the small area represented by this line, the tectonic character of the environment of deposition of D and E is difficult to assess but they lie conformably above C with uniform thickness. Overall the history appears to be one of prolonged platform sedimentation punctuated by periods of extensionrelated subsidence, all culminating in a period of

64 J. W. GRANATH ET AL. Fig. 9. Detail of line 4700 (Fig. 6). (a) area around well NSA-1F. A strong angular unconformity shown in gold colour (the ‘F ’ unconformity (Fig. 5), separates section below (deformed in the Fitzroy movement before departure from Australia) from overlying rift, drift/transit, and overlap sections. (b) area around and to the SE of NSA-1F well flattened on the gold horizon ‘F ’ showing different relationships above and below. A thin section stands out between that of the keels and the inverted Late Cretaceous and Cenozoic section, representing the transit section. Circled area shows the region in a. (c) a similar box flattened on the Cretaceous reflector in the vicinity of the SG P-1 well, showing the thickness variations in the pre-Cenozoic transit section below the top Cretaceous and those within the overlap assemblage above the Cretaceous. The red horizon near the gold ‘F ’ is a model horizon showing the internal geometry of the interval rather than a specific unconformity. It may represent either the arrival or the departure unconformity, but it is not clear which. Compare Figure 13 in which both surfaces are apparent. Circled area shows penetration of SG P-1 well to just above the Cretaceous.

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Fig. 10. Flexural slip (bed length preservation) restoration of line 4700 focused on the five horizons A– E in the keel structure below the Cretaceous unconformity. (a) line diagram of the structure shown in Figure 4. (b) top of unit A restored to horizontality, showing the influence of extension on the deposition of the interval. (c) units A and B restored to horizontality with top of unit B as the datum; continued extension controlled the thickness of unit B. (d) restoration of units C, D, and E, showing uniformity in the thickness of the units except for the thickening of C across a fault in the section below.

strong inversion with significant shortening. The setting may well have been an intraplate one, well away from active margins. Evidently, the keel section from basement to above E accumulated, was deformed, and subsequently eroded prior to the Cretaceous. No definitive dates can be assigned to these events: the youngest age is controlled by the age of the section above the ‘F ’ unconformity. The thickness, seismic character and geological history of the EJT keels, however, bear similarity (1) to offshore NW Australia for the upper parts (Etheridge & O’Brien 1994) and (2) to the Arafura Platform for the entire section (Struckmeyer 2006). Figure 11 is a comparison of Arafura time seismic lines to the PSTM of the well developed keel in line 4700, with the ABCDE nomenclature used here assigned to units in the Arafura lines. This correlation is the basis for tentatively assigning the Arafura stratigraphic designations to ABCDE in Figure 5. Granath et al. (2010, fig. 3) drew an additional comparison between the compressional and specifically inversion geometry of Line 4700 and the Goulburn graben of the Arafura shelf (Struckmeyer 2006). If this interpretation is correct and these features are broadly time equivalents to those on the Australian Arafura Shelf then the ‘F ’ unconformity correlates to the angular unconformity generated during the ‘Fitzroy Movement’ (as defined in the onshore Canning Basin (e.g. Smith et al. 1999)), dated as Triassic.

Stratigraphy of the EJT: pre-departure, transit and overlap successions Overlying the angular unconformity ‘F ’ and below the Eocene rift-related unconformity is a stratigraphically complicated section that is key to understanding the timing of deformation within the keels below the ‘F ’ unconformity and the timing of arrival and accretion of the EJT to the Sundaland margin. Well penetrations in the East Java Terrane. A number of hydrocarbon exploration wells have been drilled to depths below the Middle Eocene unconformity, into so-called ‘pre-Ngimbang’ units. The results of many of these wells are compiled by Bransden & Matthews (1992) and in the Pagerungan field area by Phillips et al. (1991). We have reviewed many of the original well completion reports. Although what was assumed to be economic basement often gets cursory attention in hydrocarbon exploration, the data in those reports are important to the history of the EJT. The well reports, for example, sometimes include whole rock K –Ar radiometric dates, which in those reports are taken to be the age of the lithologies themselves. The well data fall into three broad categories: (1) Wells in the western part of the area (Fig. 2) that penetrate basement along the Karimunjawa Arch (Fig. 1) generally sample ultramafic and

66 J. W. GRANATH ET AL. Fig. 11. Seismic slice of pre-Cenozoic strata below the ‘F’ unconformity compared with seismic-stratigraphic sections from the Arafura area. See Figure 4 for lithostratigraphic assumptions. (a) the section on the north limb of the synform in line 4700 illustrated in Figure 6, PSDM, (b) same slice extracted from the pre-stack time migration (PSTM), in TWT (two-way travel time). Note that the deformation may have altered the stratigraphic thicknesses, (c) a portion of Geoscience Australia seismic line 094r-08 in the Goulburn graben, (d) a portion of Geoscience Australia seismic line 094r-09 in the Arafura Basin. (c) and (d) modified from Figure 3 of Struckmeyer (2006); A possible correlation of units A through E in East Java is indicated by the letters A through E in (c) and (d).

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(2)

(3)

mafic volcanic rocks that are probably associated with obducted ophiolites, similar to those exposed onshore in SE Kalimantan. Igneous rocks are arc-like in character, notably andesite, dacite, rhyodacite, monazite, and undifferentiated tuffaceous rocks and lavas. Granitoids occur in several wells, as do undifferentiated metamorphic rocks which may be roof pendant assemblages or fore-arc metamorphic facies. Ages of 92 to 105 Ma are reported in extrusive and metamorphic rocks. Wells from the longitude of Bali eastward generally encountered sedimentary rocks at total depth, some of which are called metamorphic in well reports, although this might simply mean well indurated. Ages are invariably reported as Cretaceous or Palaeogene at the very youngest. The age determination is sometimes supported by microfauna, but may in many cases be an assumed age. The Pagerungan Field area in particular is important: a mid Cretaceous clastic section was penetrated that consists of well lithified sandstones, quartzite, siltstone and shale of preCenomanian age (Phillips et al. 1991). The Upper Cretaceous is largely missing at Pangerungan although a thin section of clastic rocks lies between these Cretaceous rocks and the Middle Eocene with unconformable relationships both above and below. Strata appear to be variously oriented with some high dips as recorded on dipmeter logs. Vitrinite reflectance data suggest that this section had been more deeply buried, with about 4500 m of section removed (Phillips et al. 1991). Similarly the West Kangean field is well documented. The section in the West Kangean-2 well is dated by palynomorphs as no older than Late Aptian and no younger than Cenomanian. Again, the younger Cretaceous is absent in the wells. Unpublished well reports about the West Kangean-2 strata indicate that over 580 m of Cretaceous siliciclastic strata were drilled. The section consists mainly of claystones with occasional interbedded siltstones and sandstones grading to quartzite and is non-marine with minor marine influence. Bransden & Matthews (1992) cite numerous penetrations in the Kangean/Lombok area of a Cretaceous section, which is mud-dominated, with some siltstone and interbeds of tightly cemented, lithic to sub-lithic sandstone. The sections appear to become less marine upward and are reported to be highly indurated. In the same area as category 2 are wells that penetrated igneous or higher grade metamorphic rocks. The lithologies are dominated

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by granitoids. Mafic rocks are absent from this suite, except that in one case amphibolite was reported as was a tonalite on the far eastern edge of the EJT. Radiometric dates also give Mesozoic ages, 92 Ma for an extrusive in the Pagerungan area, but the tonalite north of Sumbawa gave an age of 65 Ma (Fig. 2). Granath et al. (2009) interpreted this latter sample to represent a volcanic complex developed on the trailing active margin of EJT after collision with Sundaland. The ages of the protoliths for the igneous and ‘metamorphic’ rocks require some interpretation. Our approach is to regard the radiometric dates as cooling ages rather than intrusive or peak metamorphic ages. Equivalents of the rocks of category 1 are exposed onshore in Kalimantan and Java (e.g. Wakita 2000 and references therein) where the structural relationships can to some degree be mapped, indicating the trend of the accretionary complex across the East Java Sea. We regard this area as the fore-arc prism against which the EJT docked on the edge of Sundaland. The rocks of category 2 represent the lithologies in various structural positions within the supracrustal succession. For example, the NSA-1F well (Figs 6 & 9) intersected the lower portion of the succession, near crystalline basement. Although the well description lists this lithology as metamorphic with no age control, it penetrated the lower portion of preCenozoic sequence B. Despite the Mesozoic radiometric ages, category 3 is interpreted to be the basement to category 2 rocks, with probable ultimate ages of Precambrian by analogy with potential provenance regions for the terrane, as discussed below. Linkage of well penetrations to seismic horizons. Three wells on and near JavaSPAN lines 4600 and 4700 provide constraints on the age of the seismostratigraphic units above the ‘F ’ unconformity. The NSA-1F well on line 4700 (Fig. 9a) penetrated a pre-Middle Eocene section lying on Cretaceous before it crossed the ‘F ’ unconformity. It is hard to pick the ‘d ’ and ‘a’ unconformities within the interval as they appear to be parallel to bedding, but flattening on the ‘F ’ unconformity (Fig. 9b) shows thickness variations both in the Cretaceous and the pre-Middle Eocene sections above ‘F ’ and a very different structure below ‘F ’. The SGP-1 well on Line 4700 (Fig. 9c) bottomed in Eocene sediments below the mid-Eocene unconformity, thus locating the top Cretaceous deeper than the well bore and suggesting that locally preNgimbang rocks of Cenozoic rather than Cretaceous age do exist in the area. The red surface below the top of the Cretaceous may be either the arrival or the departure unconformity, or both superimposed.

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Fig. 12. Detail of line 4600 (Fig. 7) near the Kangean West-2, which penetrated a late Aptian to Cenomanian age section, thus establishing a Cretaceous age below the arrival unconformity ‘a’, which is shown in blue (cf. Fig. 5). Red surface is interpreted as the departure unconformity ‘d’. Inset summarizes interpretation: l-mJ Lower– Middle Jurassic, lK þ uJ Lower Cretaceous and Upper Jurassic, uK Upper Cretaceous.

A portion of line 4600 (Fig. 7) passed near the Kangean West-2 well (Fig. 12). This penetration is important because it crossed the Cretaceous – Cenozoic boundary and thus also locates the top of the Cretaceous section in the seismic data. Figure 12 shows a downthrown block next to the high block penetrated by the well bore. The fault between the blocks is confined to the pre-Middle Eocene, and thus relates in time to the collision and intraplate adjustments immediately thereafter. The low block contains a peak-trough cycle that is not represented on the horst, indicating some of the section is lost to erosion on the high block during its uplift. The interpretation is that the Upper Cretaceous is lost to erosion on the high block, and that pre-Cenomanian rocks in the wellbore are immediately overlain by an unconformity we identify as the arrival ‘a’ unconformity. The Upper Cretaceous is preserved above the arrival unconformity in the low block. Thus the oldest date for the arrival of the EJT at Sunda is at the beginning of the Cenomanian or slightly younger. The inset in Figure 12 summarizes this interpretation. It is likely that the uppermost Cretaceous is largely absent in much of the EJT, and that time-wise the Cretaceous unconformity lies below the top of the Cretaceous Period and may, as here in the Kangean West-2 well be coincident with the

arrival unconformity. The thin section between the departure unconformity ‘d ’ (red) and the ‘F ’ unconformity (orange) represents the post-Fitzroy section overlapping the keel prior to departure from Australia, presumably Lower and/or Middle Jurassic rocks. Departure and arrival unconformities. The locations of the departure and arrival unconformities are subtle and often complicated by local structure as in the case of Figure 12. They are shown somewhat better in Figure 13, which is the southern end of line 4600 (Fig. 7) at the location of the ST Alpha-1 well. The well only penetrated Miocene limestones at TD, but this location shows better separation between the top Cretaceous and the ‘F ’ unconformities. The unconformities have angular relationships where they can be recognized. Because the pre-Cenomanian section in the area is mud dominated, we place it in the transit section between ‘d ’ and ‘a’ and predict it would be a Cretaceous open marine, possibly pelagic section accumulated during the drift across Neo-Tethys. The section above ‘a’ would therefore be the overlap section, and predictably it would be a discontinuous marine and non-marine section thickening toward the Meratus suture. The thickened Eocene section (between the bright green and orange horizons) in

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Fig. 13. Detail on line 4600 (Fig. 6) in the area around the ST Alpha-1 well. (a) The interpreted arrival ‘a’ and departure ‘d ’ unconformities above the major angular unconformity ‘F ’. (b) same section as in d, flattened on the ‘F ’ unconformity, showing the differences in orientation above and below.

line 4600 (Fig. 7) may in fact represent some of the apron of Meratus-related molasse, particularly if the bright green reflector, the mid-Eocene is placed somewhat too low. The seismic character of this section is relatively featureless and may represent non-marine rocks. The departure unconformity is not penetrated by any wells. Immediately beneath the transit section is a seismo-stratigraphic sequence of variable thickness that is fault controlled. This section is characterized by normal faults that cut the ‘F ’

unconformity but that do not penetrate the top Cretaceous reflector. This is interpreted to represent a syn-rift section that developed prior to departure of the EJT that is truncated by the departure unconformity ‘d ’. Figure 9b, c of line 4700 and Figure 13b (line 4600) have been flattened at unconformity ‘F ’ and Cretaceous levels. In each case a thin section lies between the flattened horizon and the red ‘d ’ unconformity, and the section above these levels is more conformable with the section above than that below. A small section of the syn-rift/pre-departure

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sequence is, therefore, inferred to be preserved in some parts of the JavaSPAN survey, even though the departure unconformity lies at or near the ‘F ’ unconformity or near the top of crystalline basement over much or the area. A more complete Mesozoic rift section may be anticipated near the edges of the block.

(5)

post-Middle Eocene cycle of rifting, subsidence, and inversion. Post-Middle Eocene. Above the mid-Eocene unconformity is a history of rift-related sedimentation and its inversion in the Miocene. This history is common throughout the EJT.

Discussion Summary: Pre-Eocene history The geological history of the EJT can be subdivided into five chapters, or megasequences as illustrated on the right column of Figure 5: (1) Precambrian through Palaeozoic. The seismic stratigraphy within the keels suggests the rocks are predominantly of sedimentary and, in the deeper parts of the keels, metasedimentary character. The sedimentary sequences reveal a history of intermittent extension and subsidence resulting in the accumulation of more than 8 km of sediments (Fig. 11). A major compressional event terminated the sedimentation. It is responsible for the formation of a major angular unconformity, referred to in this paper as unconformity ‘F ’. This event is thought to correlate with the Fitzroy Movement known from several locations in northern Australia. Outside of the keels crystalline basement lies immediately beneath unconformity ‘F ’. (2) Early and Middle Jurassic. Above unconformity ‘F ’ lies a thin section that pre-dates departure from the Australian location. It is poorly known, apparently only from seismic data, but would correlate with the rift-related section in Australia. Presumably if this is true then the section would be thicker on the southern margin of the EJT where it is hidden by younger geology. This section is capped by a Jurassic departure unconformity, labelled ‘d ’ in this study. (3) Upper Jurassic through mid-Cretaceous. Muddominated Cretaceous rocks with subordinate coarser clastic lithologies known from wells in the EJT are assigned to a transit section post-dating departure from Australia. They are capped by an unconformity marking the collision of the EJT with Sundaland, the arrival ‘a’ unconformity in this paper. (4) Late Cretaceous through Early Eocene. Cretaceous rocks above ‘a’ comprise an overlap assemblage marking the suture of the EJT, followed by poorly understood and probably only locally developed Paleocene and Early Eocene rocks for which the best evidence is in seismic data. This section is capped by the unconformity marking the beginning of a

The EJT, as the southern part of the East Java –West Sulawesi Block (Hall 2009; Hall et al. 2009) lies among a number of terranes with Gondwanan affinity. Hall (2009) and Hall et al. (2009) have attributed an Australian origin to SW Borneo, and Smyth et al. (2007) concluded that the likely source of the inherited Archaean– Cambrian zircons in igneous rocks of southeastern Java Island is a Precambrian to Palaeozoic basement with western Australian affinity, or a similar Gondwanan crustal province. Palaeozoic metamorphism of the Malino complex of northwestern Sulawesi suggests it has affinities with eastern Australian and more specifically the Bird’s Head of New Guinea (van Leeuwen et al. 2007). The southeastern Sulawesi terrane (including Buton) originated in northern New Guinea (e.g. Hamilton 1979; Pigram & Panggabean 1984; Struckmeyer et al. 1993) or off the northwestern margin of Australia/western New Guinea (e.g. Milsom et al. 2000; Hall 2002). Surono & Bachri (2002) compute a palaeolatitude of about 208S for Triassic palaeomagnetic samples from southeastern Sulawesi, compatible with the Late Triassic palaeolatitude for the northern NW Shelf of Australia and northwestern New Guinea (Struckmeyer et al. 1993; Wakita & Metcalfe 2005). It would be difficult to argue that the EJT was derived from a drastically different provenance, and in any case the source area is constrained to that region by reason of plate reconstructions (e.g. Hall 2002, 2009), Indian Ocean magnetic striping (Royer et al. 1989), and modern plate motion models, particularly Nuvel-1 (DeMets et al. 1990).

Significance of onshore stratigraphic sections A number of authors have previously drawn attention to ‘older’ sedimentary sections in the region. In this section we infer age and approximate stratigraphic position for these sequences based on the new JavaSPAN seismic lines presented in this paper. One of the earliest indications of continental crust beneath the East Java Sea area came from SE Kalimantan, where Palaeogene rocks occur on the Paternoster block and in the neighbouring ‘Meratus graben’ (Rose & Hartono 1978). The section is composed of coarse clastic rocks derived

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from the west that are overlain by a generally fining-upward sequence of sandstones, claystones, and finally carbonate facies of the Upper Eocene Berai formation. They designate the section a ‘preBerai’ formation. The lower part of the section is poorly dated but overlies metamorphic and intrusive rocks, and thus appears to overlap the forearc complex of the Meratus suture zone. The lithostratigraphy resembles post-orogenic molasse-like deposits, and in so far as the section grades up into post-collisional Eocene rocks, it appears to be equivalent to part of the overlap assemblage above the arrival unconformity for the EJT. Similarly, in southwestern Sulawesi, the Upper Cretaceous Balangbaru Formation also appears to be an overlap representative. It has been interpreted as a deep sea fan about 3300 m thick by Hasan (1991). The base lies in angular unconformity with the underlying metamorphic basement. Internal bedding is undeformed and dips slightly to the east, and the top is disconformable with the overlying Eocene Malawa Formation. The Balangbaru Formation is composed predominantly of interbedded sandstones and silty shales with some conglomerates, pebbly sandstones and basal conglomeratic breccias deposited in bathyal to abyssal water depths. Environments of deposition are interpreted to range from inner fan to basin plain. The channelized and sheet flow sandstones show fairly good lateral continuity. Provenance is interpreted to be from uplifted metamorphic basement, predominantly from the NW which would be internal to the EJT. Bathyal conglomerates and breccias suggest tectonic instability and we infer from this that the Balangbaru was deposited on the trailing edge of the EJT as part of the overlap assemblage.

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Basin where Neoproterozoic basins are largely confined to the south and west of the Pilbara (Archaean) craton (Fig. 14) (Cawood & Korsch 2008). The age of the crystalline basement in the EJT is poorly constrained. K– Ar dates represent cooling ages, so that the zircon data of Smyth et al. (2007) probably represents the best available information on the nature of the basement. Their data set spans the entire Neoproterozoic to Archaean with four well separated Precambrian peaks. As they point out, the spectrum of ages represented in Java compares with the spectrum of data from an alluvial dataset in the Perth Basin, but with rather stronger representation in several of the peaks older than 900 Ma. The spectrum suggests that an efficient mixing process supplied zircons to the crust beneath Java, perhaps a sedimentary succession with provenance from several adjacent basement terranes. Neoproterozoic (500 –1000 Ma) crystalline rocks are only well represented in the Pinjarra Orogen along the west coast of Australia (Fig. 14), but there are no sources of the older zircons immediately in that terrane. Neoproterozoic intrusive rocks also occur on the eastern margin of the Pilbara and within the Tasman orogenic belt (Cawood &

Possible origins for the EJT The stratigraphic succession of the EJT pre-Tertiary best compares with the section known from the Arafura Shelf of Australia (Fig. 11) especially considering distances over which significant stratigraphic variation is to be expected. Both sections are on the order of 8.5 km in thickness (Fig. 11) and can be divided similarly into five packages. If that correlation is correct, the section spans the entire Palaeozoic, and contains sedimentary rocks as old as Neoproterozoic. This contrasts with the Australian continental fragments of the eastern Indonesia islands, where the mid-Palaeozoic and older section is metamorphosed and deformed (Pigram & Panggabean 1984; Milsom et al. 2000; van Leeuwen et al. 2007), which would compare best with the Tasman orogen of eastern Australia rather than the NW Shelf of Australia. The thick section of old sedimentary rocks in the EJT also contrasts with the NW Shelf west of the Bonaparte

Fig. 14. Potential fit and size comparison of the East Java– West Sulawesi Block in three possible positions against the northwestern continental shelf of Australia. Rotations and internal distortions that may improve fit are not considered. The EJT is filled in solid colour. The block may be somewhat smaller when Eocene extension in the Makassar Basins is compensated. Relative merits of these positions is discussed in the text. Base map is a 155 Ma reconstruction provided by R. Hall. Grid lines are 108 squares. Phanerozoic basins noted along NW Shelf, epicratonic Proterozoic basins outlined, Pilbara craton in cross-hatched symbol; interior Precambrian complexes omitted.

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Korsch 2008), but only as igneous bodies rather than metamorphic complexes. Mesoproterozoic (1000– 1600 Ma) sources of zircons are similarly restricted to intrusive bodies on the east side of the Pilbara craton south of the Canning Basin and in central Australia. Orogenic Palaeoproterozoic rocks (1600–2500 Ma) are well distributed across northern Australia, as are Archaean. The Archaean cratons of Western Australia are obvious, but inliers in the Northern Territory are also well known (Plumb 1979; Worden 2007; Cawood & Korsch 2008). The best catchments to gather zircons from the surrounding terranes might be the Precambrian basins. In the west, basins surrounding the Pilbara could supply zircons of all the requisite ages. In the north, the Kimberley and McArthur Basins could equally supply the requisite suite as well as the Wessel Group offshore in the Arafura shelf, with which the lower section of the EJT keel succession is correlated here. It seems unlikely that the EJ–WS Block fits tightly against Western Australia as even the most conservative estimates of the extent of Greater India leave little room for such a fit (Hall 2009). Other possible positions on are shown in Figure 14. Position 1 is similar to the favoured one of Hall et al. (2009): it places the EJT against the Canning and Browse Basins, which accounts for the older zircons in Java (Smyth et al. 2007) but not the stratigraphic similarity of the Neoproterozoic and Palaeozoic sections in JavaSPAN to the Arafura Shelf, regardless of any rotations or adjustments of the fit. Position 2 brings that stratigraphic section northeastward where it could align with the Arafura in a belt of such late Precambrian sedimentary rocks outboard of Timor and the Bonaparte Basin, but leaves the Banda Sea Embayment as an odd shaped gap. The gap could, however, be partially filled by the Banda Ridges and an adjustment to the shape and/or orientation of the EJ–WS Block. Location 3 fills the Banda Sea gap, and brings the EJT closer to the Arafura Shelf. Any of the 3 could tap adjacent Australia for the observed mix of detrital zircons, especially if they are cycled through the epicratonic Proterozoic basins. Position 3 is viable only if the Sula Spur and the Bird’s Head have moved into their present position since departure of the EJ –WS Block. Hall (2009) and Hall et al. (2009) consider the ‘Sula Spur’ and Bird’s Head have remained in their present configuration at least since the Jurassic, but many other authors have concluded that both have been displaced to the west from locations variably adjacent to New Guinea (Granath et al. 2010, and particularly references therein), largely based on their similarity of Palaeozoic orogenic history to the Tasman belt. Locations 2 or 3, on balance, are the preferred placement of EJ –WS Block against

its ‘origin terrane,’ no doubt with adjustments to shape necessary to make a good fit.

Conclusions Structural relationships in the pre-Tertiary basement and overlying supracrustal section in the East Java Sea have been imaged clearly and systematically for the first time. These new data substantiate the pre-Cenozoic history involving a long period of sedimentation from late Precambrian through the Permo-Triassic that culminated in a compressional event and inversion of the section. The character of the section and its history suggests the terrane was derived from Australia, and particularly has character in common with the Arafura Shelf. The East Java Terrane was accreted to Sundaland in the mid Cretaceous, after which time it shared a well-known history of Eocene extensional tectonics and inversion in common with Sundaland. The fact that much of the pre-Cenozoic section may be preserved suggests heretofore unknown hydrocarbon system(s) may be present below current production levels. The authors wish to thank ION/GX Technology for its support and permission to show the JavaSPAN data at SAGE. In addition the help of Dale Bird of Bird Geophysical and the processing team at GXT, Ika Novianti and Vijay Singh, without whom the project would be entirely impossible, is gratefully acknowledged. Robert Hall kindly provided the base for Figure 14, a reconstruction with the East Java– West Sulawesi Block plotted in three potential restored positions against Australia. Benjamin Clements and Jurgen Adam provided thorough reviews that much improved the paper.

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Australia –SE Asia collision: plate tectonics and crustal flow ROBERT HALL SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK (e-mail: [email protected]) Abstract: The Sundaland core of SE Asia is a heterogeneous assemblage of Tethyan sutures and Gondwana fragments. Its complex basement structure was one major influence on Cenozoic tectonics; the rifting history of the north Australian margin was another. Fragments that rifted from Australia in the Jurassic collided with Sundaland in the Cretaceous and terminated subduction. From 90 to 45 Ma Sundaland was largely surrounded by inactive margins with localized strikeslip deformation, extension and subduction. At 45 Ma Australia began to move north, and subduction resumed beneath Sundaland. At 23 Ma the Sula Spur promontory collided with the Sundaland margin. From 15 Ma there was subduction hinge rollback into the Banda oceanic embayment, major extension, and later collision of the Banda volcanic arc with the southern margin of the embayment. However, this plate tectonic framework cannot be reduced to a microplate scale to explain Cenozoic deformation. Sundaland has a weak thin lithosphere, highly responsive to plate boundary forces and a hot weak deep crust has flowed in response to tectonic and topographic forces, and sedimentary loading. Gravity-driven movements of the upper crust, unusually rapid vertical motions, exceptionally high rates of erosion, and massive movements of sediment have characterized this region.

Eastern Indonesia is at the centre of the convergent region between the Eurasian, Australian and Pacific plates (Fig. 1). It is the site of the gateway between the ancient deep Pacific and Indian Oceans which disappeared in the Early Miocene as Australia began to collide with the Sundaland margin of Eurasia. Today it is the passageway for water which continues to move from the Pacific to the Indian Ocean, by complex routes reflecting the evolution of the collision zone since the Early Miocene. This tectonically complex region is known to biologists as Wallacea, with a biota and diversity as complex as the geology. Wallace (1869) recognized in the 19th century that biogeographical patterns in some way reflected geology but we are still very far from understanding the links between geology, palaeogeography, ocean –atmosphere circulation and climate which may have influenced the evolution of life. Unravelling the geology is a first step, but remains a difficult one. Here I discuss this first step: the geological development of the Australia –Asia collision, particularly in eastern Indonesia. The Cenozoic, particularly Neogene, development was strongly influenced by what was present before collision, so this paper begins with an outline of the Mesozoic and Early Cenozoic history of SE Asia, the Jurassic breakup of the northern Australian part of Gondwana and the assembly of Gondwana fragments in SE Asia in the Cretaceous. Rifting of fragments, now in Indonesia, from Gondwana was the first control on the Australian margin and the character of Sundaland, affecting

both the shape of the continental margins and the distribution of different types of crust within them. The nature of the Mesozoic Pacific margin is also touched upon, and the possible contribution of Cathaysian fragments to SE Asia. In contrast to most previous reconstructions of the region, the docking of different fragments is interpreted to have terminated subduction around SE Asia from the mid-Cretaceous until the Eocene, except for a short Paleocene episode of subduction beneath West Sulawesi. The effects of the assembly of different blocks, with their different internal structures, and separated by sutures, is then considered. The history of subduction resulted in an unusual lithosphere, and a high regional heatflow, and these features, combined with the heterogeneous nature of the basement were a major influence on Cenozoic deformation. It is argued that the Sundaland continent is not a craton or shield, but is a large region of generally weak lithosphere with weak and strong parts responding in a complex way to movement of the rigid plates that surround it. This determined the way in which the Australia–Asia collision proceeded and the deformational response of the upper crust to the movements of major plates, and the collision history is next reviewed, particularly the important subduction rollback into the Banda embayment of the Australian margin. Finally, I consider if plate tectonics can be reconciled with deformation of the crust and suggest that the region is not behaving as plates or microplates, as illustrated by different parts of eastern Sundaland

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 75– 109. DOI: 10.1144/SP355.5 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Geography of SE Asia and surrounding regions. Small black filled triangles are volcanoes from the Smithsonian Institution, Global Volcanism Program (Siebert & Simkin 2002), and bathymetry is simplified from the Gebco (2003) digital atlas. Bathymetric contours are at 200 m, 1000 m, 3000 m and 5000 m.

and Wallacea. I outline an alternative model explaining why the surface topography and bathymetry, and palaeogeography, have changed very rapidly during the late Neogene with important consequences for ocean currents, local climate, and probably global climate.

Assembly of SE Asia It is now generally accepted that the core of Sundaland (Fig. 2) was assembled from continental blocks that separated from Gondwana in the Palaeozoic and amalgamated with Asian blocks in the Triassic (Metcalfe 2011). The position of the eastern boundary of the Indochina–East Malaya block, the nature of crust to the east of it, and when this crust was added to Sundaland, are not known because much of this area is now submerged or covered with younger rocks. Only in Borneo are there rocks exposed that are older than Mesozoic. Most workers have assumed or implied that the

continental core of SW Borneo was attached to Sundaland well before the Cretaceous. Hamilton (1979) drew a NE –SW line from Java to Kalimantan widely accepted as the SE limit of Sundaland continental crust, implying much of Borneo was part of Sundaland by the Cretaceous and considered the region external to this core, from Sarawak to East Java, as Cretaceous and Tertiary subduction complexes. Many workers, including Hamilton (1979), Metcalfe (1988, 1990, 1996), Williams et al. (1988) have suggested broadly south-directed subduction beneath north Borneo during the Cretaceous and Early Cenozoic. Cretaceous north-directed subduction beneath south Borneo is indicated by the distribution of ophiolites and HP-LT metamorphic rocks in Java and SE Kalimantan (Parkinson et al. 1998). However, it is also possible that SW Borneo was added to Sundaland in the Cretaceous, much later than commonly assumed. Metcalfe (1996) shows most of the area north, east and south of Borneo as accreted crust, including

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Fig. 2. The Mesozoic and Cenozoic growth of Sundaland. It is suggested here that Sundaland grew in the Cretaceous by the addition of two main fragments: SW Borneo and East Java– West Sulawesi. In the Early Miocene new continental crust was added to Sundaland by collisions in Borneo and East Indonesia. Hamilton’s (1979) SE limit of Cretaceous continental crust is the part of the light blue line crossing the Java Sea from Java to Borneo.

several small continental blocks. There have been a number of suggestions for the origin of these continental fragments and Borneo crust, and when they became part of SE Asia. Ben-Avraham & Emery (1973) suggested a suture west of Borneo along the Billiton Depression interpreted as a transform fault associated with Cretaceous opening of the South China Sea. Metcalfe (1988, 1990, 1996) identified the SW Borneo and Semitau blocks, both with a South China origin, that moved south after rifting in the Late Cretaceous, opening the proto-South China Sea. Although the history of the Asian margin, and the interpreted age of the South China Sea, have changed (cf. Ben-Avraham & Uyeda 1973) an Asian origin for offshore Sarawak and much of Borneo has been supported by obvious Cathaysian characteristics of faunas and floras from the Dangerous Grounds (Kudrass et al. 1986), NW Kalimantan (Williams et al. 1988) and Sarawak (Hutchison 2005). Others have suggested an Australian origin for parts of Borneo. Luyendyk (1974) suggested the entire islands of Borneo and Sulawesi separated from Australia during Gondwana breakup in the Jurassic. Johnston (1981) proposed that a fragment rifted from the NW Shelf in the Late Jurassic collided with SE Asia in the mid-Cretaceous and underlies the area from Java to the eastern Banda

Arc. Smaller blocks have been interpreted as rifted from NW Australia in the Jurassic (Hamilton 1979; Pigram & Panggabean 1984; Audley-Charles et al. 1988; Metcalfe 1988; Powell et al. 1988). One major fragment was named Mt Victoria Land (Veevers 1988) or Argoland (Powell et al. 1988). Ricou (1994) suggested that Argoland corresponds to the Paternoster ‘plateau’ which he interpreted to have collided with Borneo in the Paleocene. However, most authors have interpreted the rifted Australian fragments to be much further away than Indonesia. Audley-Charles (1983, 1988) and Charlton (2001) suggested Argoland is now as far away as south Tibet, but it has most commonly been identified with West Burma. This view has been repeated so often that it has become received wisdom (Fig. 3) despite the fact that Metcalfe (1990, 1996), who first proposed it on the basis of Triassic (quartz-rich) turbidites above a preMesozoic schist basement similar to the NW Shelf, observed it was ‘speculative’ with ‘as yet no convincing evidence for the origin of this [West Burma] block’. Metcalfe (2009) has since abandoned the interpretation. In contrast, for other authors West Burma has been part of SE Asia since the Triassic and is therefore not Argoland. Mitchell (1984, 1992) argued that the Triassic turbidites in Burma were deposited on the southern

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Fig. 3. The most common interpretation of the origin and present position of the rifted blocks from NW Australia shown on reconstructions for 165 and 80 Ma, modified from Wakita & Metcalfe (2005). These interpret the Argo block as rifted from the NW Australian margin in the Late Jurassic and added to Asia as the West Burma (WB) block in the Cretaceous. SW Borneo is interpreted as separated from Asia during the Cretaceous by formation of the Proto-South China Sea (Proto-SCS).

margin of Asia, and Barber & Crow (2009) interpreted West Burma as a continuation of the West Sumatra block, now separated from it by opening of the Andaman Sea, which was part of Sundaland from the Late Palaeozoic. Pulunggono & Cameron (1984) proposed that north Sumatra includes the Sikuleh and Natal continental fragments, either rifted from Sundaland or accreted to it, and Metcalfe (1996) suggested these had a NW Australian origin. Barber (2000) and Barber & Crow (2005) reviewed these suggestions and argued that there is no convincing evidence for any microcontinental blocks accreted to the margin of Sundaland in the Cretaceous. They interpreted the Sikuleh and Natal fragments, like Mitchell (1993), as part of the Woyla intra-oceanic arc thrust onto the Sumatran Sundaland margin in the mid-Cretaceous. If the Australian rifted fragments are not in Tibet, West Burma or Sumatra then where are they? There is increasing evidence that they are in Borneo, West Sulawesi and Java, with some Cathaysian continental crust forming part of NW Borneo and the offshore shelf to the north of Sarawak and east of Vietnam, and that all these fragments arrived in their present positions during the Cretaceous.

Origin of crust of east Sundaland It is suggested here that the SE Asian promontory east of the Indochina–East Malaya block has grown by the addition of continental crust in two major stages: during the Early to mid-Cretaceous, and during the Neogene (Fig. 2). Some continental fragments have an Asian origin, but most are Australian. I suggest that an Asian fragment collided with east Sundaland, between Vietnam and northern Borneo, in the mid-Cretaceous and that Australian fragments also docked against the East Malaya block in the Early to mid-Cretaceous. A new reconstruction (Hall et al. 2009a) shows how these fragments moved into SE Asia (Figs 4 & 5).

Offshore Vietnam to Borneo It is commonly assumed that there was an eastfacing Andean margin with subduction of Pacific oceanic crust throughout the Mesozoic (e.g. Taylor & Hayes 1983; Metcalfe 1996) in the west Pacific. For South China and Indochina there is evidence for subduction in the Jurassic and Early Cretaceous but not in most of the Late Cretaceous. In the SE China margin Jahn et al. (1976) suggested that a Cretaceous (120– 90 Ma) thermal episode was

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Fig. 4. Reconstructions at 150 and 135 Ma. In the Late Jurassic the Banda blocks had separated forming the Banda embayment and leaving the Sula Spur. The Argo block separated slightly later, accompanied by a reorientation of spreading in the Banda embayment. Spreading propagated west, possibly along the continent– ocean boundary of Greater India to form the Woyla Arc. The arc and continental fragments moved away from the Gondwana margins as the subduction hinge rolled back. At 135 Ma India had begun to separate from Australia. Spreading in the Ceno-Tethys was predominantly oriented NW –SE and the Banda, Argo blocks and the Woyla Arc moved towards Sundaland as the Ceno-Tethys widened. Ex P, Exmouth Plateau.

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Fig. 5. Reconstructions at 120 and 90 Ma. There were numerous ridge jumps during India– Australia separation. In the Early Cretaceous the Banda block docked with Sundaland along a strike-slip suture at the Billiton Depression to become SW Borneo. Subduction continued beneath the Woyla Arc and probably south of Sumatra. At 90 Ma the Argo block docked with SW Borneo along the strike-slip Meratus suture, forming East Java and West Sulawesi, and the Woyla Arc docked with the Sumatra margin of Sundaland. The collisions terminated subduction. However, India continued to move north by subduction beneath the Incertus Arc (Hall et al. 2009a) which required formation of a broadly north– south transform boundary between the Indian and Australian plates. At about this time, Australia began to separate from Antarctica but at a very low rate.

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related to west-directed Pacific subduction. In South China, around Hong Kong, felsic magmatism ceased in the Early Cretaceous (Sewell et al. 2000), but it is not known whether magmatism continued outboard because this area is submerged. Further to the north in South China, there are younger Cretaceous magmatic rocks in a belt further east interpreted as subduction-related but magmatism had ceased by 80 Ma (Li & Li 2007). This suggests that there was a trench associated with west-dipping Pacific subduction, east of the present South China coast, in the Jurassic and Early Cretaceous but not in the Late Cretaceous after 80 Ma, and the subduction zone may have continued south across the South China Sea. Zhou et al. (2008) used geophysical data to trace a Jurassic – Early Cretaceous subduction complex south from Taiwan along the present northern margin of the South China Sea which they interpret to have been displaced to Palawan by opening of the South China Sea. This belt probably continued into Vietnam where there are Early Cretaceous granites (Nguyen et al. 2004; Thuy et al. 2004) with youngest ages of 88 Ma and may have terminated in south Vietnam or, less probably, continued into northern Borneo. From east Vietnam northwards there is no evidence for east-directed subduction after 80 Ma. There have been many suggestions of west- or south-directed subduction beneath north Borneo in the Late Cretaceous and Early Cenozoic (e.g. Hamilton 1979; Taylor & Hayes 1983; Williams et al. 1988; Tate 1991) although Moss (1998) identified problems with the common interpretation of the Rajang Group deepwater clastic sediments as subduction-related. He suggested that subduction had ceased by about 80 Ma after arrival of microcontinental fragments now beneath the Luconia Shoals and Sarawak, leaving a remnant ocean and a foreland basin in northern Borneo in which the Rajang Group was deposited. There is little evidence anywhere of subduction-related magmatism younger than about 80 Ma, and the Late Cretaceous was a period of rifting and extension of the South China margin (e.g. Taylor & Hayes 1983; Zhou et al. 2008). Although subduction has been interpreted in Sarawak (Hutchison 1996, 2005) and NW Kalimantan (Williams et al. 1988, 1989), Late Cretaceous and Early Cenozoic sequences are fluviatile and marginal marine. Dredged crust (Kudrass et al. 1986) from the Dangerous Grounds indicates the presence of a continental sedimentary rocks with Cathaysian affinities and metamorphic rocks with Early Cretaceous ages. In the rest of the region of offshore Malaysia and Vietnam little is known of the basement which is deep below a thick sediment cover. Recent offshore studies suggest a suture could continue towards the SW of Vietnam (Fyhn et al. 2010; Pedersen et al.

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2010). Hall et al. (2009a) interpreted a Luconia– Dangerous Grounds block of Asian origin, similar to that named Cathaysia by Zhou et al. (2008). Collision of this block between 90 and 80 Ma with a suture broadly in the position identified by Zhou et al. (2008) can account for Cathaysian continental crust, subduction melanges and magmatism in South China, Vietnam and NW Borneo. It does not require SW Borneo to have been part of Sundaland before this time.

South Borneo and Sulawesi There have been many suggestions that there was a collision between a Gondwana continental fragment and the Sundaland margin in the mid Cretaceous (e.g. Sikumbang 1986, 1990; Hasan 1990, 1991; Wakita et al. 1996; Parkinson et al. 1998) with a suture located in the Meratus region. Geochemical evidence (Elburg et al. 2003) and zircon dating (van Leeuwen et al. 2007) indicate continental crust may lie beneath much of west Sulawesi, and it has an Australian origin (van Leeuwen et al. 2007). Recent studies in East Java show that at least the southern part of the island is underlain by continental crust (Smyth 2005; Smyth et al. 2007, 2008). The igneous rocks of the Early Cenozoic Southern Mountains volcanic arc contain Archaean to Cambrian zircons and suggest a west Australian origin for the fragment (Smyth et al. 2008). Continental crust is also suggested to underlie parts of the southern Makassar Straits (Hall et al. 2009b) and East Java Sea between Kalimantan and Java, based on basement rocks encountered in exploration wells (Manur & Barraclough 1994). The evidence for the origin of SW Borneo is admittedly limited. Palaeomagnetism indicates it has been at its present latitude since the Cretaceous (Haile et al. 1977; Fuller et al. 1999). The Schwaner Mountains are dominated by Cretaceous igneous rocks which intrude a poorly-dated metamorphic basement suggested to be Permo-Triassic (e.g. Williams et al. 1988; Hutchison 2005) or older. The interpreted older ages are based on correlation of metamorphic rocks from Sarawak to Kalimantan (e.g. Tate 1991, 2002) across important sutures (Lupar Line and Boyan melange). However, there are convincing links to Australia. Devonian limestones from the Telen River in the Kutai basin (Rutten 1940) have a fauna resembling that of Devonian limestones from the Canning Basin (M. Boudagher Fadel, pers. comm. 2009). Alluvial diamonds from Kalimantan have many similarities to diamonds from NW Australia (Taylor et al. 1990). Interpretations of an Asian origin for SW Borneo discussed above were based on Cathaysian faunas and floras found in Sarawak and NW Kalimantan, but all these are within the Kuching zone (Hutchison

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2005) or NW Kalimantan Domain (Williams et al. 1988) in, or closely associated with, melanges and deformed ophiolites. These rocks are interpreted here as fragments of Asian material accreted during the Cretaceous which are not part of the SW Borneo block. SW Borneo is interpreted (Hall et al. 2009a) to be a block separated from the Banda embayment at about 160 Ma and added to Sundaland in the Early Cretaceous. This is consistent with the evidence for its origin discussed above, its size, and the age of rifting on the NW Shelf (Pigram & Panggabean 1984). The northern edge of the block was a south-dipping subduction zone as proposed by many authors (e.g. Hamilton 1979; Williams et al. 1988; Tate 1991; Hutchison 1996; Moss 1998) but was not continuous with the South China –Vietnam suture. A small Inner Banda block is interpreted (Hall et al. 2009a) to have followed the Banda block but to have moved relative to it during a later collision event, which may now underlie part of Sabah and northern West Sulawesi. SW Borneo accreted to Sundaland in the Early Cretaceous between about 115 and 110 Ma along the Billiton lineament that runs south from the Natuna area (Ben-Avraham 1973; Ben-Avraham & Emery 1973). The East Java–West Sulawesi block is interpreted as the Argo block, including the offshore continuation of the Canning Basin, whose detrital sediments provided the Palaeozoic to Archaean zircons found in East Java. The East Java –West Sulawesi block separated from NW Australia at about 155 Ma as rifting propagated west and south (Pigram & Panggabean 1984; Powell et al. 1988; Fullerton et al. 1989; Robb et al. 2005). East Java and West Sulawesi may include a number of separate fragments, rather than a single block, added to Sundaland at about 90 Ma at a suture running from West Java towards the Meratus Mountains and then northward (Hamilton 1979; Parkinson et al. 1998). Collision of the Woyla arc with the Sumatran Sundaland margin occurred at the same time as the East Java– West Sulawesi fragment docked (Hall et al. 2009a).

Termination of subduction The rifting of fragments from Australia determined the shape and character of the Australian margin which was to have a major influence on the Neogene development of Australia – SE Asia collision. The arrival of the rifted blocks also had a profound effect because they terminated subduction (Smyth et al. 2007; Hall 2009a, b; Hall et al. 2009a) around Sundaland in the mid-Cretaceous for 45 million years, and when subduction resumed in the Eocene their deep structure

influenced Cenozoic deformation of SE Asia. For the period 90 Ma to 45 Ma around most of Sundaland, except north of Sumatra, there was no subduction. Australia was not moving north, and there was an inactive margin south of Sumatra and Java until 45 Ma. Thus, no significant igneous activity is expected and little is recorded (Hall 2009a). The new reconstruction (Hall et al. 2009a) does, however, predict NW-directed subduction beneath Sumba and West Sulawesi between 63 Ma and 50 Ma where, in the latest Cretaceous and Paleocene, there was calc-alkaline volcanism interpreted as subduction-related (e.g. van Leeuwen 1981; Hasan 1990; Abdullah et al. 2000; Elburg et al. 2002; see Hall 2009a, for review).

Consequences for SE Asian lithosphere At present the interior of Sundaland, particularly the Sunda Shelf, Java Sea and surrounding emergent, but topographically low, areas of Sumatra and Borneo are largely free of seismicity and volcanism (Hamilton 1979; Hall & Morley 2004; Simons et al. 2007). This region formed an exposed landmass during the Pleistocene, and most of the Sunda Shelf is shallow, with water depths less than 200 m and little relief which has led to a misconception that it is a stable area. Sundaland is often described as a shield or craton, but seismic tomography, geological observations and heat flow (Hall & Morley 2004; Currie & Hyndman 2006) show that these terms are not appropriate. Unlike well-known shields or cratons Sundaland is not underlain by a thick cold lithosphere stabilized early in the Precambrian. P and S wave seismic tomography (Bijwaard et al. 1998; Ritsema & van Heijst 2000) show it is an area of low velocities in the lithosphere and underlying asthenosphere, in contrast to Indian and Australian continental lithosphere to the NW and SE (Fig. 6). Such low mantle velocities are commonly interpreted in terms of elevated temperature, and this is consistent with regional high heat flow, but they may also partly reflect the mantle composition or elevated volatile contents. Also unlike cratons, there has been significant deformation within Sundaland during the Mesozoic and Cenozoic. During the Cenozoic there was widespread faulting, the formation of numerous sedimentary basins, many of which are very deep, and localized but significant elevation of mountains (Hall & Morley 2004). Much of the Sundaland interior has high surface heat flow (Fig. 7), with values typically greater than 80 mW/m2, much greater than cratons (Artemieva & Mooney 2001). Likely causes are upper crustal heat flow from radiogenic granites and their erosional products,

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Fig. 6. 150 km depth slice through the S20RTS shear wave tomographic model of Ritsema & van Heijst (2000). High velocities are represented by blue and low velocities by red. Cratons are easily identified; SE Asia is not among them.

the insulation effects of sediments, and a large mantle contribution. The upper mantle velocities and heat flow observations suggest the region is underlain by a thin and weak lithosphere (Hall & Morley 2004) that extends many hundreds of kilometres from the volcanic margins but is probably a consequence of subduction (Currie & Hyndman 2006) beneath Sundaland throughout much of the Mesozoic until the midCretaceous and from the Eocene to present day. Critically, such ‘subduction back-arc’ lithosphere (Hyndman et al. 2005; Currie & Hyndman 2006) is not only significantly weaker than cratonic lithosphere but is likely to deform internally in response to plate boundary forces (Fig. 8) and to within-plate

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forces generated by topography (Lynch & Morgan 1987; Whittaker et al. 1992; Zoback et al. 2002). Regionally, the entire area north of the JavaSunda trench and west of the Philippine trench is underlain by weak lithosphere and is very responsive to plate boundary forces, but it is also heterogeneous. The long accretionary history of the region means that it is a composite mosaic of continental fragments (Fig. 9) with varying lithospheric thickness, different internal structures, crossed by sutures with different orientations, and cut by strike-slip faults of different ages (e.g. Allen 1962; Hamilton 1979; Sieh & Natawidjaja 2000; Barber & Crow 2009). For example, much of East Java, South Borneo, West Sulawesi and possibly parts of Sabah, are underlain by continental crust of Australian origin, and the rifted blocks brought with them the deep structure now observed. Deep structural lineaments, now oriented approximately NW–SE, are often traced across the whole of Borneo and commonly into Sulawesi (e.g. Satyana et al. 1999; Fraser et al. 2003; Gartrell et al. 2005; Puspita et al. 2005; Simons et al. 2007). Most are not active faults at present, although they are commonly represented in this way. Most of these lineaments show no signs of having been active faults during much of the Cenozoic, although a few have been reactivated. However, they do appear to have influenced the development of the region during the Cenozoic, and there are indications of changing basement character, depth to basement, and changes in sedimentary thicknesses across them. The lineament orientations are what would be expected if they are basement structures inherited from Australia (Fig. 10) where there are deep and old structures that can be traced offshore across

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Fig. 7. Contoured heat flow map for SE Asia, modified from Hall & Morley (2004).

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Back-arc

Fig. 8. Relative strengths of cratons and subduction backarc regions from Hyndman et al. (2005). Sundaland lithosphere is expected to be very responsive to plate boundary forces, especially when wet.

the NW Shelf and western Australia (e.g. Cadman et al. 1993; Goncharov 2004). Furthermore, deep seismic profiles between Borneo and East Java (Emmet et al. 2009) show many resemblances to the deep structure of Late Palaeozoic sedimentary basins from offshore western Australia (Fig. 11). This lithosphere was rifted from Australia and was accreted to form East Java, SE Borneo and South Sulawesi and is much thicker, cooler and stronger

than other parts of eastern Indonesia. This is reflected in the absence of significant Cenozoic deformation of these parts of Sundaland. In addition, in the eastern part of the region there are several oceanic basins of different ages. Thus, although overall the region is weak, it includes very strong parts. The complex deformation of the region during the Cenozoic reflects all these features in addition to the changing forces at the plate edges. This is illustrated particularly well by the collision of Australia with SE Asia.

Australia collision The composite character of the SE Asian lithosphere was a major influence on the way in which Australia–SE Asia collision developed, but also of great importance was the nature of the Australian margin. The Jurassic rifting led to formation of a continental promontory, the Sula Spur (Klompe´ 1954), that extended west from New Guinea on the north side of the Banda embayment. This embayment was part of the Australian plate and contained oceanic crust of Late Jurassic age. Its last remnant is the Argo Abyssal Plain SW of Timor. From the Late Jurassic to the Neogene the embayment was surrounded by a passive continental

Fig. 9. Sundaland blocks that were part of Sundaland by mid-Cretaceous, modified after Metcalfe (1996) and Barber et al. (2005). Ophiolitic sutures are shaded in green. West Sumatra, West Burma and Indochina– East Malaya were Cathaysian blocks added to Eurasia during the Palaeozoic. Sibumasu was accreted along the Raub–Bentong suture in the Triassic. West Burma and West Sumatra were subsequently moved along the Sundaland margin. The Woyla Arc was accreted in the Cretaceous. The Luconia and Dangerous Grounds blocks are interpreted to be Cathaysian fragments rifted from Asia and added to Sundaland in the Cretaceous. SW Borneo and East Java– West Sulawesi were rifted from West Australia and added in the mid-Cretaceous.

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Fig. 10. (a) Basement structure of Australia (Oz SeebaseTM Study 2005). For the continental crust the image highlights areas of exposed or shallow basement (mainly Archaean or Proterozoic crust) in shades of pink in contrast to areas with thick sedimentary cover in shades of blue. (b) Structure of the Canning Basin from Cadman et al. (1993). Pink areas are Archaean or Proterozoic basement. (c) Black lines show general trends of deep structures in NW Australia and predicted orientation of deep structures in Indonesia at the present-day if these faults were brought with accreted blocks from NW Australia according to the reconstructions of Figures 4 and 5.

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Fig. 11. (a) Deep structure of part of the NW Shelf modified from Goncharov (2004). (b)-. Deep structure of Java Sea modified from Emmet et al. (2009). Seismic lines are approximately the same horizontal and vertical scales and locations are shown on the inset map. TWT sec is two-way travel time in seconds. Both areas are characterized by old deep basement faults and thick sections of relatively undeformed sedimentary rocks of probable Palaeozoic and Mesozoic age.

margin that can be traced from the Exmouth plateau, via Timor, Tanimbar and Seram to SE Sulawesi. The shape of the embayment and great age of the oceanic lithosphere within the Australian plate were major influences on the way the collision developed. The spur, and the way in which it was fragmented during the Neogene (Figs 12 & 13), is the cause of many of the controversies about collision ages, and explains the present unusual distribution of continental crust in the present Banda Arc. Before 25 Ma there was subduction of oceanic lithosphere at the Java Trench which continued east into the Pacific north of New Guinea, south of the Sulawesi north arm, the Philippines and Halmahera. Soon after 25 Ma the Sula Spur began to collide with the North Sulawesi volcanic arc, and this is the first Australia –SE Asia collision. Ophiolites were thrust onto the continental crust, derived from the ocean north of the Sula Spur and probably from the North Sulawesi fore-arc, and are preserved today in East Sulawesi (Ku¨ndig 1956; Silver et al. 1983). Ophiolites in South Sulawesi represent other

parts of the oceanic crust between the Sula Spur and West Sulawesi; they may have been thrust east during the collision but more likely represent remnants of a Palaeogene transform margin at the eastern edge of Sundaland that have not been thrusted at all. The important points are that by the Early Miocene there was Australian crust in East and SE Sulawesi which continued east to the Bird’s Head, and there was no subduction of the embayment. Between 25 and 15 Ma the convergence between the Australian plate and Eurasia was absorbed in several ways: subduction of Indian ocean crust at the Java Trench; subduction of the Proto-South China Sea; broad non-rigid counter-clockwise rotation of Sundaland (Borneo, West Sulawesi, Java); internal deformation of Sundaland; and contraction, uplift and erosion in East and SE Sulawesi. There has been considerable controversy about reconstruction of the Banda region (see discussions in Hall & Wilson 2000; Hall 2002; Spakman & Hall 2010). Several authors have recognized, implicitly

Fig. 12. Reconstructions of the Banda region at 25 Ma and 15 Ma. Soon after 25 Ma the first stage of Australia–SE Asia collision began as the Sula Spur collided with the Sunda Arc in North Sulawesi. Farther north the Proto-South China Sea was almost eliminated by subduction beneath north Borneo. Green shading shows the extent of oceanic crust older than 120 Ma. By about 15 Ma the Java Trench propagated east along the northern continent –ocean boundary of the Banda embayment and subduction hinge rolled back to the SE.

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Fig. 12.

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or explicitly, the importance of subduction rollback in the Neogene development of the Banda arc (e.g. Hamilton 1979; Hall 1996, 2002; Charlton 2000; Milsom 2001; Harris 2006). I consider that rollback into the embayment is a key to understanding the development of the region, not only of the Banda arc but also of Wallacea, and the description here is based on a detailed model that links the tectonic reconstruction to the structure of lithosphere in the mantle (Spakman & Hall 2010). Subduction rollback into the Banda embayment began at about 15 Ma when the Java Trench became aligned with the northern side of the embayment, a tear fault developed from the western edge of the Sula Spur and propagated eastward along the continent –ocean boundary. As the tear moved east, the oceanic embayment began to sink rapidly by its own negative buoyancy and began the rollback of a subduction hinge into the Banda embayment. Australia advanced northward, and the subduction hinge rolled back into the Australian plate forming the west-plunging lithospheric fold defined today by seismicity. The exact time when rollback began is uncertain, but it was manifested by extension of the region above the Banda slab, which included parts of the pre-collision Sundaland margin in West Sulawesi and the collided Australian crust of the Sula Spur. An age between about 15 and 12 Ma is indicated by extension-related volcanic activity in West Sulawesi (Polve´ et al. 1997), core complex ages in the Sulawesi north arm (van Leeuwen et al. 2007), the beginning of spreading in the North Banda Sea (Hinschberger et al. 2000), and subsidence and volcanic activity near Sumba (Fortuin et al. 1997). Extension is interpreted to have occurred during three important phases. The earliest phase led to formation of the North Banda Sea between 12.5 and 7 Ma (Hinschberger et al. 2000). Extended continental crust from the Sula Spur was separated from that remaining in East and SE Sulawesi and transported into the region of the upper plate above the subduction hinge. Some of this crust remains in the Banda Ridges, and some forms part of the basement of the Banda volcanic arc and its fore-arc east of Flores. The eastern part of this arc, from east of Wetar to Seram, was active only during a short period (c. 8–5 Ma) of volcanic arc magmatism before a second major phase of

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extension led to formation of the South Banda Sea (Hinschberger et al. 2001). During opening of the South Banda Sea continental crust was again extended and carried into the Banda fore-arc; this crust is now found in Timor and several of the small outer arc islands from Leti to Babar (e.g. Bowin et al. 1980). Volcanic arc activity continued in the Inner Banda Arc from Flores at least as far east as Wetar, but continued rollback of the subduction hinge led to collision between the southern passive margin of the Banda embayment and the volcanic arc which began in East Timor at about 4 Ma (Audley-Charles 1986, 2004) and led to termination of volcanic activity from Alor to Wetar. The southern passive margin of the Banda embayment had an irregular shape with a number of rectilinear offsets similar to the present-day Exmouth Plateau: one south of Sumba and another in the region of East Timor. Uplift and thrusting of Australian continental crust began earliest in East Timor, whereas in the remnants of the oceanic embayment to the east subduction rollback continued. The most marked final phase of extension of the upper plate above the retreating hinge led to the formation of the Weber Deep which subsided from fore-arc depths of about 3 km to its present-day depth of more than 7 km in the last 2 million years. Two issues have plagued interpretation of the Banda arc. One is the direction and rate of convergence, and the second is the age of collision. There is no requirement for the two slabs often postulated (e.g. Cardwell & Isacks 1978; McCaffrey 1989; Das 2004; Hinschberger et al. 2005; but cf. Hamilton 1979) to account for features such as dip directions of the lithosphere and the apparent rise in the north-dipping subducted slab from depths of several hundred kilometres shown on NNE– SSW sections drawn parallel to Australia–SE Asia convergence direction. It is not Australia–SE Asia convergence, but the rollback of the subduction hinge into the embayment accompanied by deformation of the slab in the mantle, that accounts for the shape of the arc and the subducted lithosphere (Spakman & Hall 2010). During most of the Neogene the rollback direction was broadly to the SE, but after arrival of the Banda volcanic arc at the southern passive margin of the Banda embayment, rollback was east-directed forming the Weber Deep. The size and shape of the embayment

Fig. 13. Reconstructions of the Banda region at 8 and 4 Ma. Rollback of the subduction hinge into the Banda embayment stalled briefly, and spreading ceased in the North Banda basin. By this time the Sula Spur had been fragmented leaving remnants in Sulawesi and the Sula-Banggai Islands. Extended crust on the south side of the embayment was later left as the Banda Ridges in the central Banda Sea as spreading began in the South Banda Basin at about 6 Ma. At about 4 Ma collision between the Banda volcanic arc and the southern continental margin of the embayment began in East Timor. The irregular shape of the margin resulted in a complex collision between Timor and Sumba, while rollback continued to the east forming the Weber Deep.

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are similar to the slab dimensions inferred from seismicity, but seismic tomography indicates that part of the continental lithosphere underlying the western Sula Spur must have also been incorporated in the subducted slab (Spakman & Hall 2010). In the Banda region, notably in Timor but also elsewhere in the Banda arc, a variety of radiometric ages (most are K –Ar ages with a few Ar –Ar ages) has led to much confusion and claims of multiple or pre-Pliocene collisions (e.g. Berry & McDougall 1986; Richardson 1993; Reed et al. 1996; Linthout et al. 1997; Charlton 2002; Keep et al. 2003, 2009; Harris 2006). As discussed by Standley & Harris (2009) for Timor these ages record numerous episodes in the development of the Banda region. Pre-Miocene metamorphic ages represent events predating Australian collision with the SE Asia margin. Collision of the Sula Spur with the Sulawesi north arm occurred soon after 25 Ma and is now recorded by cooling ages of metamorphic rocks in Sulawesi (Parkinson 1998a, b; van Leeuwen et al. 2007), Timor (Berry & McDougall 1986; Standley & Harris 2009), dredged samples from the Banda Ridges (Silver et al. 1985), and from Kur in the Banda fore-arc (Honthaas et al. 1997). The ages do not record the time of collision at the place the rocks are now found, because they have been moved to their present positions by extension of the upper plate above the retreating subduction hinge (see reconstruction in Spakman & Hall 2010). Even younger metamorphic and igneous ages, such as those from rocks dredged in the Banda Ridges and from Timor, do not record the age of collision of the southern margin of the Banda embayment with the Asian margin. For example, many authors have followed Berry & Grady (1981) and Berry & McDougall (1986) in interpreting high grade metamorphosed rocks with cooling ages of 8 Ma as marking collision of Australian crust with the Asian margin despite failing to explain why, after collision with the arc, volcanic activity continued until 3–4 Ma. For example, a Late Miocene collision age has been interpreted in Sumba (Keep et al. 2003) and even older collision ages suggested for Timor (Keep et al. 2009). The problem cannot be solved by making distinctions between continental crust supposedly of Asian and Australian origin (cf. Charlton 2002; Harris 2006). As explained above, parts of the Cenozoic SE Asian margin were underlain by continental crust of Australian origin that arrived in the Cretaceous. On Timor Standley & Harris (2009) demonstrated an important difference between the Banda Terrane, which has detrital zircons up to midCretaceous age, and Australian continental margin basement which has no detrital zircons younger than Permian–Triassic. The Banda Terrane was part of the Asian margin from the mid-Cretaceous

but its basement includes continental rocks with a West Australian provenance (Hall et al. 2009a) probably similar to those beneath the Australian continental margin that collided in Timor in the Pliocene. More important is the Early Miocene collision of the Australian origin Sula Spur and its subsequent extension and fragmentation during slab rollback (Spakman & Hall 2010). The Banda allochthon in Timor is a complex including continental crust and arc rocks that formed part of the Early Cenozoic Asian margin, their overlying sedimentary rocks, Australian continental crust that collided in the Early Miocene, and Neogene arc rocks formed during the subduction of the embayment. Elsewhere in the Banda Arc young metamorphic ages have been used to interpret complex tectonic collision-related scenarios. For example, metamorphic and igneous rocks from Seram with ages of 5.5– 6 Ma have been used to infer formation of an ophiolite at 15 Ma and obduction at about 9 Ma (Linthout et al. 1997). Not only are the rocks completely unlike any other sub-ophiolite metamorphic rocks, it is also difficult to reconcile the proposed two dimensional reconstruction with any reconstructed map of the West Pacific since it requires the Banda volcanic arc to be placed 2000 km north of Timor in the Middle Miocene. Some confusion results from use of the term collision, but more follows from the assumption that metamorphic ages must mark contractional deformation that accompanied collision – in fact the K –Ar and Ar–Ar ages simply record cooling, which in most cases resulted from extension. Neogene metamorphic ages record extension of this complex upper plate. I interpret all the post-Sula Spur collision metamorphic and igneous ages, mainly between 12 and 4 Ma, to record extension of the upper plate, including Australian-origin continental crust and the Banda fore-arc, and tearing along the northern oceanic –continent boundary of the embayment during rollback. In eastern Indonesia the first contact of the Australian continent and the Asian margin was soon after 25 Ma. Rollback into the Banda embayment began at between 15 and 12 Ma. Volcanic activity in the western Banda arc began at about 12 Ma. The tear along the northern oceanic –continent boundary stalled or ceased at about 6 Ma near west Seram, juxtaposing continental crust and hot mantle by delamination (Spakman & Hall 2010), causing melting and metamorphism, later exhumed. In Timor and Sumba the arc –continent collision age of about 4 Ma is marked by a cessation of volcanic activity in the inner Banda arc in Wetar and Alor by 3 Ma (Abbott & Chamalaun 1981; Scotney et al. 2005) and by the rapid uplift that followed collision which moved sedimentary rocks deposited at depths of several kilometres below

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sea level to their present positions of several kilometres above sea level (e.g. Fortuin et al. 1997; Audley-Charles 2011). The very young volcanoes in the eastern part of arc from Damar to Banda (Abbott & Chamalaun 1981; Honthaas et al. 1998, 1999) record the latest and final stage of rollback into the last remnant of the embayment that accompanied formation of the Weber Deep. However, the extension that accompanied rollback into the embayment formed new oceanic basins in the Banda Sea. It can be argued that Australian– SE Asia collision began about 25 million years ago, continues today, and is likely to continue for many millions of years to come as these small basins are destroyed. The East Indonesian region provides a useful perspective on the debate about India –Asia collision age, variously estimated as between 60 and 35 Ma. For a flavour of this debate see Rowley (1996), Aitchison et al. (2007), Garzanti (2008), Khan et al. (2009) and Yin (2010). If there has been disagreement about ages in East Indonesia, where collision began more recently, is less advanced, and where continental margins can be reconstructed, it is easy to see why there is controversy surrounding timing of events in the larger and much more deformed Himalayan orogenic belt.

Plate tectonics v. deformation of the crust Although the plate model used here (Hall 2002; Hall et al. 2009a) provides a good first order understanding of the history and development of the region it is less clear that plate tectonics provides an adequate

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basis for understanding the details of the present or Cenozoic deformation of the region. Clearly, plate movements have been a major control and have led to the complex distribution of blocks, sutures, and the character of the lithosphere discussed above. However, it is impossible to draw continuously connected plate boundaries between major plates (Australia, Pacific, Philippine Sea) surrounding SE Asia, for example to join the Java Trench to the Philippine Trench. The problem remains even if smaller plates are postulated, and this is critical for modelling deformation since interpreted regional stresses are critically dependent on the geometry and position of inferred plate boundaries. Furthermore, much of the SE Asian region is continental, and deformation of continents is significantly less well understood than that of the strong oceanic plates; we now know that deformation is, and has been, much more complex than interpreted from models of rigid blocks separated by narrow fault zones that cut the lithosphere (e.g. Thatcher 2009). An alternative is to consider large parts of the region as a diffuse plate boundary zone (Gordon 1998) or wide suture zone (Hall & Wilson 2000; Hall 2009b) within which there is deformation. This type of approach ‘implies that a deforming zone is bounded by two (or more) rigid or nearly rigid plates in motion relative to each other’ (Gordon 1998) but raises the question of identifying the rigid areas. Gordon (1998) represents east Asia and SE Asia (Fig. 14) as a very large deforming region with large rigid parts, such as the Yangtze, Indochina and Borneo ‘plates’ or ‘blocks’, and

Fig. 14. Diffuse plate boundary zones modified from Gordon (1998). Plate abbreviations are AUS, Australia; CL, Caroline; IN, India; PS, Philippine Sea. East and SE Asia are represented as a very large deforming region with rigid parts, such as the North China (NC), Yangtze (Y), Indochina (I) and Borneo (B) blocks.

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shows East Indonesia as a wide deforming zone east of the Makassar Straits. This is consistent with the fact that Sundaland is largely free of seismicity and volcanic activity, which supports proposals of a SE Asian (McCaffrey 1996) or Sundaland (e.g. Simons et al. 2007) plate separate from Eurasia. However, I suggest that the identification of this plate does not imply a rigid strong region, but more a contrast with the strong Pacific and Australian plates, from which the Sundaland interior is separated by poorly defined boundary zones that are characterized by intense seismicity and volcanism. On the whole, the entire region of east Asia and SE Asia is better considered as non-rigid with relatively small strong parts within it, and this has been the situation throughout the Cenozoic. The region between the South China Sea and the Bird’s Head of New Guinea exemplifies the character of the region and highlights some key features of how it is deforming.

Northern Borneo Hall & Morley (2004) drew attention to the weakness of Sundaland during the Cenozoic, recorded by the presence of numerous sedimentary basins, many very deep, and elevated areas. During the Cenozoic most of Borneo north of the Paternoster– Lupar lineaments was a weak deforming region, but southern Borneo and South Sulawesi was a stronger block, or possibly blocks. To the north of the these lineaments are the Rajang –Crocker fold belts, including thick Upper Cretaceous to Eocene, and Eocene to Lower Miocene, deepwater sediments, thick sediments of the Kutei basin, and the deep North Makassar Basin. Most of emergent north Borneo has a Neogene history of contraction and has been the source of the large volumes of sediment filling Neogene basins onshore and offshore, with clear evidence of significant vertical motions relative to sea level, of the order of kilometres, in the last 10 to 15 million years. In contrast, in southern Borneo west of the Meratus Mountains is a broad downwarp, the Barito Basin, filled by Eocene to Miocene terrestrial to marginal marine clastic sediments and shallow marine limestones, whereas to the east is the long-lived Eocene to Miocene Paternoster– Tonasa carbonate platform. With the exception of a narrow zone of deformation in the Meratus Mountains, which may be a reactivated strike-slip suture in the basement, both areas are still largely undeformed. Seismic lines across the Paternoster Platform, and field studies on land, show that Eocene to Recent largely shallow marine carbonates are of the order of 1 –2 km in thickness and record vertical movements relative to sea level of much smaller amounts over 40 million years.

Northern Borneo is the site of important and misunderstood deformation. Haile (1973) first recognized the role of subduction in the history of northern Borneo, and Hamilton (1979) identified the deep NW Borneo –Palawan trough as an extinct subduction trench. Hinz and co-workers (e.g. Hinz & Schlu¨ter 1985; Hinz et al. 1989) disputed this interpretation and argued that the trough was the site of northward thrusting but not subduction, and these alternative views continue to create confusion (e.g. Hutchison 2010). The history of northern Borneo is reviewed and discussed elsewhere (e.g. Hall & Wilson 2000; Hutchison et al. 2000; Hutchison 2005; Hall et al. 2008). Subduction of the proto-South China Sea beneath northern Borneo terminated in the Early Miocene after collision of the extended South China continental margin crust. However, the subduction zone was approximately 150 km south of the present NW Borneo Trough and is now beneath Sabah. The collision resulted in uplift and erosion in the interior of Borneo which provided sediment to the north, east and south. After collision there was a brief period of erosion which formed the Top Crocker Unconformity (van Hattum et al. 2006; Hall et al. 2008) on land and offshore. However, soon after the emergence of much of Sabah, the situation changed again. Although a narrow band of mountains probably remained along the present spine of the Crocker Ranges, the areas to the north and south subsided below sea level and sedimentation resumed. In southern Sabah there was a wide basin SE of the Crocker Ranges (Noad 1998; Balaguru et al. 2003; Balaguru & Nichols 2004). Most of the sediment fed into this basin and carried to the Sulu Sea came from the Borneo interior. River and shallow marine sediments are now preserved in a number of structures described as circular basins, which are remnants of the much larger basin supplied by a large river system, flowing NE, which deposited sand and mud in a delta and coastal plain complex. NW of the Crocker Ranges there was deposition of thick sediments in deltas and coastal plains of north Sabah and Brunei by rivers flowing to the north or NW. In Brunei and offshore Sabah, the position of the shelf edge at different times can be identified (Hazebroek & Tan 1993; Sandal 1996) showing that it moved seaward during the last 15 million years. This indicates that the Crocker Ranges were narrow about 15 million years ago, and have widened gradually with time. In offshore Brunei and NW Sabah Morley et al. (2008) noted that inversion, thrusting and uplift of the present-day onshore area and inner shelf occurred during the Middle Miocene to Pliocene, while a deepwater fold and thrust belt developed during the latest Miocene to Holocene. There was

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a seaward shift of deformation with time (C. K. Morley, pers. comm. 2010) consistent with the movement of the shelf edge. Although some of the deformation can be attributed to shallow gravitational processes in a delta, in places there is more shortening in the deepwater fold and thrust belt than there is extension within the Neogene sedimentary section. On the shelf thick-skinned contractional deformation has episodically affected the Sabah margin from the Middle Miocene to the Pliocene (C. K. Morley, pers. comm. 2010). Therefore a number of authors (e.g. Ingram et al. 2004; Morley et al. 2008; Hesse et al. 2009, 2010; King et al. 2010) interpret deformation, shortening magnitudes, stress orientations, GPS observations (Simons et al. 2007) and recent seismicity to indicate a role for tectonic stresses which they attribute to ongoing convergence of blocks or plates, subduction or inheritance from former subduction. It is noteworthy that almost all the data on which these interpretations are based, with the exception of GPS observations and seismicity (Simons et al. 2007), are from offshore NW Sabah and nearby land areas such as Brunei. Little is published or available for the Sulu Sea side of Sabah, yet on land the structural grain is completely different from west Sabah, changing from the NNE-trending Crocker to the ESE-trending Sulu direction (e.g. Hamilton 1979; Hazebroek & Tan 1993; Tongkul 1991, 1994) and offshore fold axes and thrusts are apparently broadly parallel to the Sulu trend. This is inconsistent with suggested microplates (e.g. Simons et al. 2007) based on GPS observations, and the deformation history and structural trends in southern Sabah (Balaguru 2001; Balaguru et al. 2003; Tongkul & Chang 2003) are equally incompatible. I suggest there is no plate convergence in the NW Borneo region and that deformation is largely a result of topographically-induced stresses and mobility of the deeper crust. NW of the Crocker mountains is a very thick Neogene sediment wedge, including the offshore fold and thrust belt (Fig. 15). Recent studies show thin crust beneath this sediment wedge (Franke et al. 2008; C. Foss, pers. comm. 2008), which requires thinning of crust previously thickened during the Early Miocene collision of the Dangerous Grounds microcontinental block and the Sabah active continental margin. Seismic lines from oil companies (Hutchison 2010) and new data acquired during Malaysian Law of the Sea investigations (V. R. Vijayan, pers. comm. 2008) show elevated features within the NW Borneo Trough at water depths close to 3 km which are capped by carbonates and pinnacle reefs, indicating major subsidence. There is almost no seismicity associated with the trough, no volcanic activity on land, and nothing to indicate southward subduction; nor is there evidence for

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converging plates to produce the fold and thrust belt. On the other hand there is evidence for repeated failures of the shelf edge and slumping into deepwater (e.g. McGilvery & Cook 2003). To the south, on land, there are .2 km high mountains formed of deformed Eocene to Lower Miocene Crocker deep marine sediments, intruded by the 7 –8 Ma Kinabalu granite (Cottam et al. 2010) forming a 4 km mountain, and I suggest these observations indicate a link between significant rapid and young uplift on land, evidenced by exhumation of a 7 –8 Ma granite now exposed at 4 km above sea level, and significant rapid and young subsidence offshore. The trough is thus a flexural depression due to sediment loading and is associated with a flexural bulge in the Dangerous Grounds. In this interpretation the offshore fold and thrust belt, and major shelf failures producing huge deepwater mass transport complexes observed on the sea floor and in the Neogene sequences beneath, are the result of landward normal faulting producing deepwater thrusting. However, although this is not quite the toe-thrust model of the Shell geologists (e.g. Hazebroek & Tan 1993) developed from Niger delta studies, there are important similarities in area and scale (Corredor et al. 2005). In contrast, several authors have argued that the NW Borneo margin is significantly different from the Niger Delta (e.g. Morley et al. 2008; Hesse et al. 2009, 2010; King et al. 2010) based on differences in modern stress patterns, and the observation that in parts of the deepwater fold and thrust belt there is more contraction than extension in the Neogene sedimentary section. I suggest this simply reflects the absence in west Africa of the several kilometres of elevation on land in Sabah, and the balance between contraction v. extension would be found if the section on land were included. In other words, there is no requirement for regional convergence to account for deformation which in any case can explain only deformation of offshore NW Borneo (if that, e.g. Hesse et al. 2009 show that fold orientations are inconsistent with GPS observations) and not that observed in other parts of northern Borneo and offshore. The subsidence and crustal thinning offshore and uplift onshore can be explained by movement of the deeper crust, and this is a phenomenon observed in other parts of the region.

West Sulawesi West Sulawesi has many of the features of northern Borneo except on a larger scale (Fig. 16) but there is even less information in the public domain. On land in the Lariang and Karama areas shallow marine Miocene rocks are overlain by Pliocene coarse clastic sediments derived from an orogenic belt to

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Fig. 15. Profile across the Sabah margin with crustal P wave velocities (m/s) modified from Franke et al. (2008). The section shows the Dangerous Grounds continental crust that was thrust beneath the NW Sabah margin in the Early Miocene and loaded by a thick wedge of sediments that has built out from Sabah during the Neogene; the wedge of sediment now forms an actively deforming fold and thrust belt. A critical point is thinning of the upper and lower crust beneath the thickest part of the wedge. The location of the line is shown on the DEM of satellite gravity-derived bathymetry combined with SRTM topography (Sandwell & Smith 2009). The deepest part of the trough is immediately NW of the 4000 m granite peak of Mt Kinabalu, marked by the black square.

the east (Calvert & Hall 2007). There are mountains up to 3 km above sea level which expose deep crustal rocks such as garnet granulites and eclogites, intruded by young granites, in the Palu area, and probably extensively throughout West Sulawesi (T. van Leeuwen & I. Watkinson, pers. comm. 2009). Rapid uplift and exhumation provided sediment to the broadly west-vergent offshore fold and thrust belt. From north to south the character and orientation of the fold belt changes. The trend of fold axes indicates a radial transport of material away from the mountains which terminate relatively abruptly to the south at the northern edge of the South Sulawesi Tonasa–Tacipi platform where there has been carbonate deposition since the Eocene (e.g. Wilson & Bosence 1996; Ascaria 1997; Ascaria et al. 1997). Like northern Borneo

there was a pre-existing deepwater area into which the fold belt could grow, the Makassar Straits (Hall et al. 2009b), but seismic lines across the northern margin of the Paternoster platform indicate at least 1 km of subsidence of the North Makassar basin on reactivated faults close to Sulawesi at the end of the Miocene. The subsidence is the same age as the rapid exhumation on land. Like north Borneo there is a temporal link, and I suggest a causal link, between subsidence and deformation offshore and uplift and exhumation on land.

North and East Sulawesi West Sulawesi is not the only part of Sulawesi that records rapid subsidence and uplift. This is true for most of North and East Sulawesi, possibly for

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Fig. 16. DEM of the West Sulawesi region with satellite gravity-derived bathymetry combined with SRTM topography (Sandwell & Smith 2009) merged with bathymetry from Puspita et al. (2005). Fold trends in the offshore fold belt are highlighted in yellow and suggest westwards radial vergence. On land areas of (mainly) Late Miocene and younger granites are shown in red. The area north of the NW– SE lineament is suggested to be a weak area with significant deformation whereas south of the lineament the Paternoster Platform and its equivalent in South Sulawesi is a strong area that is almost undeformed.

SE Sulawesi, and certainly for the major enigmatic inter-arm basins of Gorontalo Bay and Bone Gulf. The north and east arms of Sulawesi are striking in their exceptional elevations (up to 3 km) within short distances of the coast, and the narrow width of these emergent areas. Dating uplift and exhumation is only just beginning. K –Ar and Ar– Ar cooling ages of 23 –11 Ma from micas and hornblende are reported by van Leeuwen et al. (2007) from the Malino Complex at the west end of the north arm, which they interpret as a core complex. The ages appear to fall into two groups and although there is only a small number (23 ages from 6 samples). I speculate that the older ages record Early Miocene collision of the Sula Spur, and the cluster of ages from 14–11 Ma record rollbackinduced extension. Throughout West, North and East Sulawesi there is evidence for significant vertical motions on land at about 5 Ma, recorded by K – Ar and apatite fission track ages from granites (e.g. Bergman et al. 1996; Elburg et al. 2003; Bellier

et al. 2006), and by widespread and thick Celebes Molasse deposits which indicate rapid exhumation from about 5 Ma (e.g. Calvert 2000; van Leeuwen & Muhardjo 2005). The term Celebes Molasse is used for a variety of Neogene terrestrial or shallow marine deposits found throughout Sulawesi, but although it may include Lower Miocene postophiolite detrital sediments (e.g. Surono 1995; Surono & Sukarna 1996) there was clearly a major increase in output of clastic sediment in the Latest Miocene and/or Early Pliocene in West and East Sulawesi (e.g. van Bemmelen 1949; Garrard et al. 1988; Davies 1990; Calvert 2000; Calvert & Hall 2007). Offshore, recently acquired seismic and multibeam data in Gorontalo Bay show spectacular subsidence recorded by numerous pinnacle reefs now found within a range of water depths between 1 and 2 km many of which, despite the high rates of sediment supply, are not buried by sediment. They indicate very young and rapid subsidence. For

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example, between the north coast of the east arm and the Togian Islands there is a 50  20 km lobe of sediment with a thickness of up to 2 s TWT (two-way travel time) (Jablonski et al. 2007), that links alluvial fan deposits on land at north and south ends of the lobe. The water depths in the area between the east arm and the Togian Islands are now up to 1.5 km; 25 km to the south of the east arm coast elevations exceed 2 km. East of Poh Head at the end of the east arm are probable platform carbonates with no sediment cover at water depths of 1 km, and possible carbonates still deeper beneath bedded sediments (Ferdian et al. 2010; Watkinson et al. 2011). The platform carbonates are likely to be Middle and/or Upper Miocene by comparison with limestones beneath the Celebes Molasse in the Togian Islands, implying subsidence of 1 km or more in probably less than 5 Ma.

North Moluccas A striking feature of East Indonesia is that, although there are undoubtedly some small plates, when plate boundaries are revealed by seismicity (Engdahl et al. 1998) they terminate abruptly and cannot be traced into other boundaries, implying relatively rapid changes from subduction to distributed deformation of the same plate (Fig. 17). Much of the deformation recorded by the upper crust seems to

be almost independent of the deeper lithosphere. The Philippine Trench terminates at about 38N; to the north the slab has been subducted to at least 100 km, dips steeply and becomes almost vertical, and there is a deep trench. This plate boundary neither continues southeastwards as often shown, nor does it connect via obvious faults to the Molucca Sea. The abrupt termination of the trench implies considerable distributed deformation in the north Halmahera area. In the Molucca Sea seismicity shows the wellknown double subduction system (Silver & Moore 1978) clearly indicated by seismicity, tomography and volcanic activity. There is no trench associated with either the west or east-dipping slabs, and it appears that the ‘melange wedge’ of the central Molucca Sea (McCaffrey et al. 1980) is deforming independently of the subducted Molucca Sea plate beneath. The southern edge of the west-dipping Molucca Sea slab terminates abruptly beneath Gorontalo Bay and runs almost due east –west. There is no surface expression, although this is not to be expected in Gorontalo Bay which forms the upper plate. However, surprisingly the lineament that would be expected to mark the former southern boundary of the Molucca Sea Plate, named the North Sula Sorong Fault (Hamilton 1979), which should be a major left-lateral strike-slip fault, has no expression on the sea bed, and has no seismicity

Fig. 17. Earthquake hypocentres in East Indonesian from Engdahl et al. (1998) and the Global CMT database (CMT Project 2009). Hypocentres that can be identified with different plate boundaries are shown in colour chosen to match Figure 19. At the mantle scale these imply a number of sharply demarcated and distinct plates, but at the surface many of the plate boundaries are not connected, and the Molucca Sea Plate is completely subducted with no surface expression, being overridden by the two converging fore-arcs which are deformed into the central Molucca Sea wedge.

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in the crust. There are a very small number of hypocentres below 30 km, a few of which could indicate left-lateral displacement on a broadly-east –west fault system, but at the sea bed the most obvious structures in the plate boundary region are southward-directed thrusts. Even these do not have the displacement expected. Silver et al. (1983) suggested that thrusting was related to southward gravitydriven movement of the Molucca Sea wedge, and this is plausible north of the Sula Islands. However, close to Poh Head the thrusting occurs at the southern termination of strands in a right-lateral fault system which can be traced east from land (Simandjuntak 1986; Ferdian et al. 2010; Watkinson et al. 2011). The existence and displacement on this right-lateral fault system casts doubt on the connection between the fold belt in the east arm south of Poh Head (an Early Miocene structure), the interpreted Batui Thrust crossing Poh Head on land, and the Sangihe subduction offshore. Some of the faults shown on maps of the region do exist but are old structures (Fig. 18) that no longer have surface expression (e.g. the Molucca Sea Trenches). Some plate boundaries do not connect to others and require distributed deformation of the lithosphere to maintain an internally consistent plate model (e.g. the Philippine Trench). Some of the faults probably do not exist

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or are not connected in the ways shown on most maps, for example, the NW–SE Greyhound Fault does not cross the seabed at all, and its existence is doubtful. The improvement in the quality of remotely sensed data (notably SRTM and Aster imagery for the rain forest areas of eastern Indonesia), and new seismic and multibeam data, means that structures can be mapped with more confidence, but there are still many areas of uncertainty. Nonetheless, it is certain that if the crust in the Sulawesi –North Moluccas region is broken into a series of blocks the boundaries do not correspond to those of the known plates. In fact, it is very unlikely that the upper crust is deforming as a series of rigid blocks, and some of the relative movements of the upper crust are not those predicted by our current plate models (Fig. 19).

Deformation If the region is not deforming as plates or microplates, nor reflecting microplate movements, how is it responding to movements of the large plates? The common features of the region between offshore northern Borneo and the North Moluccas are rapid uplift of land to elevations of up to 3 km (and locally higher in the case of Mt Kinabalu at 4 km), and rapid subsidence offshore with water

Fig. 18. GPS velocities and interpreted blocks in East Indonesia modified from Socquet et al. (2006). Different colours of GPS vectors show the different blocks outlined by the faults shown in black lines and are coloured to match Figure 19. Faults that penetrate the lithosphere are shown with heavy black lines. Many of the interpreted block boundaries are upper crustal faults of uncertain character and age; dashed lines are inactive, non-existent or very doubtful. The upper crust is clearly deforming in a complex way not directly related to plates that can be identified from seismicity (compare with Fig. 17).

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Fig. 19. 3D cartoon of plate boundaries in the Molucca Sea region modified from Hall et al. (1995). Although seismicity identifies a number of plates there are no continuous boundaries, and the Cotobato, North Sulawesi and Philippine Trenches are all intraplate features. The apparent distinction between different crust types, such as Australian continental crust and oceanic crust of the Philippine and Molucca Sea, is partly a boundary inactive since the Early Miocene (east Sulawesi) and partly a younger but now probably inactive boundary of the Sorong Fault. The upper crust of this entire region is deforming in a much more continuous way than suggested by this cartoon.

depths of more than 2 km and several kilometres of sediment below the seabed. The term uplift is now used with reluctance because it is imprecise in specifying what has moved, and it is difficult to quantify amounts and rates (England & Molnar 1990). However, it is very probable that what is now the land surface in this region was very close to sea level only a few million years ago and that slopes were less steep. Areas now offshore also had less relief and were close to sea level, demonstrated for example, by the distribution of carbonates. Material that was at the top of the crust on land several million years ago was eroded, transported offshore and deposited in sediment layers, above carbonates deposited close to sea level, which are in many places now at depths of several kilometres below sea level. What is now at the surface on land at elevations of up to 3 km in Sulawesi was, a few million years ago, more than 3 km below the land surface and has been exhumed by high rates of erosion. On land we generally lack the means to be precise about the amounts of uplift of the land surface or uplift of rocks within the crust, whereas in offshore regions hydrocarbon exploration drilling

can provide accurate dating, but at present in the frontier regions of eastern Indonesia ages are generally lacking. Different parts of Borneo and east Indonesia have risen and subsided since the Early Miocene following Australia’s initial contact with SE Asia. I make the assumption, based on the best but limited evidence currently available, that very significant change in relief began at about 8 Ma in northern Borneo, and at about 5 Ma in West, North and East Sulawesi. In the Halmahera islands, and probably in Seram, significant relief changes are even younger. From the area of West, North and East Sulawesi around Gorontalo Bay it is possible to identify a number of features that are typical of the much larger region of eastern Indonesia. Uplift and subsidence are intimately interlinked in time. Uplift has been maintained despite high rates of erosion, implying that the forces causing the uplift continue to act. Sometimes high erosion rates are explained away by suggesting very weak rocks at surface – this is a possible explanation in Sabah, although many of the rocks now exposed do not seem to be

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unusually weak, but the Celebes Molasse contains abundant fragments derived from igneous rocks of the East Sulawesi ophiolite implying erosion of strong rocks. The rates of subsidence are greater than would be expected from purely thermal processes but are similar to those driven by extension in some parts of the world. However, most tectonic models suggest this region is being deformed by converging plates where contraction would be expected. There are faults, but seismicity indicates that few if any are lithosphere-scale structures. As mentioned above, several lithospheric faults have no surface expression, and of those that do only the North Sulawesi subduction zone and the Palu-Koro Fault may cut the entire crust. It is not surprising that there are faults in the upper crust, but there are no obvious fault-bounded blocks, and even if there were this would imply that they were merely upper crust features, disconnected from the deeper lithosphere. The deformation appears largely independent of plate boundaries, and the distribution, amounts and rates of vertical movements appear to be far greater than expected from conventional models of stretching or from other mechanisms such as strike-slip faulting. The region of Sulawesi around Gorontalo Bay has been deformed into elevated north and east arms and subsided central bay, and there is also a lower elevated ridge running roughly east –west through the Togian Islands. Considering a north– south section across Gorontalo Bay the scale of the deformed region can be approximated as a curve with wavelength of 200 km and amplitude 5 km. These amounts are small for deformed plates, where wavelengths are much larger and amplitudes smaller, and also appear unusual for deformed continental crust, based on comparison with other parts of the world, whether in extension or contraction, although features on this scale could be produced under certain conditions by compression (D. A. Waltham, pers. comm. 2009). GPS studies show that there are very high rates of movements (e.g. Walpersdorf et al. 1998; Vigny et al. 2002; Socquet et al. 2006; Simons et al. 2007), comparable to estimates of plate movements in the region (Fig. 18), and they also show what is recorded by geological observations: relatively abrupt changes from strongly deformed to little deformed areas, although it is not clear if the boundaries between these areas are narrow or wide. Since GPS measurements cover a period of only a few years, stations are very scattered, and some may be poorly sited, it is not clear what these results mean. Attempts have been made to outline fault-bounded blocks that may explain GPS measurements, but the blocks demarcated are bounded by structures of different ages, some

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without surface expression, and some are simply required by a block model but are surmised with no evidence.

An alternative model I suggest all these observations indicate that the region (Fig. 20) bounded by the strong parts of the Australian and Pacific –Philippine Sea Plates is not a plate, in the sense of a large rigid entity of the plate tectonic paradigm, but rather is a large region of generally weak lithosphere responding to movement of the rigid plates that surround it. The region is thousands of kilometres across. It deforms internally as the forces acting at the boundaries change in direction and magnitude. Its response to the external forces is modified by the distribution of strong areas within it which are more or less rigid. Some of these subduct, such as the Celebes Sea, and the Banda embayment, inducing deformation, and others do not, such as the strong old continental fragment(s) that form East Java, SE Borneo and South Sulawesi. These latter transmit, refract, and focus deformation. Australia collision with North Sulawesi in the Early Miocene caused some contraction, with emplacement of ophiolites in Sulawesi, and broader consequences throughout Sundaland, such as local inversion, which probably reflect the changing balance of forces on the entire region. Several new subduction zones were initiated in the Neogene, such as those at the NW, NE and southern edges of the Celebes Sea, and the Java Trench propagated into the Banda embayment (Spakman & Hall 2010). It is not clear how subduction was initiated, and whether there was a period of compression before the plate broke, although in the Banda embayment it is likely that a pre-existing lithospheric fracture simply broke along an older boundary when the opportunity became available. The deformation induced by subduction was predominantly extensional, was concentrated in the upper plate as the Banda subduction hinge rolled back, and was modified by several factors including the heterogeneous nature of the upper plate, magmatism caused by decompressional melting, and melting induced by fluid movements into the mantle wedge above the subducted slab. Subduction rollback at the North Sulawesi Trench has caused additional extension and contributed to subsidence in Gorontalo Bay and exhumation of metamorphic core complexes on land in Sulawesi (Spencer 2010). The response to rollback was predominantly subsidence in the weaker parts of the region, but stronger areas, such as SE Borneo –South Sulawesi, were almost unaffected by nearby deformation and have remained at a similar elevation relative to sea level. The stronger areas have also acted to transmit

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Fig. 20. Deformation of East and SE Asia. The region is surrounded by plates that include strong continental and oceanic lithosphere but without completely connected boundaries shown as solid black lines. The white area is a deforming region that cannot be easily assigned to particular plates. GPS vectors from Shen et al. (2005) and Simons et al. (2007) with green arrows show the upper crust is deforming in a complex way not directly connected to plate movements. SE Asia is a largely weak area with strong oceanic and continental parts.

compressional forces from the plate boundaries, and this has caused uplift in adjacent weak areas such as the Central Borneo Mountain Ranges, West Sulawesi Mountains, and North and East Sulawesi. Subsidence and uplift in weak areas were temporally linked. Morley & Westaway (2006) have shown that unusually rapid and large amounts of subsidence in basins of SE Asia can be explained by deep crustal flow with thinning of the crust beneath the basins; they argued that in these settings sediment loading can cause, or contribute significantly to, subsidence. They suggested that the deep crust moved away from the depocentre, thinning the crust, and flowed towards the sediment source regions where the crust was thickened. They focused their attention on the role of crustal flow in producing deep sedimentary basins and not on the regions towards which lower crust was

flowing, except as a source for sediment. The basins they studied were formed and largely filled since the Eocene and are now observed at a relatively late stage in their development. However, the Morley & Westaway (2006) model offers an explanation for the link between subsidence and uplift from northern Borneo eastwards at a much earlier stage in basin evolution. I suggest that deep crustal movements not only enhanced subsidence and provided sediment sources, but flowed laterally into areas already elevated and drove significant further uplift. This flow maintained relief, erosion, and provided sediment that drove further subsidence in adjacent basins. In particular, it has contributed to the formation of very high mountains in the last 5 Ma. The high short-term sediment yields in SE Asia are commonly attributed to extreme local relief plus high rainfall, with exceptionally large

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Fig. 21. Cartoon illustrating the concept of deep crustal flow linking basin subsidence and mountain elevation. Deformation of an originally irregular crust is assumed to be initiated by stresses transmitted from stronger rigid plates outside the area, and the weaker areas are most responsive. Flow of the deep crust in weaker areas then drives subsidence of basins and elevation of mountains to lead to formation of deepwater fold and thrust belts such as those of northern Borneo and West Sulawesi. Crustal flow may also enhance extension driven by other processes such as subduction rollback, or uplift resulting from delamination, both of which could be contributing to deformation in Borneo and Sulawesi.

sediment volumes carried by short mountainous rivers (e.g. Milliman & Syvitski 1992; Milliman et al. 1999). However, in many parts of SE Asia these high sediment yields have been maintained for tens of millions of years (Hall & Nichols 2002; Hall & Morley 2004) which requires a tectonic mechanism to provide relief. Deep crustal flow is a solution to this problem. Once there is movement of the deep crust a feedback process begins. Furthermore, the erosional supply may be enhanced by gravitational movement of the upper crust with detachments at depths up to several kilometres. Figure 21 shows in a cartoon form the process envisaged. Major uplift and high exhumation rates on land will promote shelf failures and produce offshore fold and thrust belts such as those documented from offshore northern Borneo and offshore West Sulawesi, similar in style and dimensions to the Niger delta, but differing in their position adjacent to significantly elevated mountains. These fold and thrust belts are largely aseismic as they are not the result of converging plates. One implication of the model is that structures in many SE Asian mountain belts previously interpreted as thrusts are actually major extensional detachments, a now relatively uncontroversial view in more accessible regions without thick rainforest vegetation cover (e.g. Coney 1980; Coney & Harms 1984; Lister et al. 1984; Lister & Davis 1989).

This model leaves unanswered the questions of why deep crustal flow starts, and why in some areas rather than others. Ultimately, it could be initiated by changes such as regional plate reorganizations, initiation of subduction and rollback, or even climate change. Areas unaffected are strong, and include relatively old oceanic lithosphere (e.g. the Celebes Sea) or areas underlain by thick old continental lithosphere (e.g. East Java –West Sulawesi). Areas that are affected are weak, which may be the result of processes such as heating associated with long term subduction, magmatism, or loss of deeper lithosphere by delamination.

Conclusions SE Asia is an unusual region. In eastern Indonesia there have been exceptionally high rates of vertical movements and rapid but varied horizontal movements that are not explicable as movements of small rigid micro-blocks, nor easily described in terms of plate tectonics. Plate tectonics provides the first order description of the region’s history, but to understand it more completely we must view it as an extensive region of very weak lithosphere, probably most of East and SE Asia, with a heterogeneous basement structure, within which are strong areas of old continental lithosphere and

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oceanic crust, all deforming in response to the changing balance of forces at boundaries with the strong surrounding plates of the Pacific, Philippine Sea, Australia and India. Seismicity shows a relatively simple plate tectonic setting with convergence of the Pacific, Philippine Sea, Australia and India plates towards SE Asia, but GPS motions record a much more complex pattern of deformation. On the whole this should not be viewed as reflecting small block movements but as a continuum of deformation in the upper crust which is partly or completely decoupled from the deeper lithosphere and the response of the upper crust to local stresses, such as those induced by topography. The location of uplift and subsidence are largely independent of plate boundaries. Hall & Morley (2004) suggested that a continuum model may offer a better description than a rigid block model and that deformation might be understood by finite element modelling of stresses originating from all the plate boundaries surrounding Sundaland combined with those from topography. However, even this approach is likely to be unsuccessful because of the heterogeneity of the ‘plate’ and because of the decoupling of upper crust and lithosphere. For most of the Cenozoic the strong areas of old continental lithosphere have remained little deformed, and several areas (SE Borneo, South Sulawesi, Banggai-Sula) contain broadly flat lying sediments deposited close to sea level. In the weaker areas are thick sedimentary successions in deep basins, commonly adjacent to deeply exhumed elevated areas, which have subsided at high rates as the adjacent mountains have risen, been deeply exhumed and supplied sediments to the basins. Rates are greater than those predicted by conventional stretching models and isostatic responses. Several features suggest that there is a connection between subsidence and uplift caused by crustal flow at depth, away from the sedimentary basins, thus causing or contributing to subsidence, and towards the mountains, promoting further uplift and maintaining sedimentation. This is different from the India collision zone where deep crust is flowing away from the thickened area (e.g. Royden et al. 1997; Clark & Royden 2000; Shen et al. 2005). In East Indonesia positive feedback has maintained subsidence, uplift, and fold and thrust deformation for more than ten million years. It explains the paradox of the high sediment yields from small land masses in SE Asia that are maintained for long periods. The magnitude and rates of vertical motions raise many questions for tectonic studies, hydrocarbon and mineral exploration, and changes in palaeogeography and their implications for life sciences, which remain to be explored. Is eastern

Indonesia an analogue for earlier Cenozoic deformation of Sundaland west of the Makassar Straits? Is this region an analogue for other orogenic belts? What initiated the most recent changes in the last 10 Ma? Work by the SE Asia Research Group has been funded over many years by a consortium of oil companies who have been generous and open-minded in their support. Several of these companies have shown me offshore data which stimulated the interpretations in this paper. Our work has also been funded at times by the University of London Central Research Fund, the Natural Environment Research Council, and the Royal Society. I thank Pusat Survei Geologi Bandung, Lemigas, Indonesian Institute of Sciences, and Institut Teknologi Bandung, for assistance and many colleagues, friends and students in the UK, Europe and SE Asia for help and discussion. I thank Clive Foss for showing me evidence of crustal thinning beneath sediments from offshore Sabah. Alfred Kro¨ner and Chris Morley provided very helpful reviews and Chris Morley made some challenging comments, notably concerning northern Borneo, for which I am grateful. I thank Jim Granath and ION-GX Technology Imaging Solutions for the deep Java Sea image of Figure 11.

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The Java convergent margin: structure, seismogenesis and subduction processes HEIDRUN KOPP IFM-GEOMAR Leibniz-Institute of Marine Sciences, Wischhofstr. 1-3, 24148 Kiel, Germany (e-mail: [email protected]) Abstract: The Java margin is characterized by a distinct variation in lower to upper plate material transfer and recurring catastrophic tsunamogenic earthquakes. Both processes are closely linked to the subduction of oceanic basement relief resulting in varying degrees of fore-arc deformation. Tomographic models of refraction seismic profiles and reflection seismic lines in combination with high-resolution multibeam bathymetric data reveal the variability in the deep structure and deformation of the Java fore-arc. Shallow subduction processes are governed by the sediment supply in the trench as well as by the nature and fabric of the oceanic lithosphere. The deep structure of the fore-arc reveals a shallow upper plate crust–mantle transition, present along the entire Java margin section. The serpentinized fore-arc mantle wedge governs the depth extent of the seismogenic zone here, which is narrower compared to its Sumatran analogue. In addition, offshore central Java, high relief oceanic basement features potentially act as asperities as well as barriers to seismic rupture, limiting the possible magnitude of subduction thrust earthquakes. However, the potential for geohazards, in particular tsunamis, is high along the entire margin. This results from tsunamogenic earthquakes, ubiquitous splay faults and potentially tsunamogenic landslides, which further increase the risk of future tsunamis.

This contribution reviews studies of the Java subduction zone (Fig. 1) with a special emphasis on the neotectonic processes and structural evolution of the fore-arc using tomographic images based on seismic refraction data (Fig. 2) and multibeam bathymetry (Fig. 3). Following introductory remarks on the tectonic setting, seismicity and geophysical data base, the observational and interpretative information gained from these studies will be presented area-specific for each tectonic segment along the margin. Going from west to east, these first-order segments are defined based on the dominating material transfer and nature of the lower plate, that is sediment accretion offshore Sunda Strait and western Java, fore-arc erosion offshore central-eastern Java, and continent–island arc collision offshore Sumba Island. A regional summary is presented in the conclusions.

Tectonic setting The Java trench forms the eastern section of the Sunda deep-sea trench and is the site where the Indo-Australian oceanic lithosphere subducts beneath the Sundaland block of Indonesia. It straddles the island of Java and the Lesser Sunda islands of Bali, Lombok and Sumbawa (Fig. 1). The Java margin is bound in the west by the transtensional regime of the southern Sunda Strait (Lelgemann et al. 2000). The Sunda Strait marks the hinge line and passage from the trench-perpendicular

convergence off Java to the oblique subduction offshore Sumatra. In the east, the Java trench terminates at 1218E where it merges into the Timor trough (Audley-Charles 1975, 2004; Hall & Smyth 2008). The transition from oceanic subduction along the Java trench to continent– island arc collision along the Timor trough of the Banda arc occurs south of Sumba Island, where continental crust is colliding with the fore-arc (Hamilton 1979) (Fig. 1). The Indo-Australian plate currently moves at 6.7 cm/a in a direction N118E off western Java and thus almost normal to the trench (Tregoning et al. 1994) (Fig. 1). Convergence speed slightly increases from western Java towards the east, however, at a very subtle rate, reaching c. 7 cm/a off Bali (Simons et al. 2007) and has been active since the Eocene (Hall & Smyth 2008). A local source tomography based on recordings over a period of almost six months from more than 100 seismic stations on Java images the steep dip angle of the subducting slab, reaching c. 708 underneath the island (Koulakov et al. 2007). The age of the incoming plate increases from Late Cretaceous offshore Sunda Strait (1058E) to Late Jurassic at 1208E. Accordingly, water depth in the trench increases from c. 6 km off the Sunda Strait to more than 7 km off Lombok and Sumbawa in the east and correlates with the decrease in sediment supply from c. 1500 m off western Java to the starved trench segments along the eastern Java trench.

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 111–137. DOI: 10.1144/SP355.6 0305-8719/11/$15.00 # The Geological Society of London 2011.

112 H. KOPP Fig. 1. Morphology of the Java margin based on satellite altimetry data (Smith & Sandwell 1997). A large bivergent accretionary wedge is expressed as a continuous bathymetric ridge fronting the Java fore-arc basin offshore western Java. This ridge structure is broken and highly deformed offshore central Java, where the oceanic Roo Rise is colliding with the margin. The eastern Java trench offshore Bali to Sumba is characterized by the subduction of smooth oceanic crust of the Argo Abyssal Plain. The transition from oceanic subduction to continent– island arc collision occurs south of Sumba where the Scott plateau enters the trench. Black lines show wide-angle refraction profiles. Black dots show locations of ocean bottom instruments displayed in Figures 4, 5, 8, 9 and 12. White lines show extent of MCS data shown in Figures 6, 7 and 11. Stars denote earthquake hypocentres of 1977 (white), 1994 (yellow) and 2006 (red). C.-J.: Ciletuh-Jampang block.

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Fig. 2. Tomographic images and velocity– depth distribution along seven refraction seismic dip lines crossing the fore-arc between western Java and east of Sumba island. The profiles document the variation from the accretionary domain (a and b) to the erosional seamount/plateau subduction regime off central to eastern Java (c and d). To the east, the transition from oceanic subduction offshore Lombok (e) to continent –island arc collision (f and g) occurs. All profiles west of Sumba show a shallow hydrated upper plate mantle, which limits the downdip extent of the seismogenic zone. Profiles are approximately aligned along the vertical stippled line. Vertical exaggeration in all profiles is 2.5.

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Fig. 3. High-resolution sea floor bathymetric mapping along the Java margin. The upper panel shows the grid of swath data underlain by global bathymetry; the lower panel displays the high resolution data coverage. Full coverage was achieved along the deformation front and lower trench offshore Lombok, Sumbawa and Sumba. RV SONNE’s EM120 multibeam echosounder system sends successive frequency-coded acoustic signals. Data acquisition is based on successive emission –reception cycles of this signal. The reception is obtained from 191 beams, with the footprint of a single beam of a dimension of 28 by 28. Achievable swath width on a flat bottom is up to six times the water depth dependent on the character of the sea floor.

The Java margin exhibits two prominent features on the incoming Indo-Australian plate: (1) the Christmas Island Seamount Province including the Roo Rise and (2) the Argo Abyssal Plain (Fig. 1). The Christmas Island Seamount Province forms a

broad, irregular topographic swell of 135–140 Ma old oceanic lithosphere off central Java (Moore et al. 1980; Mueller et al. 1997) and is dotted with numerous seamounts. It extends in an east –west direction from 958E to 1158E, where it terminates

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abruptly. The evolution of the Christmas Island Seamount Province remains enigmatic and is associated with a series of individual magmatic events. Rock sampling resolved varying formation times across the volcanic province without a clear formation time trend (Werner et al. 2009). The eastern segment of the Christmas Island Seamount Province features the oceanic Roo Rise (Fig. 1), a 400 km wide plateau rising 2–3 km above the abyssal plain. Dredged rock samples from the plateau retrieved strongly altered olivine phyric lava fragments with Mn-crusts (Werner et al. 2009). The northern flank of the Roo Rise is currently entering the trench south of eastern Java. Isolated volcanic summits on the plateau and adjacent to it represent high relief gradient features which upon subduction cause frontal erosion of the fore-arc south of central Java (Masson et al. 1990). A tomographic study of passive and active seismic data images the complex crustal structure of the fore-arc, segmented into distinct blocks (Wagner et al. 2007). The second prominent feature in the oceanic domain is the Argo Abyssal Plain, with mean water depth around 5500 m (Fig. 1) and a crustal age of 160 Ma (Mueller et al. 1997). Though large-scale topographic features are not observed on the Argo Abyssal Plain, the sea floor nonetheless exhibits a distinct structure, comprising the original spreading fabric and a pervasive pattern of trench parallel normal faults where the plate bends into the trench (Masson 1991). The horst-and-graben structures on the outer trench wall show a maximum throw of 500 m. Individual fault segments reach lengths of more than 60 km and cut deep into the oceanic basement (Lueschen et al. 2010). 530 m of sediment on the Argo Abyssal Plain have been drilled in DSDP site 261 before reaching Late Jurassic oceanic basement (Heirtzler et al. 1974) on which Cretaceous claystones, Upper Miocene and Pliocene nannofossil oozes and Quaternary radiolarian clays have been deposited. In the east, the Argo Abyssal Plain is bordered by a continental promontory of the Australian lithosphere: the Scott Plateau (Fig. 1) governs subduction zone processes at the transition from the Sunda margin to the Banda arc.

Recent earthquake activity The Java trench was the site of three major earthquakes (Mw  7.0) in the past decades (Fig. 1). In 1977, a Mw ¼ 8.3 normal fault earthquake ruptured the underthrusting plate (Spence 1986; Lynnes & Lay 1988) and caused a tsunami on the Lesser Sunda islands and Australia with maximum run-up heights of 8 m on Sumbawa (Kato & Tsuji 1995). The Mw ¼ 7.8 event of 1994 was a megathrust rupture, which occurred offshore eastern Java

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(Ammon et al. 2006; Bilek & Engdahl 2007). The triggered tsunami reached run-up amplitudes of 5 –8 m along the southern coast of Java, causing more than 700 casualties. Similarly, in 2006, a Mw ¼ 7.7 megathrust earthquake offshore western Java (Fig. 1) triggered a deadly tsunami with maximum run-up heights of 20 m (Fritz et al. 2007). Both thrust events show a rupture pattern characteristic of tsunami earthquakes (Kanamori 1972; Bilek & Engdahl 2007) and both events display aftershock sequences dominated by normal faulting, suggesting relatively complete stress release on the interplate thrust (Ammon et al. 2006). The influence of the plate structure at depth, particularly regarding subducted basement features, has been discussed for the 1994 event, for which slip over a subducted seamount was proposed to have triggered the earthquake (Abercrombie et al. 2001). This model is still debated for the 2006 event (Bilek & Engdahl 2007).

Geophysical data Three major marine experiments were conducted on the Java margin since 1997 using the German RV SONNE as platform. A total of 289 ocean bottom stations [hydrophones and seismometers: OBH/S (Bialas & Flueh 1999)] were deployed along 13 transects. In 1997/1998, the GINCO cruises focused on the transition zone between frontal and oblique subduction off the Sunda Strait and acquired three coincident seismic refraction/reflection profiles offshore western Java (P05, P06, P07) (Kopp et al. 2001; Schlueter et al. 2002). It was followed by the MERAMEX cruises off central Java in 2004 during which two long refraction dip lines (P16 and P18) were shot from the oceanic lithosphere to the continental slope in addition to a coast-parallel strike-line (P19) (Fig. 1). During this onshore– offshore or amphibious project, 100 landstations were deployed onshore Java for a period of 120 days (Koulakov et al. 2007; Wagner et al. 2007). The eastern section of the margin was subsequently investigated during the SINDBAD experiment in 2006 and was covered by four coincident refraction/reflection dip lines (P42, P31, P22 and P11) and three strike lines (P41, P33, P21) located between offshore eastern Java and east of Sumba island, including the transition from oceanic subduction to continent–island arc collision. Figure 2 displays the velocity –depth models of all acquired trench perpendicular refraction dip lines and serves as a basis for the following discussion of individual margin segments. All models were achieved by tomographic inversion of the refraction data (Shulgin et al. 2009, 2010; Planert

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et al. 2010; Wittwer 2010), except for the GINCO profile for which forward modelling by raytracing for the broadly spaced instruments was conducted (Kopp et al. 2001). Figure 4 shows an ocean bottom seismometer recording displaying typical phases from stations near trench locations on the upper plate: Pn is the refraction through the upper mantle, PmP is the reflection from the crust – mantle boundary, Poc is the refraction through the oceanic crust, PtocP is the reflection from the top of the oceanic crust and Psed denotes sediment phases. Pg fore-arc is the refraction through the fore-arc crust. The information contained in the seismic record sections thus returns velocity data for the different margin units from the sedimentary cover to the upper mantle. The bathymetry data (Fig. 3) were acquired during the MERAMEX and SINDBAD experiments using RV SONNE’s Simrad multibeam echosounder system, which provides accurate depth measurements and bathymetric mapping in areas at depths down to 11 000 m.

Offshore Sunda Strait and western Java: sediment accretion Observations Along the western margin segment (longitudes 105 to 1108E), the incoming oceanic crust is 7.5 –8.5 km thick with velocities typical of mature oceanic crust, increasing from 4.7 km/s at the basement to 7.2 km/s at the crust –mantle boundary. Upper mantle velocities are in the order of 7.8–8.0 km/s as documented by Pn mantle phases recorded by the ocean bottom seismometers (e.g. station OBS 42 on line P18 in Fig. 5). The morphology of the offshore fore-arc of western Java is dominated by a massive fore-arc high above the underthrusting plate (Fig. 2, panels a, b). The tectonic features observed here include the sediment-filled deep-sea trench, an actively accreting prism and mature fore-arc basin. They are morphological manifestations of the accretionary regime continuous to offshore Sumatra (Kopp et al.

Fig. 4. Data example of an ocean bottom hydrophone (OBH) located on the lower slope off central Java on profile P16. Phase nomenclature is as follows: Pg refraction through fore-arc crust, Poc refraction through oceanic crust, Pn refraction through upper mantle, PtocP reflection from oceanic basement, PmP Moho reflection.

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2008), where the fore-arc high is subaerially exposed on the fore-arc islands. Offshore Sunda Strait, the topographic expression of the fore-arc high is subdued as a result of the Neogene transtensional tectonics and the increasing trench-parallel slip component towards Sumatra (Huchon & Le Pichon 1984; Malod et al. 1995). Malod & Kemal (1996) estimated about 50–70 km of extension in the Sunda Strait during the Pliocene (Diament et al. 1992). The internal architecture of the fore-arc is characterized by multiple kinematic boundaries between the trench and Java continental slope. The deformation front marks the transition from the trench to the frontal prism which then transitions into the Neogene accretionary prism, that rapidly increases in thickness. In Figure 6, an c. 0.5 km thick sheet of trench sediment is underthrust below the frontal prism. A similar decollement zone is also observed on profile P05 (Kopp et al. 2009), though discontinuous due to subducted seamounts of 1 –2 km height attached to the oceanic crust (Fig. 7). This distinct lower plate relief, which is ubiquitous along the Java margin (Masson et al. 1990), potentially breaches the subduction channel and is in contact with the upper plate. A pronounced megasplay or backstop thrust separates the Neogene prism from the Palaeogene prism (Figs 6 & 7). Seismic velocities increase from the Neogene prism to the older, more consolidated material of the Palaeogene prism. Station OBH 47 on profile P18 (Fig. 8) documents this increase in seismic velocities in the Pfp and Pg phases. The Palaeogene prism is covered by a slope apron with seismic velocities not exceeding 2.5 km/s (Fig. 2, panels a, b). This unit displays little permanent deformation. During the GINCO cruise, Pliocene to recent sediments were sampled at various locations in the fore-arc basin and on the fore-arc high (Beiersdorf 1999). Samples retrieved hemipelagic muds with intercalated turbidites and dacitic to rhyolithic ash layers. Dredge samples from outcrops of the accretionary prism revealed that they consist of silty and micaceous mud and tectonized mudstone as well as of arenitic limestone and calcareous sandstone (Beiersdorf 1999). Unfortunately, rock dredge or drill samples have not been recovered from the core of the fore-arc high, so that its composition must be inferred from seismic velocities. Offshore western Java, the Palaeogene prism fronts the Java fore-arc basin (Fig. 7). The Java basin is expressed as an elongated, 500 km long subsiding belt with an average water depth of 3500 m. Sediment thickness of the basin infill reaches 4 km (line GINCO p05, Fig. 2), decreasing towards the basin fringe (line Meramex P18, Fig. 2). The basin is underlain by a unit characterized by seismic

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velocities rapidly increasing from 5.5 km/s to values larger than 6 km/s (Fig. 2). OBH 53 is situated on the northern rim of the fore-arc high on line P18 and records phases through the Palaeogene prism as well as through the fore-arc crust and mantle (Fig. 9). A number of stations record the fore-arc basement, which shows an ophiolitic character. The basement is exposed in western Java, where outcrops of peridotites, gabbros, pillow basalts and serpentinites are observed (Sukamto 1975; Schiller et al. 1991; Susilohadi et al. 2005). This Cretaceous –Paleocene complex is imaged underneath the fore-arc basin as a seaward dipping unit with the crust –mantle boundary located at c. 15 km depth (Figs 2 (panels a, b & 7). It is bounded to the NW by the Cimandiri fault zone, which cuts the Java fore-arc basin at a direction N70.88E (Susilohadi et al. 2005) and is traced onshore along the Cimandiri river near Pelabuhan Ratu (Dardji et al. 1994).

Interpretations At the deformation front, trench sediments are offscraped from the oceanic basement and transferred to the upper plate (Figs 6 & 7). The change from tensional to compressional stress within the trench initiates thrust faulting and accretion of material to the lower slope. Discrete frontally accreted imbricate thrust slices and compressional folding are characteristic of this margin segment (Kopp et al. 2009) (Figs 6 & 7). The imbricate thrust belt and detachment folds form the frontal prism sandwiched between the trench and the Neogene accretionary prism (Kopp et al. 2001, 2002). The frontal prism forms the apex of the upper plate wedge and consists of frontally accreted, fluid-rich and thus mechanically weak material (von Huene et al. 2009). It is within the frontal prism that tectonic addition of trench sediment fill occurs by uplift displacement along the frontal thrust above the decollement (Figs 6 & 7). The imbricate thrust zone of the Neogene prism is primarily composed of trench sediment transferred from the frontal prism (Figs 6 & 7). The outer Neogene prism pronouncedly contrasts in style from the inner, less compressive Palaeogene prism (Kopp et al. 2001, 2002; Schlueter et al. 2002) (Fig. 7). This contrast in style has been related to the seismogenic behaviour of the subduction fault at depth (Wang & Hu 2006), predicting that the inner wedge never experiences compressive failure, thus providing a stable tectonic regime. The transition from the active Neogene prism to the tectonically more quiescent Palaeogene prism occurs along a distinct zone, where a splay fault system offsets the sea floor (Kopp et al. 2009) (Figs 6 & 7). The surface trace of this thrust fault system is

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Fig. 5.

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observed over a distance of c. 600 km along strike of the margin (Kopp & Kukowski 2003) and implies a continuous segmentation at an average distance of 30– 35 km from the deformation front. This transition is also related to a change in surface slope, which decreases from the outer to the inner prism (Figs 6 & 7). The Palaeogene prism, which forms the fore-arc high, serves as backstop to the material accreted in the Neogene prism. Both prisms show active deformation, however, to a much lower degree in the Palaeogene prism compared to its Neogene analogue. The Palaeogene prism forms the core of the large bivergent wedge, which shapes the fore-arc high along the central Sunda margin (Kopp & Kukowski 2003) (Fig. 2, panels a, b). The Late Palaeogene rising of the Himalayan orogenic zone, which is the source for the majority of sediment in the Sumatran and western Java sector of the Sunda trench (Susilohadi et al. 2005), contributed to the evolution of the accretionary fore-arc high, which is directly related to the sediment supply from the Himalayas.

Remaining issues The internal structure of the Palaeogene prism is not imaged well in reflection seismic data owing to limited energy penetration (Figs 6 & 7). This is a common phenomenon in many accretionary margins and may possibly be associated with internal deformation related to strong shear stress along the underlying seismogenic portion of the subduction fault. In the absence of deep drilling data, information on the composition is primarily gained from wide-angle seismic surveys. Seismic velocities increase from the frontal prism to the Neogene and Palaeogene prisms due to the greater rigidity of the consolidated and lithified material (Fig. 2, panels a, b). Mass balance calculations imply that the Palaeogene prism formed by accretion since the Eocene – Oligocene (Kopp & Kukowski 2003). The deformational segmentation as manifested in the kinematic discontinuities imaged by the refraction data suggests that accretion is non-linear. The rapid landward increase of the wedge thickness accompanied by backthrusting and uplift of the fore-arc high (Fig. 7) are indicative

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of basal accretion of the underthrust sequences (Gutscher et al. 1998), however, this process is not unambiguously imaged by seismic methods. Backthrusting along reverse faults is observed on the northern flank of the fore-arc high and initiates a successive northward thrust over the southern portion of the fore-arc basin (Fig. 7) (Schlueter et al. 2002; Susilohadi et al. 2005). The evolution of the Java basin is governed by the accretiondriven uplift of the fore-arc high, which forms a barrier to the trench and abyssal plain and by tectonically induced subsidence forming a rapidly filled depression (Susilohadi et al. 2005). The mechanisms for subsidence remain unclear. In the Sunda Strait, subsidence is likely to be linked to graben formation related to the transtensional regime here (Lelgemann et al. 2000) and the loading effect of Krakatau volcanoclastics (Susilohadi et al. 2005). In western Java, the Ciletuh–Jampang block (Fig. 1) is tilted to the SW. Its subsidence could reflect basal subduction erosion, however, our seismic data fail to image this process. Based on the interpretation of 20 reflection seismic profiles, Susilohadi et al. (2005) tentatively interpreted the oldest sequences in the Java basin to be of Middle Eocene to Late Oligocene age. A regional Upper Oligocene unconformity is traced as an erosional surface in the fore-arc basins from northern Sumatra to central Java (Fig. 7) and indicates that prior to the Neogene the shelf area was dominated by subaerial exposure or shallow water conditions (Susilohadi et al. 2005). This is also supported by well data from the shelf and onshore outcrops (Susilohadi et al. 2005). Later, sediment supply increased during the late Middle Miocene with the rising volcanic activity of the arc (Susilohadi et al. 2005).

Offshore central-eastern Java: seamount subduction and fore-arc erosion Observations The central-eastern Java segment is characterized by the subduction of an oceanic plateau, the Roo Rise, which is dotted with abundant basement

Fig. 5. Seismic wide-angle section for ocean bottom seismometer OBS 42 of line P18 (upper panel), where 23 instruments were deployed (black triangles in lower panel). The gap in instrument spacing in the trench is due to the great water depth exceeding the instrument’s pressure tolerance. The middle panel illustrates the calculated travel times on top of the seismic data shown in the upper panel. Rays in the lower panel are shot through a forward model with a velocity–depth distribution based on tomographic inversion. Shaded areas are not well resolved. A pronounced Pn mantle phase is recorded through the oceanic upper mantle reaching offsets of 160 km. The bending geometry of the oceanic crust underneath the trench inhibits ray coverage and results in a gap in first arrivals at 60–80 km offset. Consecutive inversion of different phases using a top-to-bottom approach recovered crustal and mantle velocities successively. See Figure 4 caption for phase information.

120 H. KOPP Fig. 6. Pre-stack depth migrated multichannel seismic data crossing the deformation front and trench offshore Sunda Strait. Location is shown in Figure 1. The lower panel shows the underthrusting of the Indo-Australian crust underneath the lower and middle slopes of the overriding plate. In the upper panel, the fold-and-thrust belt of the frontal prism is imaged with at least four consecutive pairs of forethrusts and backthrusts. Trench sediment is uplifted along the frontal thrust and subsequently rotated as it is incorporated into the frontal prism. The lowermost 400– 500 m thick unit of trench fill is underthrust beneath the frontal prism. A splay fault marks the transition from the Neogene accretionary prism to the Palaeogene prism and coincides with a change in slope angle. Seismic velocities are interval velocities obtained during pre-stack depth migration. The MCS data were acquired by the BGR, Hannover, during the GINCO cruise.

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relief (Masson et al. 1990). Relief elevation above the surrounding sea floor ranges from hundreds of metres to over 2 km. Oceanic crustal thickness is increased where the crust is altered by the emplacement of the Roo Rise. Profile P16 (Fig. 1) documents a crustal thickness of 9 km on the flank of the Roo Rise seaward of the trench (Fig. 2, panel c). Line P42, located off eastern Java and covering the Roo Rise south of the trench (Fig. 1), shows a pronounced Moho topography, with crustal thickness decreasing from 18 km in the oceanic domain to 11 km underneath the trench (Fig. 2, panel d). The high-resolution bathymetry maps the underthrusting of basement relief underneath the upper plate, for example the incipient subduction of a small ridge (Fig. 10b) currently positioned in the trench. Larger topographic features on the oceanic plate are resolved by the global bathymetry (e.g. a moderate-sized seamount of 70 km diametre at a distance of 40 km from the trench (Fig. 10a)). The surface effect of seamount subduction and the corresponding deformation of the lower slope are revealed well in the absence of a thick sediment apron, as is the case for central Java (Fig. 10). Frontal erosion has sculpted the lower slope off central and eastern Java and is associated with a northward retreat of the deformation front by up to 60 km (Kopp et al. 2006). This segment of the Java margin shows extremely high surface slope values at the lower slope of the overriding plate, reaching values .138 (Kopp et al. 2006). Sea floor topography based on global satellite altimetry data (Smith & Sandwell 1997) reveals that the fore-arc high is characterized by isolated topographic summits between 109 –1158E trending in a NW–SE direction parallel to the trench at roughly 108S (Kopp et al. 2006) (Fig. 1). They rise to water depth of 750 –1000 m and are roughly 1 km higher in elevation than the surrounding fore-arc high. This is documented by the two adjacent seismic profiles P16 and P18 (Fig. 2, panels b, c). The western line is located in the accretionary domain as described above and crosses the easternmost portion of the Java fore-arc basin. The eastern profile is positioned on the flank of the Roo Rise (Fig. 1). The difference in elevation of the fore-arc high on these two lines is c. 1000 m (Fig. 2, panels b, c). Further landward, a broad swell located between profiles P18 and P42 (Fig. 1) is anomalous along the Sunda Arc’s coastlines. The tomographic inversion of profile P16 (Fig. 2, panel c) images a strongly deformed fore-arc and basin, which extends for 50 km in a north–south direction (compared to 80 km on line P18 (Fig. 2, panel b)). Approximately 4 km of volcanic ashes and sediment are trapped in the basin, which is carried by the landward tilted basement above the Java unconformity (Fig. 11). At a depth of

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c. 14 km, the inversion images a subducted seamount around profile km 190 (Fig. 2, white arrow in panel c). The accretionary prism offshore Java is characterized by velocities generally not exceeding 5.0 km/s (Fig. 2). The subducted relief is inferred from the higher velocities (vp . 5.4 km/s) at the base of the accretionary prism retrieved along P16. OBH 30 (Fig. 12) covers the entire fore-arc and records the internal structure of the accretionary prism and subducting slab. Imaging, however, is intricate due to the severe deformation in this domain.

Interpretations The transition from frontal accretion along the western Java segment to frontal erosion off central Java occurs over a short distance of some tens of kilometres and is documented by the two adjacent seismic profiles P16 and P18 (Fig. 2, panels b, c). The central Java margin segment is currently experiencing frontal erosion associated with the underthrusting of the Roo Rise. The northward migration of the Java Trench and deformation front above the leading edge of the Roo Rise has exposed an area of c. 25 000 km2 of deeper sea floor formerly covered by the upper plate (Kopp et al. 2006). The corresponding northward shift of the axial position of the trench by about 60 km is moderate and may reflect a relatively recent onset of plateau subduction coupled with the arrival of the Roo Rise and Christmas Island Seamount Province at the trench. Based on the global satellitederived bathymetry (Fig. 1), Shulgin et al. 2010 infer that the edge of the plateau, which already subducted, could be located as far as 70 km north of the trench, which would correspond to an onset of plateau subduction at 1.1–1.3 Ma ago. However, there is no direct evidence. Topographic basement relief is abundant on the lower plate offshore Java (Masson et al. 1990) and modulates the structure and morphology of the overriding plate at various scales. The morphological perturbations of the lower slope resulting from subduction of oceanic relief depend on the size and structure of the subducted feature and on the nature of the overriding plate. Seamount subduction has been investigated at erosive margins (e.g. von Huene et al. 2000) where the seamounts leave pronounced re-entrant grooves as they plough through the small frontal prism before being subducted beneath the continental framework rock (von Huene 2008). Comparable embayments are not as distinct offshore Java (Fig. 10), where the accretionary material behaves more plastically. Topographic perturbations resulting from seamount subduction within the frontal prism are transient and the prism will heal after the relief is

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subducted to greater depth. Frontal erosion coincides with a steepening of the lower slope angle in the central Java sector compared to the neighbouring segments (Kopp et al. 2006), bringing the taper into the unstable domain here. This results from subduction of basement relief, which at first causes deformation and uplift of the thin leading edge of the fore-arc (Taylor et al. 2005). The unstable frontal prism is marked by small-scale re-entrant scars (Fig. 10), by mass failure and extensional normal faulting (Fig. 10a). Erosive processes are enhanced by the lack of sediment in the trench and the pronounced horst-and-graben structure in the trench where the plate bends underneath the fore-arc (Fig. 10b). Tectonic modification of the fore-arc offshore central to eastern Java is expressed in regional uplift affecting the entire marine fore-arc as well as in isolated zones of increased elevation (Fig. 1). The regional uplift pattern is caused by the subduction of the buoyant oceanic plateau, which results in uplift of the shelf as also described for the Hikurangi margin offshore New Zealand (Litchfield et al. 2007). Although due to the lack of independent data the onset of plateau subduction cannot be verified, it seems likely that it has been occurring since the late Pliocene when uplift and deformation of the upper plate intensified (Shulgin et al. 2010). Crustal thickening occurs mainly in the lower crust and seismic as well as gravity data confirm the presence of a crustal root here (Shulgin et al. 2010) as postulated by Newcomb & McCann (1987) to explain the absence of a correlated gravity anomaly. These results confirm numerical models, which predict crustal thickening to be concentrated in the gabbroic/basaltic layers (van Hunen et al. 2002). The observed thinning of the oceanic crust on profile P42 (Fig. 2, panel d) may either represent a local volume variation or may image the northern rim of the igneous expression. Subduction of smaller scale high relief gradient features likely accounts for the short wavelength anomalies observed along the fore-arc high (Masson et al. 1990; Abercrombie et al. 2001; Kopp et al. 2006). The observed uplift on profiles P18 and P42 (Fig. 2 panels c, d) is inferred to be caused by the impingement of oceanic basement relief and the associated

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deformation. A trench perpendicular compressive force is applied on the fore-arc by the relatively buoyant and thick subducting Roo Rise and its volcanic summits. This effect has also been reported for other margins, for example the Ryukuyu margin (Font & Lallemand 2009), Hikurangi margin (Litchfield et al. 2007), Costa Rica margin (Fisher et al. 2004) or the New Hebrides/Solomon arcs (Taylor et al. 2005). Uplift results from isostatic adjustment and is enhanced by crustal shortening of the overriding plate. The trench perpendicular compression leads to surface elevation of the forearc high, which greatly exceeds the original height of the seamount, as predicted by numerical modelling (Gerya et al. 2009). Surface uplift of 1 km is observed on P16 (Fig. 2, black arrow in panel c) and correlates with the position of the seamount at depth. Uplift is generated by crustal shortening and thickening of the overriding plate over a locked segment of the subduction thrust (Taylor et al. 2005). Backthrusting of the fore-arc high onto the fore-arc basin (Fig. 7) is observed along the entire segment and partially accommodates fore-arc convergence (e.g. Taylor et al. 1995). In addition to the deformation of the overriding plate, a subducted seamount at depth experiences faulting and possible rupture. Baba et al. (2001) investigate the stress field associated with seamount subduction and conclude that shear failure and fracturing or dismemberment of subducting seamounts occur. This will in turn affect seismic velocities and limit the velocity contrast between the accretionary prism and the subducted seamount. Regarding the subducted seamount detected on line P16, gaps in the ray coverage along the profile certainly inhibit the imaging (Fig. 2, panel c). The presence of a seamount is, however, supported by a number of very distinct surface effects that document the dynamic influence of seamount subduction on the fore-arc morphology. These effects are associated with the subduction of moderate-sized features (Dominguez et al. 2000; Gerya et al. 2009) and include local surface uplift, topographic perturbation of the lower slope, intensification of subduction erosion, and landward trench displacement. All of these key indicators are recognized off central Java (Fig. 2, panels c, d) (Kopp et al. 2006).

Fig. 7. Pre-stack depth migrated multichannel seismic data offshore western Java. Location is shown in Figure 1. The upper panel shows the southwestern extent of the profile from the trench to the Palaeogene prism, the lower panel displays the landward part of the line from the fore-arc high to the fore-arc basin. In the frontal prism, shortening is accommodated by imbricate thrusting of the frontally accreted sediment. Approximately 1/3 of the trench material is underthrust beneath the frontal prism in a 500– 900 m thick decollement zone, characterized by discontinuous high amplitudes. Two subducted oceanic basement highs of c. 1.4 and 0.8 km height, respectively, are imaged along the decollement zone. As off Sunda Strait (Fig. 6), a splay fault separates the Neogene and Palaeogene prisms and connects to the megathrust at depth. In the lower panel, sedimentary sequences above an Upper Oligocene unconformity are deformed by thrusting at the seaward part of the basin. The MCS data were acquired by the BGR, Hannover, during the GINCO cruise. After Kopp et al. (2009).

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Fig. 8. Seismic wide-angle section for ocean bottom hydrophone OBH 47 of line P18 (upper panel), located on the lower slope off western Java. The middle panel illustrates the calculated travel times on top of the seismic data shown in the upper panel. The landward increasing velocities of the fore-arc are documented by phases Pfp and Pg, which travel through the frontal prism and Neogene/Palaeogene prism, respectively. See Figures 4 and 5 captions for additional phase and display information.

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Fig. 9. Seismic wide-angle section record section of OBH 53 placed on the transition from the fore-arc high to the fore-arc basin on line P18 offshore western Java. Record phases through the Palaeogene prism (Pg prism) reveal slower velocities here compared to the fore-arc crust (Pg margin) and mantle (Pn). Strong reflections from the upper plate Moho (PcontP) specify the depth to the crust –mantle boundary below the fore-arc basin. See Figures 4 and 5 captions for additional phase and display information.

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Fig. 10. High-resolution bathymetric mapping offshore central-eastern Java. The swath lines represent the ship tracks, for example, along profiles P16, P18 and P42. Black background is not covered by data. Black pings are data artifacts. Light colours represent shallow water depth, dark blue colours represent deeper water. (a) shows the trench and lower slope offshore central Java. Location is indicated in Figure 3. The lower slope is heavily sculpted by subducting sea floor relief. The oversteepened slope locally fails and mass wasting onto the trench floor occurs. A large, 20 km wide re-entrant scar along the track of profile P16 indicates subducted sea floor relief, resulting in extensional faulting related to uplift. (b) images the trench floor disrupted by plate-bending induced normal faulting, which also affects basement relief. Type I landslides are observed along the lower slope. Two locations exhibit a concave surface slope, indicative of re-entrant scars, which have partially healed. Material is effectively transported from the fore-arc high into the fore-arc basin along extensive canyon systems.

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Fig. 11. Four-channel streamer section across the central Java fore-arc basin. Location is shown in Figure 1. The fore-arc basin strata onlap the fore-arc high and are tilted landward, indicating syndepositional and postdepositional vertical movement of the seaward portion of the basin and the fore-arc high.

Remaining issues Profile P42 crosses the hypocentre location of the 1994 Java Tsunami earthquake (Abercrombie et al. 2001; Polet & Thio 2003; Bilek & Engdahl 2007) (Fig. 1). The reverse mechanism event, which is associated with slip at a previously wellcoupled subducted seamount, showed normal faulting aftershocks that have been related to extension in the outer rise area (Abercrombie et al. 2001). This concept is supported by the high-resolution bathymetry of the trench area, which resolves platebending induced normal faulting (Fig. 10b) with vertical offsets of up to 500 m. The tomographic inversion of P42 as well as a corresponding multichannel line show indications of deeply penetrating faults (Lueschen et al. 2010; Shulgin et al. 2010) affecting the oceanic crust in the vicinity of the trench (Fig. 2, panel d). The deep structure of the fore-arc resolves the intricate geometry of the accretionary complex, which is characterized by heterogeneous uplift and deformation patterns. The velocity– depth distribution (Fig. 2, panel d) suggests the presence of accreted oceanic crustal fragments or detached oceanic basement relief. Remnants of accreted seamounts have been proposed to be present in the Japanese island arc (Isozaki et al. 1990), indicating shearing off and crustal underplating of oceanic basement material (Uchida et al. 2010). This scenario would also explain the fore-arc structure along P42. An anomalous high velocity structure is present at a depth of 13 km (Fig. 2, panel d). It is unlikely that a subducted seamount would still be intact under these conditions. Figure 10b maps the incipient subduction of a small seamount, which currently collides with the deformation front. This seamount as well as other bathymetric features in the trench and on the outer rise is broken by the bending-related normal faulting. The surface traces of the faults are continuous across the sea floor relief. Dismemberment of a seamount or oceanic crustal fragment will decrease seismic velocities and lower the

velocity contrast to the surrounding accretionary prism. As a consequence, seismic imaging will be distorted.

Offshore the Lesser Sunda islands: transition from oceanic subduction to continent– island arc collision Observations The margin segment south of the Lesser Sunda islands shows a different structure compared to its western counterpart. Here, a mature fore-arc basin, the Lombok basin, is observed at a water depth of 4400 m, which is limited to the west by the uplift associated with the Roo Rise subduction to the east and by collision of the Scott Plateau with the crystalline crust of the Sumba Island (Fig. 1) (Shulgin et al. 2009; Planert et al. 2010). The fore-arc high and accretionary prism are much more uniformly developed than in the neighbouring sector off Java, where isolated bathymetric elevations dominate the fore-arc high topography. Off Bali and Lombok, the fore-arc high is dominated by two elongated tectonic ridges (Fig. 13a) (Mueller et al. 2008; Krabbenhoeft et al. 2010) and diminishes in size and volume to the east. The age of the oceanic lithosphere decreases from Late Jurassic at 1208E to Early Cretaceous around 1108E (Heine et al. 2004; Mueller et al. 2008). On the incoming plate, Planert et al. (2010) determine a crustal thickness of 8.6 km off Lombok, increasing to 9.0 km to the east near the transition to the Scott Plateau. The transition from the oceanic crust of the Argo Abyssal Plain to the subducting crust of the Scott Plateau occurs over short distances of less than 50 km (Fig. 13c) and concurs with an increase in crustal thickness of c. 5 km, mainly accommodated by the upper crust (Fig. 2, panels f, g) (Planert et al. 2010). Convergence occurs at a rate of c. 70 mm/a in a direction N138E offshore Bali (Simons et al. 2007).

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Fig. 12. Seismic wide-angle section record section of OBH 30 located on the southern edge of the fore-arc basin offshore central Java on profile P16. This station covers the entire subduction complex and reveals the velocity structure of the accretionary prism (Pg prism) and the deep structure of the fore-arc (Pg margin and PcontP). PgP is the fore-arc crust basement reflection. The oceanic Pn phase to the south is reverse to the according phase on station OBS 42 displayed in Figure 5. See Figures 4 and 5 captions for additional phase and display information.

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Figure 13 displays oceanic basement structures in the Argo Abyssal Plain trending at angles of 45 – 608. Planert et al. (2010) argue that these structures trace the original sea floor spreading fabric as they trend parallel to the magnetic anomalies (Lueschen et al. 2010). Approaching the trench, plate bending induced normal faulting starts to dominate the sea floor fabric within 40 km of the trench axis (Fig. 13) (Planert et al. 2010). The resulting rough topography of the oceanic basement can be traced to several kilometres depth underneath the accretionary prism. Riffling of slope debris subparallel to the underthrust horst-and-graben relief is observed for parts of the frontal prism (Fig. 13a), similar to processes observed in northern Chile (von Huene & Ranero 2003). This region corresponds to reduced upper mantle velocities, which reach values of 7.5 km/s within a distance of 30 – 50 km from the trench (Fig. 2, panels e, f ). The fore-arc high rises steeply from the trench to water depth of less then 2500 m (Fig. 13a). Localized slope failure is observed on Figure 13 and is associated with the oversteepening of the lower slope. The sea floor morphology is dominated by two distinct, east –west trending ridge structures (Ridge A and Ridge B in Fig. 2, panel e) spaced c. 25 km apart (Fig. 13a). Uplift and tilting of piggyback basins hosted between the ridges (Fig. 13a) document active deformation and vertical displacement (Mueller et al. 2008). Seismic velocities of the fore-arc high do not exceed 5.5 km/s where it is in contact with the underthrusting oceanic crust at a depth of c. 13 km (Fig. 2, panels d, e), indicating a sedimentary composition as inferred for other parts of the margin. This is also supported by the relatively smooth magnetic response of the forearc high (Mueller & Neben 2006). The fore-arc high is fronting the Lombok basin, which carries 3–4 km of sediment above a seaward dipping basement (Fig. 2, panels d, e). The c. 9 km thick basement crust underneath the basin shows a high velocity gradient in its upper portion, decreasing in the lower crust. The upper plate Moho is located at a depth of c. 16 km underneath the fore-arc basin (Fig. 2) with upper mantle velocities of 7.5– 7.6 km/s.

Interpretations Sediment recycling is the principal process of mass flux along the lower slope south of the Lesser Sunda islands, where mass wasting of the fore-arc high supplies upper plate material to the trench (Fig. 13b), which is subsequently incorporated into the frontal prism. The oceanic crust is progressively faulted and altered as it approaches the trench. The complex shape of the thrust interface as imaged in the tomographic models (Fig. 2, panels d–g)

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suggests a high degree of fracturing of the oceanic crust with potential dissection into singular blocks. Horst-and-graben structures with vertical offsets of up to 500 m are recognized along the outer trench wall (Fig. 13a–c). Where the lower plate relief is not as pronounced, the lower slope is not impacted by subducted seamounts and is characterized by thrust faults (Fig. 13a). Offshore Bali and Lombok, the middle slope largely remains undisturbed, however, local undulations in the topography may result from lower plate fabric subducted beyond the frontal prism. A number of landslides have been identified along this margin sector (Brune et al. 2010a) and are classified into two types (Types I and II) as proposed by Yamada et al. (2010). Type I slides are of smaller dimension, developing on the lower slope and occurring frequently. The frequency of Type II slides is much lower compared to Type I failures. They occur on the middle and upper slope and are of larger dimension compared to Type I slides. Offshore Sumbawa at 1178520 E/ 11840 S, a Type II slide on the middle and lower slopes at 5300 m water depth with a width of 23 km affected a volume of c. 15 km3 (Brune et al. 2010a) (westernmost slope failure in Fig. 13b). From the absence of a deposition lobe it may be inferred that the landslide sediment deposit has been frontally accreted and incorporated into the frontal prism seaward of the headwall scarp. A northward offset of the headwall scarp indicates segmentation of the slide and collapse in successive events. Moving to the east, at least three adjacent slope failures (Type I) are identified in Figure 13b. They are of much smaller volume and only affect the lower slope. Lateral migration of slope failures has been predicted by analogue modelling (Yamada et al. 2010). The primary single slide will lead to changes in slope topography due to sediment displacement. Adjacent areas then become instable due to the resulting topographic undulations and another event is triggered in adjacent areas. The largest slide (Type II) with a volume of 20 km3 (Brune et al. 2010a) is encountered at 1198150 E/ 11830 S where it has left a significant deposit lobe resulting from failure of the middle and lower slope (Fig. 13b, c). This is located in the transition area from oceanic subduction to continent–island arc collision at the easternmost end of the Java trench. A sequence of seaward vergent normal faults on the outer trench wall relays the deepening of the sea floor from the Argo Abyssal Plain with a water depth of ,5000 m to the trench at 6500 m depth below the sea surface (Fig. 13b) (Planert et al. 2010). Near the deformation front, two landward vergent faults with a strike of c. 658 and thus subparallel to the magnetic anomalies (Mueller et al. 2008) are sculpting the deformation front and lower slope as they are subducted (Fig. 13c).

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The deformation front is additionally affected by transpressional deformation related to the southward curvature of the trench as it merges into the Timor trough (Fig. 13c). The relocated epicentre of the 1977 earthquake (Engdahl & Villasen˜or 2002) is shown in Figure 13b. The main event and the aftershock sequence likely resulted from slip along the re-activated inherited sea floor fault fabric as it bends underneath the upper plate and ruptures the oceanic lithosphere to a depth of 30 – 50 km (Spence 1986; Lynnes & Lay 1988). One of the most prominent features on the wideangle profiles is the shallow upper plate mantle, which is found at a depth of c. 16 km underneath the fore-arc basin. The low seismic velocities of c. 7.5 km/s detected here (to c. 2 km below the Moho) are attributed to hydration and serpentinization of mantle peridotite (Faccenda et al. 2009), requiring faults to penetrate the oceanic crust and reach deep into the mantle. This is supported by the hypocentre relocation of the 1977 Sumba earthquake sequence, which resolved normal faulting to affect the oceanic lithosphere to a depth of 34 km (Spence 1986) as indicated above. The shallow position of the upper plate mantle may fundamentally affect seismogenesis along the Java margin as it limits the extent of the seismogenic zone. The interface contact with weak, hydrated minerals such as serpentinite, which mechanically cannot support stick-slip behaviour, would result in stable sliding downdip of the seismogenic zone (e.g. Hyndman et al. 1997; Oleskevich et al. 1999). However, exceptions to this concept may exist, for example offshore Sumatra, where earthquake nucleation has been proposed to occur in the mantle (Dessa et al. 2009; Klingelhoefer et al. 2010) or offshore Japan, where earthquake clusters below the depth of the fore-arc Moho are related to seamount detachment (Uchida et al. 2010). Along with the depth extent of the seismogenic zone, the size of the lateral rupture zone determines the potential magnitude of subduction thrust earthquake. Local asperities that may act as barriers to rupture thus will also influence the magnitude. The eastward propagating minimum 200 km rupture of the 1977 Sumba normal faulting event was likely limited by

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the thick crust of the Scott plateau, acting as a barrier to further rupture propagation.

Remaining issues The transition from oceanic subduction to the collisional regime along the Timor trough occurs south of Sumba Island (Fig. 13c), where the crystalline crust of the Sumba block is imaged in the tomographic model (Fig. 2, panel f ). The easternmost profile is located east of Sumba Island around 1218E (Fig. 1) (Shulgin et al. 2009) and documents the early stages of continent–island arc collision along the westernmost extension of the Timor trough (Fig. 2g). The relative motion along this segment of the Timor trough has slowed to c. 15 mm/a. The incoming crust of the Scott plateau reaches a thickness of 15 km as it subducts below the fold-and-thrust belt of the upper plate. The increased sediment thickness is reflected in decreasing water depth in the trough (Fig. 13c). Sediments are likely sourced from the Australian continent and contribute to the evolution of a large accretionary body with a width of c. 130 km. This evolving collisional system is dominated by the Sumba Ridge, a high velocity block that acts as backstop to the accretionary prism in the south (Fig. 2, panel g). Backthrusting onto the Savu Basin in the north originates from the compressional deformation caused by the northward propagation of the Australian lithosphere (Bock et al. 2003).

Conclusions This study investigates contrasting modes of deformation in three segments of the Java convergent margin, defined by varying processes of mass transfer. Sediment supply to the trench acts as the principal factor governing lower to upper plate material transfer. The decreasing sediment supply to the Java trench from west to east correlates with a changing pattern of mass flux: from sediment accretion offshore western Java to tectonic erosion off central Java. Sediment accretion characterizes the Sumatra sector of the Sunda margin, where sediment input on

Fig. 13. High-resolution bathymetric mapping offshore Lombok to Sumba. Location is indicated in Figure 3. (a) The accretionary prism is dominated by two elongated ridges which host piggy-back basins. The topography of the lower plate is dominated by plate-bending induced normal faulting locally overprinted by original sea floor spreading fabric. Oceanic crust topography causes riffling of lower slope material upon subduction underneath the frontal prism. Mass wasting occurs to the fore-arc basin in the north. (b) Slope failure results in landslides affecting the lower slope (Type I failure) or the middle and lower slope (Type II failure). The headwall scarp of the Type II failure is offset northwards, indicating failure in successive phases. Three adjacent Type I slides are indicated and associated upper material wasting is observed on the trench floor. White star indicates hypocentre location of 1977 Sumba earthquake. (c) Transition to the Timor trough coincides with a shallowing of the sea floor and is associated with transpressional deformation and uplift of the overriding plate. A debris lobe is observed in the trench and will eventually be re-incorporated into the frontal prism. Refer to Figure 10 caption for additional display information.

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the oceanic plate continuously increases to the north with closer proximity to the Ganges –Brahmaputra system (Moore et al. 1980). The western Java segment is characterized by a net addition of material from the lower plate to the upper plate and by an oceanic sea floor topography smoothed by a sediment apron. Offshore western Java, frontal sediment accretion dominates and c. 2/3 of the trench sediment sequence is incorporated into an imbricate thrust belt (Schlueter et al. 2002) (Figs 6 & 7). The thickness of material subducting beyond the frontal accretionary prism ranges from 500 to 1000 m per trench km here. Basal accretion likely occurs below the fore-arc high, contributing to the evolution and uplift of a .100 km wide bivergent wedge (Fuller et al. 2006; Kopp et al. 2009). To the east, offshore central Java, the transition from sediment accretion to tectonic erosion occurs over a distance of less than 100 km around 1108E (between profile P18 and P16 in Fig. 1). Here, the trench is devoid of sediments except for isolated sediment ponds (Masson et al. 1990). A complex canyon system traverses the continental slope and supplies material to the Java and Lombok fore-arc basins. Sediment discharged from Java and the Lesser Sunda islands does not reach the trench, but is trapped in the fore-arc basins as the distance from the trench to the active volcanic arc increases from west to east. The nature of the basement of the arc framework crust underneath the Java and Lombok fore-arc basins remains enigmatic. Based on previous work by Curray et al. (1977) and Kopp et al. (2002), it is proposed that the fore-arc crust could be composed of an altered oceanic terrane, which resisted subduction due to increased positive buoyancy (Planert et al. 2010). An alternative view is based on the rock record of Sumba Island: the seaward extent of continental crust south of Java and the Lesser Sunda islands could be the lateral continuation of the Late Cretaceous arc massif (Rutherford et al. 2001). High relief oceanic basement is subducting offshore central and eastern Java, leading to a rough sea floor and causing frontal erosion of the fore-arc (Fig. 10). The margin geometry is influenced by the subduction of an oceanic plateau, the Roo Rise, underneath the Java fore-arc. The Nusa Tenggara segment offshore the Lesser Sunda islands experiences the transition from oceanic subduction to continent island arc collision (Fig. 13), with a rapid change in upper plate structure along strike. The eastern Sunda margin is prone to large potentially tsunamigenic landslides. Landslides are categorized in Type I and Type II slides, following the nomenclature of Yamada et al. (2010). Type I slides are of smaller dimension and occur on the lower slope, while Type II slides affect a larger

area/volume and are observed on the middle and upper slopes. Both types are triggered by the oversteepening of the slope either due to the subduction of relief or near a thrust surface in the frontal imbricate thrust fan (Fig. 13). While the smaller Type I slides are ubiquitous along the Java margin, the larger Type II slides are only observed in the easternmost segment. Brune et al. (2010b) have identified 12 landslides along the Sunda margin from high-resolution multibeam bathymetry. While the volume of the Type I landslides is typically less than 5 km3, the three largest ones, which are located in the transition zone from the Java trench to the Timor trough, show large volumes of 15 – 20 km3 and have originated on the middle slope, which qualifies them as Type II slides (Fig. 13b). A potential contribution of the Type II slides identified in the vicinity of the 1977 Sumba earthquake is not verified. Tsunami propagation modelling successfully predicted the observed run-up heights from the earthquake tsunami alone and does not necessarily require a further contribution from a landslide tsunami (Brune et al. 2010a). An additional process for tsunami generation is the potential activation of splay faults during the co-seismic phase. Splay faults connect to the megathrust at depth and dip steeply to the surface, as imaged offshore western Java (Kopp et al. 2009) (Figs 6 & 7) and off the Lesser Sunda islands (Lueschen et al. 2010). Thus the low-angle slip of the megathrust will potentially be transferred to a higher angle, which may greatly enhance sea floor displacement (Tanioka & Satake 1996). Due to the lack of deeply penetrating multichannel seismic data, the role of potential splay faults in the generation of the 1994 and 2006 Java tsunamis remains unresolved. Deep-seated subduction processes excerpt control on the structure and deformation of the upper plate as well as on the seismogenesis of the fore-arc. The most dramatic effects are observed in the central-eastern Java segment, where deformation of the sedimentary units in the fore-arc basin (Figs 7 & 11) and backthrusting of the fore-arc high onto the basin (Fig. 2, panels c, d) are documented. A decrease in the subduction angle of the underthrusting plate, as detected off South America, however, is not observed here (Koulakov et al. 2007). This concurs with results from numerical modelling, which predict that a moderate-sized plateau will not significantly alter the subduction angle (Gerya et al. 2009). In addition, the modelling also predicts that a decrease in magmatic activity is unlikely. Tomographic inversion has revealed the interplay between the fore-arc and the volcanism on Java, where the high vp/vs ratio of a pronounced low velocity anomaly in the Javanese crust is indicative of fluid

THE JAVA CONVERGENT MARGIN

ascent from the underthrusting plate to the volcanic arc (Koulakov et al. 2007) and has been interpreted to image the related fluid ascending paths (Wagner et al. 2007). Other predictions based on numerical modelling regarding the fore-arc morphology are also matched: a local increase in topography is observed in the overriding plate as well as a northward displacement of the deformation front (Kopp et al. 2006), indicative of erosive processes here. Stress fluctuations govern erosion, which requires a strong subduction thrust fault and a mechanically weak overlying wedge. On short timescales, earthquakes are a common mechanism to cause variations of stress, which then occur from the interseismic phase to the co-seismic activity (Wang et al. 2010). Changes in basal fault strength may also be caused by the rough topography of an oceanic plate lacking a significant sediment cover. Deformation of the wedge caused by the impinging bathymetric features will mechanically weaken the prism, which is then overlying a strong basal detachment, providing conditions favouring subduction erosion. Basal subduction erosion would pose a tectonic mechanism for basin subsidence; however, this would require the underthrusting plate to remain in contact with the upper plate from the trench to underneath the fore-arc basin. The tomographic images of Figure 2 clearly demonstrate that the underthrusting plate dips into the upper mantle beneath the outer fore-arc high. Further evidence for this configuration comes from earthquake hypocentres distribution (Wagner et al. 2007; Wittwer 2010) and gravity modelling (Grevemeyer & Tiwari 2006; Planert et al. 2010; Shulgin et al. 2010). In addition, satellite magnetic data record a significant anomaly extending seaward, resulting from a hydrated mantle wedge underneath the fore-arc (Blakely et al. 2005). This then raises the question if subduction erosion of the upper plate’s lithospheric mantle wedge occurs and if this accomplishes basin subsidence. This issue, however, is beyond the scope of this paper.

Seismogenesis The Java margin is characterized by a notable absence of Mw . 8 earthquakes compared to its Sumatran counterpart, leading to the question of what controls seismic rupture and consequently the potential size of earthquakes offshore Java. The magnitude of an earthquake is associated with the size of its rupture zone. Slip motion on a fault will depend on the tectonic environment of the source region (Bilek 2007). Two aspects are related to slip motion: (1) the role of the decollement zone and (2) the role of sea floor relief acting as asperities or barriers to rupture.

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The notion that trench sediments affect seismogenesis was brought forward by Larry Ruff (Ruff 1989) and is here extended to the concept of the subduction channel in general. On a global scale, giant megathrust earthquakes (Mw  8.5) are observed in systems characterized by sediment-flooded trenches [e.g. Sumatra (1883, 2004, 2005), Southern Chile (1960, 2010), Alaska/Aleutians (1964, 1965, 1986)] as well as at erosional margins [Kamchatka (1952), Kuril Islands (1963), Northern Chile (1868, 1877)], which show a subduction channel of several hundred metres in thickness. The existence and geometry (thickness) of a subduction channel thus influences rupture propagation to a greater degree than the nature of the material in it (trench sediment v. upper plate erosional debris) (Tanioka et al. 1997; Bilek & Lay 1999). A discussion on the role of fluids in this context, however, is beyond the scope of our data. The second aspect regards sea floor roughness and the question of whether basement relief acts as an asperity or barrier to seismic rupture (Bilek & Lay 2002). Certainly this will not play a role where basement highs are deeply buried in the subduction channel. Here, subduction channel material smoothes sea floor relief and cushions upper plate contact. Where this is not the case, underthrusting seamounts or ridges may pose a limit to lateral rupture propagation, as do crustal faults (Collot et al. 2004). Bathymetric relief on the underthrusting plate will lead to variations in mechanical coupling and high friction models as well as low friction models have been proposed (Cloos 1992; Mochizuki et al. 2008). Large seamounts (3 km height) may increase the normal stress between the plate and raise interplate coupling (Scholz & Small 1997). On the contrary, reduction of normal stress has been proposed to result from elevated pore pressure of entrained fluid-rich sediment during erosion (von Huene et al. 2004). Weak interplate coupling may be related to the damage caused by erosion that inhibits the accumulation of elastic strain energy (Mochizuki et al. 2008). Recurring Mw c. 7 earthquakes are related to seamount subduction in the Japan trench (Mochizuki et al. 2008) where weak coupling has been linked to fluid-rich sediment and migration of fluids at the base of the seamount. Along the Java subduction zone, different tectonic features exert a first-order control on the seismogenesis of the margin and govern the lack of Mw . 8 megathrust earthquakes: (1)

Our tomographic images reveal a shallow upper plate Moho with low mantle velocities, indicative of hydrated minerals (Fig. 2). Hydration is caused by fluids, which are released from the subducting slab and

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(2)

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entrenched sediments, leading to serpentinization of the mantle wedge (Hyndman & Peacock 2003). The limited downdip extent of the seismogenic zone is also supported by gravity data and thermal modelling (Grevemeyer & Tiwari 2006) as well as by the fore-arc morphology (Krabbenhoeft et al. 2010). In the central Java segment, seismic rupture would additionally be limited along strike by subducted basement relief acting as barriers that will resist co-seismic slip. Erosional damage related to seamount/Roo Rise plateau subduction may hinder the accumulation of elastic strain (Mochizuki et al. 2008). Local elevation in pore pressure of sediment entrained during underthrusting of a seamount may also be expected. The uneven slip distribution recorded during the co-seismic phases of the recurring tsunami earthquakes on this margin sector (Fig. 1) (Ammon et al. 2006; Bilek & Engdahl 2007) document the highly heterogeneous plate coupling of the fore-arc. The structural diversity of the underthrusting plate in conjunction with fluid-related processes governs the heterogeneous plate coupling offshore Java.

A heterogeneous structure has also been documented for the onshore portion of the upper plate fore-arc. Two high velocity, rigid blocks sandwich a low velocity anomaly in southern Java, which is interpreted as a weakened contact zone. The predominantly strike-slip focal mechanism of the Mw ¼ 6.3 Java event in May of 2006 corresponds to the orientation of this contact zone (Wagner et al. 2007). However, our data lack the resolution to precisely determine the role of upper plate heterogeneity in seismic rupture propagation patterns. In summary, the complex megathrust interface geometry is the main factor for the observed absence of large (.8) magnitude earthquakes offshore Java (Newcomb & McCann 1987), while smaller magnitude earthquakes frequently occur. The interplay between the tectonic habitat of the source region and the seismogenesis of large megathrust earthquakes is only crudely understood to date. The topic invites further research in the future to better understand the seismogenic segmentation and the specific geohazard potential of convergent plate boundaries. I am indebted to numerous colleagues who have been involved in the Java margin studies. The concepts presented here have evolved from the many discussions with the members of my group at IFM-GEOMAR (D. Hindle, L. Planert, M. Scherwath, A. Shulgin, A. Wittwer, J. Zhu and M. Zillmer). The data acquisition would not have been possible without the enthusiastic

support of J. Bialas, A. Krabbenhoeft, C. Papenberg, J. Petersen, I. Trummer, D. Wagner. Special thanks go out to E. R. Flueh for his readiness to lead the cruises. W. Weinrebe is warmly thanked for processing of the multibeam data. The studies were conducted in collaboration with Y. Djajadihardja (BPPT Jakarta), B. Luehr and O. Oncken (GFZ Potsdam), C. Mueller, C. Gaedicke, C. Reichert (BGR Hannover), and W. Rabbel (CAU Kiel). This work would not have been possible without the continuous support of the SONNE program by the German Federal Ministry for Science and Technology BMBF. We are indebted to the captains and crews of RV SONNE for the excellent support and performance at sea. Cruises SO137, SO138, SO176, SO179 and SO190 were supported by grants 03G0137A (GINCO I), 03G0138A (GINCO II), 03GO176A (MERAMEX), 03G03G0190A and 03G0190B (SINDBAD project). Additional support was supplied by the DFG through the SUNDA project (grant KO2961/1– 2) and by the GEOTECHNOLOGIEN program of BMBF and DFG for the SUNDAARC project (grant 03G0579B). I thank reviewer D. Scholl for many discussions and thoughtful comments that greatly helped to improve the manuscript. Comments by an anonymous reviewer additionally tightened the manuscript. I kindly acknowledge volume editor R. Hall and M. Cottam for their editorial guidance.

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Subducting slab structure below the eastern Sunda arc inferred from non-linear seismic tomographic imaging S. WIDIYANTORO1*, J. D. PESICEK2 & C. H. THURBER2 1

Faculty of Mining and Petroleum Engineering, Bandung Institute of Technology, Jl. Ganesha 10, Bandung 40132, Indonesia

2

Department of Geoscience, University of Wisconsin-Madison, 1215 W Dayton St., Madison WI 53706, USA *Corresponding author (e-mail: [email protected]) Abstract: Detailed P-wave speed velocity structure beneath the Sunda arc has been successfully imaged by applying a non-linear approach to seismic tomography. Nearly one million compressional phases from events within the Indonesian region have been used. These include the surface-reflected depth phases pP and pwP in order to improve the sampling of the uppermantle structure, particularly below the back-arc regions. We have combined a high-resolution regional inversion with a low-resolution global inversion to minimize the mapping of distant aspherical mantle structure into the study region. In this paper, we focus our discussion on the upper mantle structure beneath the eastern part of the Sunda arc. The tomographic images confirm previous observations of a hole in the subducted slab in the upper mantle beneath eastern Java. The images also suggest that a tear in the slab exists below the easternmost part of the Sunda arc, where the down-going slab is deflected in the mantle transition zone. In good agreement with previous studies, the properties of the deflected slab show a strong bulk-sound signature.

The Sunda arc, located in the western part of the Indonesian region, extends from northwestern Sumatra to Flores, that is, to the west of the Banda arc-Australian continent collision zone. The Sunda arc marks the subduction of the Indo-Australian plate beneath the Eurasian plate. The age of the incoming plate varies laterally. It is relatively young along Sumatra, where subduction is highly oblique. In contrast, below the eastern part of the Sunda arc (i.e. Java and small islands east of it), the age of the subducted plate is significantly older and the convergence direction is almost perpendicular to the arc. Mueller et al. (1997) suggest that the age of the incoming oceanic plate ranges from 40 Ma beneath northern Sumatra to 110 Ma south of Java. The lateral variation of the nature and age of the incoming plate influences the style of deformation and seismicity along the Sunda arc (Cloetingh & Wortel 1986). The character of subduction-related seismicity changes abruptly from Sumatra to Java. Seismicity does not exceed a depth of 300 km beneath Sumatra, except for some small events in the southeastern part of the island, but earthquakes occur at depths of up to c. 670 km below Java to the east (Fig. 1). The convergence rate of the Indo-Australian and Eurasian plates in general increases from Sumatra to the easternmost part of the Sunda arc (Minster & Jordan 1978). Tregoning et al. (1994) measured

convergence rates of 6.7 + 0.7 cm a21 across the Java trench between Christmas Island, SW of Java, and west Java in a direction of N118E + 48. This is similar to the relative plate velocity between Australia and Eurasia predicted by the NUVEL-1A plate motion model (DeMets et al. 1994). The movement of Australia northward caused rotation of blocks and accretion of microcontinental fragments to SE Asia (Hall 2002). With this complex tectonic setting, it can be expected that the Sunda arc overlies a heterogeneous mantle partly evident from its seismicity (Fig. 1). Previous seismic tomographic studies of mantle structure below the study region focused on the deep subduction of the Indo-Australian plate. The imaged slab penetrates directly into the lower mantle, where it deflects in the uppermost lower mantle and sinks almost vertically to a depth of at least 1200 km (Fukao et al. 1992; Puspito et al. 1993; Widiyantoro & van der Hilst 1996, 1997; Bijwaard et al. 1998). The aim of this study is to explore the detailed structure of slabs in particular in the upper mantle beneath the eastern Sunda arc, where a pronounced seismic gap exists in a depth interval of c. 250–450 km (e.g. Newcomb & McCann 1987) and a hole in the subducted slab has been reported (Hall et al. 2009). We present new P-wave seismic images produced by means of an improved tomographic imaging technique (Pesicek

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 139–155. DOI: 10.1144/SP355.7 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Map of the study region. Circles depict the epicentres of relocated events (Engdahl et al. 1998, 2007) in the region occurring between 1964 and 2007, with colours denoting the hypocentre depths. Black lines indicate the position of the vertical cross sections displayed in Figures 4 and 8.

et al. 2010). We note that P-wave tomographic images of the subduction zone west of the present study area (i.e. below the western Sunda arc) have been presented in detail by Pesicek et al. (2008, 2010). Herein we discuss only the eastern Sunda arc and present new S-wave images to enhance our discussion and interpretations.

Data and method Data Engdahl et al. (1998) carefully relocated nearly 100 000 earthquakes that occurred between 1964 and 1995 by using a nonlinear scheme and the radially stratified ak135 velocity model developed by Kennett et al. (1995). These data consist of c. 13 million P, pP, pwP, PKP and S phase arrival times reported by almost 6000 globally distributed seismographic stations. In this study, we have used an updated data set covering the period 1964 to 2007 (Fig. 1). In the western Sunda arc region, the new data set has been further groomed to provide more accurate depths (Engdahl et al. 2007). The updated dataset consists of 957 262 compressional phases from events within the Indonesia region, including 10 640 pP and 4239 pwP phases. Detailed data selection criteria are presented in Pesicek et al. (2010).

Method We have used a cellular representation of mantle structure by discretizing the entire mantle using cells of 58  58 (with 16 layers down to the bottom of the mantle), but in the study region we have employed a finer grid of 0.58  0.58 (with 19 layers down to 1600 km) in order to allow the resolution of relatively small-scale structures. Such a model parameterization minimizes contamination by structures outside the volume being investigated (Fukao et al. 1992; Widiyantoro & van der Hilst 1996, 1997). We solved for perturbations to 178 272 model slowness cells using the iterative LSQR algorithm (Paige & Saunders 1982), a conjugate gradient technique first used in seismic tomography by Nolet (1987); see also Spakman & Nolet (1988). Following Bijwaard & Spakman (2000) and Widiyantoro et al. (2000), we used a step-wise procedure to solve the non-linear travel-time inversion for seismic velocity variations. In the main step, ray paths and travel times are updated by 3D ray tracing through intermediate realizations of the model. The 3D ray tracing is based on the pseudobending method of Koketsu & Sekine (1998), originally developed by Um & Thurber (1987). In this study, this procedure was improved upon by use of an a priori global crustal model (CRUST 2.0; Bassin et al. 2000) in order to reduce the initial residual variance of the data (Pesicek et al. 2010).

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In the Indonesian region, we initially traced rays from stations to sources through the global spherically symmetric model ak135 (Kennett et al. 1995). We then replaced the shallowest layer with velocity values from the a priori crustal model. We remark that the images produced by this nonlinear inversion are comparable to those of a one-step linearization (Pesicek et al. 2008), but the progressive updating of the slowness field and ray paths results in larger magnitudes of the wave speed perturbations of c. 30%.

Presentation of seismic tomograms We interpret P-wave travel-time residuals in terms of velocity perturbations relative to the ak135 reference velocity model (Kennett et al. 1995). In this section we present the images achieved after conducting five iterations with sources fixed to update the 3D velocity model (Pesicek et al. 2010). In Figure 2, we present P-wave velocity anomaly maps for depths representing the upper mantle and transition zone. From the tomographic inversions, we infer that the subducted slab is defined by a laterally continuous region of higher-thanaverage P-wave velocity in the upper mantle and transition zone. Most parts of the region of interest, in particular along the island arc, are sufficiently sampled by seismic rays and reasonably resolved (Fig. 3). In the upper mantle and transition zone, the dimension of the smallest feature that is resolved is about 100 –200 km. Notice that resolution generally degrades with increasing distance away from the slab due to irregular and/or poor sampling. With this and the limitations of the resolution tests in mind we only interpret the large-scale structures. The image of the slab in the uppermost mantle resembles the present-day Java trench and parallels the present-day Sunda arc. A high velocity slab is detected beneath Java and islands to the east, but hardly seen beneath eastern Java at depths around 250 –450 km (Fig. 2). Further examination of anomaly maps for different depth intervals indicates that the fast slab is also absent at shallower depths below the easternmost part of the Sunda arc. The complexity of the inferred slab structure is further illustrated by vertical sections across the eastern Sunda arc (Fig. 4). The images in Figure 4 suggest that the Indo-Australian plate dips steeply beneath the Java arc and is only partially outlined by a seismic zone. The magenta dashed lines on the cross sections depict our preferred interpretation of the subduction angles. In general, the velocity images and the seismicity reveal a gently dipping slab (10 –308) from the trench to the arc. Then the slab dips more steeply (c. 60 –708) down to the transition zone. However, it appears that the dip changes

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to nearly vertical at depths of about 400 km below the eastern Sunda arc, where the slab is deflected in the transition zone (cross sections O –R in Fig. 4). The seismogenic slab seems to be continuous, except beneath eastern Java (see cross sections E –G in Fig. 4) where a pronounced seismic gap is observed between 250 and 450 km depth. Further east (cross sections K –M in Fig. 4), we detected higher-than-average seismic velocities, but with smaller amplitudes than elsewhere. The amplitude reduction in the seismic gap may suggest a ‘necking’ of the slab (cf. Widiyantoro & van der Hilst 1996, 1997). The summary of our observations and interpretations is represented by the 3D smoothed isosurface plot for þ0.85% velocity perturbation illustrating the subducted slab in the upper mantle beneath the eastern Sunda arc and the western part of the Banda arc (Figs 5 & 6). Here, the hole in the slab below eastern Java and a tear in the slab beneath the easternmost part of the Sunda arc and perhaps also beneath the Banda arc are clearly depicted, while the small hole below Flores is related to the ‘necking’ of the slab. Additionally, we present S-wave tomograms derived by using an updated version of the reprocessed S arrival time data of Engdahl et al. (1998). Information from S-wave data helps characterize the inferred velocity variations but previous studies have not produced good constraints on slab structure (e.g. Zhou & Clayton 1990). This is likely due to the relatively high noise level of the International Seismological Centre (ISC) S-wave data. However, refinement of the travel time data, particularly the reprocessing of multiple data sets to extract improved S-wave information conducted by Engdahl et al. (1998), has greatly benefited the current study. Engdahl et al. (1998) used S phases in the initial source location and an appropriate S-wave reference velocity model (ak135; Kennett et al. 1995). Most aspects of the S-wave tomographic imaging technique employed here are similar to those of the P-wave data, except that for a ray to be included in the inversion, the travel-time residual for S relative to the ak135 reference model has to lie in the range +15.0 s, in contrast to the dynamic reweighting of P-wave residuals (Pesicek et al. 2010). Figures 7 and 8 contain the resulting S-wave images displayed in the same style as the P-wave images given in Figures 2 and 7 in order to provide direct comparison. Notice that in general the S-wave images depict the slab in the upper mantle well, although some smearing occurs in particular in the uppermost mantle. In the following section we discuss our observations based on the P- and S-wave images and relate them to tectonic processes in the region.

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Fig. 2. Layer anomaly maps depicting results of the inversion using full P-wave arrival time data for upper-mantle and transition zone structures below the eastern Sunda arc. Velocity perturbations relative to ak135 developed by Kennett et al. (1995) are shown from 22% to þ2%. For each map, mid-layer depths are listed in km at the top. The top layer anomaly map has had a crustal correction applied using the global crustal model CRUST 2.0 (Bassin et al. 2000) in the study region, as discussed by Pesicek et al. (2010). The image of the slab in the upper mantle (from 52.5–450 km) parallels the present-day Sunda trench. Notice the broadening in map view of the slab beneath the easternmost part of the Sunda arc at depths of around 530–615 km.

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Fig. 3. Resolution tests calculated using the same parameters as the real data inversion. Noise was added based on the actual, but randomized, residual distribution following Pesicek et al. (2010). (a) Spike model resolution tests after two iterations. Synthetic 4% velocity perturbation input spike anomalies (2.5  2.58; black contours) are separated by 2.58 in latitude and longitude and by 2 layers in depth. Depths with no input anomalies are shown and the perturbations in these layers are an indication of vertical smearing.

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Fig. 3. (Continued) (b) Alternate spike model, the same as (a) except the input pattern is shifted to be the opposite of (a), that is, layers with (without) anomalies in (a) now lack (have) them.

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Fig. 3. (Continued) (c) Synthetic slab resolution test. In general, the synthetic slab (4% velocity perturbation) is well recovered north of the Java trench. The geometry of the synthetic slab is shown by magenta contours in each layer where it is present.

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Fig. 4. Vertical sections across the convergent margin in the eastern part of the Sunda arc through the P-wave model plotted as velocity perturbations relative to ak135. Contour scales are from –2% to þ2%. Circles depict earthquake hypocentres projected from a distance of up to 55 km on both sides of the plane of section. Magenta dashed lines depict our preferred interpretation of the subduction angle.

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Fig. 4. (Continued) For each region, the cross section length is listed at the top-right in km, and in degrees at the bottom-right. Note that there are some imaging artifacts, for example, the dipping structures south of the slab depicted in cross sections B–D, that look like slabs. On the other hand, the flat lying slab on the 660 km discontinuity in cross sections P– R seems to be a real feature (see also Widiyantoro & van der Hilst 1997).

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Fig. 5. Three-dimensional smoothed iso-surface plots for þ0.85% P velocity perturbation relative to ak135 illustrating the subducted slab in the upper mantle beneath the eastern Sunda arc and the western part of the Banda arc viewed from the NW.

Fig. 6. Same as Figure 5, but viewed from the north. Notice the existence of a hole in the subducted slab below eastern Java, a tear in the slab beneath the easternmost part of the Sunda arc, and the small hole below Flores interpreted as a ‘necking’ of the slab.

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Fig. 7. Same as Figure 2, but derived using S-wave data and displayed using a 4% perturbation scale. Notice that the slab in the upper mantle parallel to the present-day trench is well imaged.

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Fig. 8. Same as Figure 4, but derived using S-wave data and displayed using a 4% perturbation scale. Notice that the subducted slab is reasonably well imaged. However, the shallow structural feature beneath the back-arc region is less resolved due to lack of ray sampling.

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Fig. 8. Continued.

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Discussion Based on the results of our resolution tests (Fig. 3), the tear in the slab beneath eastern Java as depicted by the tomographic images (Figs 2–6) is judged to be a real feature. The tear is also observed in the high resolution P-wave tomographic model of Bijwaard & Spakman (2000). A number of processes could induce a tear in the down-going slab involving both crustal buoyancy variations and mantle processes acting on the slab. For example, Hall et al. (2009) suggested that the tear beneath eastern Java forms a hole in the subducted slab. This is related to the trench stepping back to the south after the slab broke due to the arrival of a buoyant plateau at c. 8 Ma that was unable to subduct. The convergence initially caused contractional deformation before the slab broke in front of the plateau. After the trench stepped back, subduction resumed behind the plateau causing the hole to develop (Hall et al. 2009). Similarly, a hole in the subducted slab has also been observed below the Izu – Bonin–Mariana trench by Miller et al. (2006). Miller et al. (2004, 2005) proposed that the Ogasawara Plateau collision and subsequent subduction are related to distortion and heterogeneity in the Pacific plate at depth. Using their new reconstruction of the subducted Pacific plate with a 3D visual model, Miller et al. (2006) inferred that the initiation of the oceanic plateau subduction at c. 8 Ma and the resulting complex slab morphology are related. For the Java case we agree with and support the interpretation by Hall et al. (2009). In addition, we note that any other process that either slows the rate of input into the trench relative to the rate of subduction of the slab within the Benioff zone or causes a trench jump would open a gap in the slab. The tear in the slab beneath the easternmost part of the Sunda arc may be explained in a similar way. When Timor, part of the Australian continental plate, together with small islands to the east of it arrived at the former trench and collided with the Banda arc at c. 3 Ma (Hall 2002), it could not be subducted due to its high buoyancy. As a result, contractional deformation developed in the convergence zone and the slab broke, followed by the stepping back of the trench southward, forming the present-day Timor trough (B. Sapiie, pers. comm. 2009). This may have created the tear at shallow depth in the slab as the buoyant plateau arrived at 3 Ma. Presently, the region above the tear is marked by an aseismic zone around east Timor centred at 1268E and –98S (see the seismicity plot in Fig. 1). In addition, the observed small hole in the subducted slab below Flores is interpreted as a thinning in or ‘necking’ of the slab, which may be related to the maximum tensional stress

perpendicular to the trench in the region (Cloetingh & Wortel 1986). In order to enhance our discussion, we compare the P-wave images with the S-wave ones. In spite of relatively noisy data, the results of the S data inversion are qualitatively in good agreement with those from the P data. The slab in the upper mantle is well imaged by the S data (Figs 7 & 8). One prominent difference is the absence of the deflected slab in the mantle transition zone below the eastern Sunda arc in the S images (cf. layer anomaly maps in Figs 2 & 7 at depths of around 530– 615 km, and cross sections O –R in Figs 4 & 8). Such a difference was reported from regional seismic tomography of the NW Pacific island arcs using P- and S-wave arrival time data with similar ray path coverage (Widiyantoro et al. 1999). This implies that the stagnant lithospheric slab in the transition zone is more likely a bulk-sound structure, which is strongly supported by the results from joint inversions for bulksound and shear wavespeed (Kennett et al. 1998; Gorbatov & Kennett, 2003). For the Izu Bonin region, where trench migration has been reported, the properties of the deflected slab lying on top of the 660 km discontinuity show strong bulk-sound and weak shear signatures in contrast to the descending slab itself. A similar feature has also been observed below the Aegean region (Widiyantoro et al. 2004), where the slab is intensely distorted in the transition zone, with lower mantle penetration spatially confined to a relatively small area. The joint P- and S-inversion results also indicate that the deformed part of the slab is likely to be a bulk sound feature (Kennett et al. 1998). The difference between the shear and bulk sound signature of slabs that either deflect in the transition zone or penetrate to larger depths thus seems robust and may contain important information about the interaction of down wellings with the upper mantle discontinuities. The deflection of the slab in the transition zone below the eastern Sunda arc and the western Banda arc may resemble that observed beneath the Izu Bonin arc, the Aegean Sea as well as the Tyrrhenian Sea (Spakman et al. 1993), and can be related to the slab roll-back that has accompanied back-arc extension. Because the P-wave data sampling is much denser than the S-wave sampling, one may wonder if the differences in the P- and S-wave images are due to differences in data coverage. To answer this question we have conducted additional P-wave tomographic imaging using similar ray path coverage as the S-wave data for the Indonesian region, as conducted for the NW Pacific island arcs by Widiyantoro et al. (1999). In general, the results depict structural features similar to those from the P-wave inversion results using full data coverage given in Figures 2–6. An example of vertical

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sections across Bali through the P- and S-wave models derived using the common data coverage is given in Figure 9a, b. The P- and S-wave images depict not only the subducted Indo-Australian plate below Bali, but intriguingly also the south-dipping feature of the back-arc lithosphere directly north of Bali. The hypocentres of local earthquakes recorded by the Meteorological, Climatological and Geophysical Agency (MCGA) of Indonesia seismographic stations also form an image of southward dipping lithosphere underlying the Java Sea north of Bali (Widiyantoro & Fauzi 2005). Fault plane solutions of events at depths less than 50 km north of Bali in the period 1963–2001 compiled from several catalogues (mainly the Harvard Centroid Moment Tensor solution catalog; Global CMT Catalog 2009) generally depict thrust events (Figure 9c). We envisage that the southward dipping lithosphere underlying the Java Sea to the north of Bali may represent a south-dipping thrust due to back-arc shortening similar to that to the east (Flores). The back-arc thrust of Flores generated a large tsunamigenic earthquake in 1992. In 2004, a devastating earthquake and subsequent strong tsunami occurred in Alor to the east of Timor. This event also occurred on a back-arc thrust that may represent the eastward extension of the Flores back-arc thrust fault. The tsunami catalogue of Indonesia shows that

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tsunamigenic earthquakes occurred in the area to the north and NE of Bali in 1816 and 1979, respectively (Hamzah et al. 2000). We interpret these events to be related to the south-dipping lithosphere in the back-arc region of Bali as revealed by the seismic tomograms (Fig. 9a, b).

Concluding remarks We have presented seismic tomographic models of the 3D upper mantle velocity structure of the eastern Sunda arc from improved methods and data. The new images provide a more detailed structure of the subducted slab partly evident from the seismicity and enrich our understanding of the lithospheric processes governing its geodynamical evolution. The subducted slab seems to be continuous in the upper mantle along most of the eastern Sunda arc with some exceptions: (i) a hole in the slab in the mid upper mantle below eastern Java, and (ii) a tear in the slab in the uppermost mantle below Timor and small islands east of it. These observations may be related to the arrival of a buoyant plateau near eastern Java at c. 8 Ma and the arc – continent collision around Timor at c. 3 Ma, respectively. In addition, we also observed a ‘necking’ in the subducted slab below Flores. The observed southward dipping feature below Bali is intriguing. This feature is in excellent

Fig. 9. (a) and (b) P- and S-wave tomograms respectively derived using similar ray path coverage, and (c) CMT solutions of shallow events in the back-arc region of Bali, courtesy of Fauzi, MCGA.

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agreement with local seismicity and is also depicted clearly by P- and S-wave images derived using similar ray path coverage (Fig. 9), although some degree of smearing may exist due to the smoothing applied in the inversion. The intriguing south-dipping feature in the back-arc region to the north of Bali has caused some tsunamigenic events. Following the great Andaman– Sumatran earthquake of 26 December 2004 with its attendant devastating tsunami, there have been calls for a tsunami early warning system for the Indian Ocean. Inferences from this study also urgently call for such an early warning system to mitigate tsunami hazards not only in the fore-arc, but also in the back-arc regions of Bali and small islands to the east. The deflected slab below the easternmost part of the Sunda arc and most likely also below the Banda arc indicates that these regions have undergone slab roll-back. The deflected slab is clearly imaged in the P-wave model, but not in the S-wave model. This supports previous observations that the deflected slab in the mantle transition zone is likely a bulksound feature. Besides the excellent similarities depicted by the P- and S-wave images, the differences between the two models need to be investigated further. In order to explore the bulk and shear moduli, which have differing sensitivity to temperature and mineral composition, we need comparable high-resolution images of both P and S velocity distributions from high quality P-wave and S-wave data. They may then help to constrain the nature of processes that produce the observed variations. The next generation of subduction zone S-wave tomographic models would need more precise S-wave arrival times, such as those carefully processed by Grand (1994) for global tomography. We thank E. R. Engdahl, R. D. van der Hilst and R. Buland for the updated hypocentre and phase data set used in this study, and Fauzi for fruitful discussion especially on the seismicity around the Bali region. Thanks also go to J. Granath and W. Spakman for helpful reviews, and R. Hall for useful comments. This material is based upon work supported in part by NASA, under award NNX06AF10G. S. W. would like to thank the ITB alumni association for a Fellowship (2009/2010) to conduct research on the structure of subduction zone and tectonics of the eastern Sunda arc.

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Ductile flow in the metamorphic rocks of central Sulawesi IAN M. WATKINSON SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK (e-mail: [email protected]) Abstract: Metamorphic rocks exposed along the Palu-Koro Fault of west-central Sulawesi, Indonesia, show abundant evidence of non-coaxial ductile deformation. The deformed rocks include gneisses, amphibolites and schists, that form part of a regionally metamorphosed basement complex of Mesozoic– Precambrian Australian (Gondwanan) origin. In the Palu and Neck regions of Sulawesi, ductile shear fabrics record low-angle westward extension. Further south in the Palu valley, extension is directed towards the south and SW, along with gently-dipping ductile thrust fabrics. Vergence exceptions are common at both outcrop and kilometre scale. Cross-cutting granitic dykes place some constraint on the timing of ductile foliation formation. In the neck region of Sulawesi, it occurred before c. 44– 33.7 Ma. In the central and northern Palu valley, to the south, it occurred before 5 –3.5 Ma. The timing and orientation of non-coaxial strain precludes its origin as a result of Palu-Koro Fault activity. Instead, ductile flow occurred during either Eocene–Miocene mid-crustal extension above a metamorphic core complex, Cretaceous subduction-related deformation in the over-riding plate, or intracontinental deformation within Gondwana.

Ductile flow and associated shear zones are common in metamorphic basement rocks around the world. Examples exposing small-scale, polygenerational shear zones have been recorded from the Variscan of Spain (e.g. Carreras 2001); a regional shear zone which controls fluid flux in the Bohemian Massif (Vra´na & Ba´rtek 2005); shearing controlling melt extraction in the Canadian Cordillera (Nyman et al. 1995); both discrete and diffuse shear zones (Goscombe et al. 2006), and mid-crustal channel flow (Langille et al. 2010), in the Himalayas. Lower- to mid-crustal rocks in which these processes occur can be uplifted by orogenic folding, thrusting, oversteps/bends along major strike-slip faults, or by metamorphic core complex exhumation. Of these mechanisms, major strike-slip faults in particular can provide important information on deep crustal non-coaxial strain, because they can facilitate rapid vertical movement of their roots, which often include ductile shear zones penetrating most, or all of the lithosphere (e.g. Hanmer et al. 1992; Leloup et al. 1995; Vauchez & Tommasi 2003; Watkinson et al. 2008). The Palu-Koro fault of central Sulawesi is an active, north–south-trending high strain rate sinistral strike-slip fault, which cuts through a suite of rapidly exhumed metamorphic rocks of low to high grade (e.g. Egeler 1947; Katili 1970; Sukamto 1973; Helmers et al. 1990; Sukido et al. 1993; Walpersdorf et al. 1998; Bellier et al. 2001, 2006). It is an ideal place to examine the timing and kinematics of ductile deformation in the metamorphic basement of this under-studied area, to help understand the complex history of Sulawesi and east Indonesia.

Geological setting Sulawesi lies at the convergence of the Eurasian, Indo-Australian and Philippine tectonic plates. Its tectonic evolution has been influenced not just by broad external forces exerted by their convergence, but by a complex history of subduction, extension, ophiolite obduction and collision of continental fragments (e.g. Katili 1978; Hamilton 1979; Silver et al. 1983a, b; Hall 1996, 2002; Parkinson 1998; Calvert & Hall 2003; van Leeuwen & Muhardjo 2005). The island is composed of four elongate ‘arms’, which broadly correspond to lithotectonic units (e.g. Sukamto 1975; Hamilton 1979; Fig. 1). The north arm consists of a Neogene island arc underlain by oceanic crust, with small fragments of continental crust (Taylor & van Leeuwen 1980; Elburg et al. 2003; van Leeuwen et al. 2007). The north arm is linked to the rest of the island by the ‘neck’, a narrow, mountainous ridge largely underlain by metamorphic basement (e.g. Sukamto 1973). The east arm is dominated by a highly tectonized ophiolite, inter-thrust with Mesozoic and Cenozoic sediments (e.g. Hamilton 1979; Simandjuntak 1986; Parkinson 1991), which also crops out on the SE arm. West Sulawesi, commonly referred to as a magmatic arc of Miocene –Pliocene age, is dominated by very young granitoid intrusions (c. 14 –3 Ma) (Elburg et al. 2003) which have been intruded into a suite of Cenozoic volcaniclastic and Precambrian– Mesozoic metamorphic rocks (Fig. 2). The metamorphic basement originated in Gondwana (Bergman et al. 1996; Smyth et al. 2007; van Leeuwen et al. 2007). These continental fragments

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 157–176. DOI: 10.1144/SP355.8 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Summary of the geology of Sulawesi, showing principal structures and geographical features. Modified after Hall & Wilson (2000). MMC, Malino Metamorphic Complex; BC, Bantimala Complex.

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Fig. 2. Geological map of the Palu area. See box in Figure 1 for location. Modified after SULAROX; Sukamto (1973, 1975); Sukido et al. (1993); and van Leeuwen & Muhardjo (2005).

accreted to the SE Sundaland margin during the Middle Cretaceous (Hamilton 1979; Manur & Barraclough 1994; Parkinson et al. 1998; Hall et al. 2009). All these rocks are cut by the Palu-Koro Fault, a major, strike-slip fault which may penetrate the whole lithosphere (e.g. Brouwer et al. 1947; Hamilton 1979; Walpersdorf et al. 1998; Bellier et al. 2001). Limited existing fission-track data indicates that the Neogene intrusives and their

metamorphic hosts along the Palu Fault were rapidly uplifted during the Pliocene, and geomorphic studies indicate that uplift continues to the present day (Bellier et al. 2006). Most of the basement rocks have been subjected to regional metamorphism, and many have a thermal overprint due to the Neogene magmatism (e.g. Egeler 1947; Sopaheluwakan et al. 1995). In between these metamorphic periods, evidence of

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non-coaxial ductile deformation can also be recognized in many places. This study aims to describe its nature, and to speculate on its timing and tectonic causes.

Metamorphic basement of west-central Sulawesi Large regions of west Sulawesi’s metamorphic basement rocks crop out along the Palu valley and along the neck that connects the main island to the north arm. Names, locations and contacts between metamorphic complexes in the area vary significantly between authors. In this study, the map and nomenclature of van Leeuwen & Muhardjo (2005) is used for the area north and east of Palu, and of Sukido et al. (1993) for the area south of Palu and west of the fault valley, with small modifications following recent fieldwork (Fig. 2). For simplicity, when discussed together, all the rocks of the present study are termed the Palu metamorphic rocks, though this should not be considered a new litho-stratigraphic term. In the following descriptions of the individual units, a distinction is made between a metamorphic complex (an association of metamorphic rocks in any tectonic setting) and a metamorphic core complex (an association of metamorphic rocks exhumed by supra-crustal extension along a low-angle fault). Six metamorphic complexes have been defined from areas adjacent to the Palu-Koro Fault. From north to south, these are:

Malino Metamorphic Complex The Malino Metamorphic Complex lies at the northeastern end of the Palu fault valley, along the northern edge of Tomini Bay (Ratman 1976; van Leeuwen et al. 2007; Fig. 1). It is composed of mica schists, gneisses, greenschist, amphibolite, marble and quartzite, which formed under conditions of regional metamorphism ranging from 300– 350 8C and 0.3–0.5 GPa to 646 –617 8C and 0.75–0.96 GPa (van Leeuwen et al. 2007). Zircons from metagranitoids that intrude the Malino Metamorphic Complex indicate that it is in part of Devonian to early Carboniferous age, and includes inherited Proterozoic and Archaean zircons. Isotopes of Sr and Nd have similar characteristics to northern Australian river sediments, and geological similarities to the Bird’s Head region of New Guinea all suggest that the complex is an allochthonous terrane derived from Australia (van Leeuwen et al. 2007). A greenschist facies carapace around the Malino Metamorphic Complex core is derived from the adjacent Palaeogene Tinombo Formation. van Leeuwen et al. (2007) interpret the contact between the core and carapace to be a dome-shaped low angle normal fault, formed

during exhumation of the Cretaceous accreted metamorphic core as a metamorphic core complex during the Miocene. Those authors also describe an alternative whereby the Malino Metamorphic Complex is a Bird’s Head-derived fragment subducted beneath the north arm, and immediately exhumed during the Late Oligocene– Middle Miocene. This is supported by evidence of deformation and uplift in the western and central north arm (van Leeuwen et al. 2007).

Palu Metamorphic Complex The Palu Metamorphic Complex extends along the neck to about 40 km south of Palu, on the eastern side of the Palu-Koro Fault only (Fig. 2). It is composed of biotite schists, paragneisses, amphibolitic schists, marble and orthogneisses (Egeler 1947; Sukamto 1973). The metasediments and metagranitoids are of Permo-Triassic continental Australian age and affinity, but rocks of Sundaland and MORB affinity also occur (van Leeuwen & Muhardjo 2005). The Palu metamorphic rocks are overlain by the Late Cretaceous Latimojong Formation and the Palaeogene Tinombo Formation, a folded sequence of volcanic and marine sedimentary rocks which have been metamorphosed to greenschist facies (van Leeuwen & Muhardjo 2005).

Karossa Metamorphic Complex The Karossa Metamorphic Complex lies west of the Palu-Koro Fault, in the Lariang region. It is dominated by metapelites, and also contains MORB affinity metabasites. The Palu and Karossa Metamorphic Complexes are possibly young metamorphic core complexes (van Leeuwen & Muhardjo 2005).

Wana and Gumbassa Metamorphic Complexes The Wana and Gumbassa Metamorphic Complexes contain metamorphic rocks similar to the Palu Metamorphic Complex, and are combined with the latter by van Leeuwen & Muhardjo (2005). Nonetheless, a distinction is convenient from a geographical perspective at least. The Wana Metamorphic Complex crops in small areas on both sides of the Palu-Koro Fault about 20 km south of Palu, and more extensively west of Gimpu (Fig. 2). It is dominantly composed of schistose rocks, including mica schist, amphibole schist, quartzite and gneiss, inferred to be of Triassic age (Sukido et al. 1993). The Gumbassa Metamorphic Complex crops out west of Gimpu and in slivers along the centre and east side of the Palu-Koro Fault (Fig. 2). It is dominantly composed of gneissic rocks, including gneissic granite and diorite, gneiss and schist, inferred to

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be of Triassic –Jurassic age (Sukido et al. 1993). Ages assigned to the complexes should be viewed as provisional, as Sukido et al. (1993) provide no supporting evidence for them. Lenses of granulite, eclogite and garnet peridotite are tectonically intercalated with the lower grade metamorphic rocks along the southern parts of the Palu fault valley (Kadarusman & Parkinson 2000; van Leeuwen & Muhardjo 2005). They represent the deepest rocks exposed in the area. Helmers et al. (1990) determined rapid uplift of the garnet peridotites from peak P–T conditions of 1050–1100 8C and 1.5–2 GPa, or about 60 km depth, Kadarusman & Parkinson (2000) estimated peak P– T conditions of 1025–1210 8C and 1.9– 3.2 GPa. The eclogites may have experienced pressures of 2.8 GPa (Liou & Zhang 1995; Kadarusman & Parkinson 2000). Kadarusman & Parkinson (2000) described ductile deformation of olivine, garnet, clinopyroxene and orthopyroxene in the garnet peridotites, and inferred that these formed between garnet – lherzolite assemblage metamorphism and spinel –garnet –lherzolite assemblage metamorphism. Both of these stages preceded granulite and amphibolite facies events. Younger, brittle deformation occurred during greenschist facies retrograde metamorphism, and later, under shallow level, serpentinite-forming conditions.

Pompangeo Schist Complex The Pompangeo Schist Complex is part of Sulawesi’s north –south striking central metamorphic belt which lies east of the Palu fault valley, adjacent to the plutono-metamorphic belt of central and western Sulawesi (Fig. 1). It is composed of marble, calcareous phyllite, quartz –mica schist, phyllite, metaconglomerate, metabasic intrusions and metatuffs. West of lake Poso, quartzo-feldspathic schist and quartzite also become abundant. Metabasic and serpentinitic intercalations (but not ophiolitic rocks) increase in abundance westward. The contacts between many lithologies are defined by eastverging thrusts, parallel to a strong, compositional layering-parallel transpositional foliation, which dips WSW (Parkinson 1998). Asymmetric folds and quartz segregations in the schist show that early folding and transposition occurred during noncoaxial deformation, associated with localized mylonitic zones which indicate a top-to-the-east shear sense (Parkinson 1998). Metamorphic grade increases from east to west across the metamorphic belt. The highest grade rocks, closest to the Palu-Koro fault, have been interpreted as underthrusted accretionary complex slices above a west-dipping subduction zone, metamorphosed during the Middle Cretaceous (Parkinson 1998; Parkinson et al. 1998). The surface contact

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between the Pompangeo Schist and the plutonometamorphic rocks of central and western Sulawesi is inferred to be an east-dipping thrust, which must have formed after the Mio–Pliocene intrusion of the granitoids (Simandjuntak et al. 1991; Parkinson 1998).

Non-coaxial strain in the metamorphic rocks of the Palu-Koro fault zone Northern Palu valley Metamorphic rocks in the northern Palu valley include quartz and quartz-biotite slates, schists, amphibole schists, garnet–mica schists, and bands of foliated granitoid. Along the northwestern side of the valley, young and mostly unfoliated granitoids of the Dondo Suite dominate, and metamorphic rocks are limited to a few small relics which have a strong thermal overprint (Egeler 1947). Schists composed of acicular green amphibole, K-feldspar and plagioclase possess a strong L-S tectonite fabric. Their foliation generally dips steeply east, and mineral lineations plunge moderately towards the east and NE (Fig. 3a). Many amphiboles are euhedral, but in places form asymmetric lenticular fish, indicating non-coaxial strain (Fig. 4a). Sigma-type porphyroclasts of individual feldspar crystals or feldspar aggregates, with recrystallized tails, deflect the aligned amphiboles (Fig. 4b). Large quartzo-feldspathic masses with asymmetric tails up to 0.5 m long have a similar geometry, and may be boudins of pre-kinematic veins or dykes (Fig. 4c). Kinematic indicators along the NW Palu valley show top-to-the WSW and NW ductile thrusting, locally with a significant dextral strikeslip component, together with top to the NE and SE ductile extension (Fig. 3a, Table 1). In places where compositional banding is prominent, intense folding tends to form instead of the laminar flow of more homogeneous parts of the schist. Folds are disharmonic and often ptygmatic. On the eastern side of the valley, north of Pandere, metamorphic rocks are much more widespread. These are mapped as the Wana Complex, Gumbasa Complex, and Latimojong Formation by Sukido et al. (1993). van Leeuwen & Muhardjo (2005) class all of the metamorphic rocks on the eastern side of the valley as the Palu Metamorphic Complex. A suite of schistose rocks crops out along the road in the linear, NW-trending valley which intersects the Palu valley from the east near Bora. In the north, these possess a moderately south to SW-dipping foliation defined by alternations of layers rich in garnet, quartz and feldspar, and layers rich in aligned biotite and sillimanite. A coarse but discontinuous lineation is formed by

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(a)

I. M. WATKINSON N

(b)

N

(c)

(f)

N

N

River near Pandere (Northern Palu valley)

Bora region (Northern Palu valley)

Northwest Palu valley (Northern Palu valley)

(e)

N

(g)

(d)

N

River near Tuwa (Central Palu valley)

N Pole to foliation Stretching lineation Pole to fold axial plane Fold hinge line Fracture cleavage Pole to dyke margin Tectonic transport direction

Momipi river (Central Palu valley)

Tawaeli-Toboli road (Neck area)

Alindau river (Neck area)

Fig. 3. Equal area lower hemisphere stereonets showing the orientation of mylonitic foliation, stretching lineation, and other structural elements for regions referred to in the text. Planes are represented as poles. Black arrows indicate the tectonic transport direction where known, up-plunge (towards the centre of the stereonet) or down-plunge (away from the centre of the stereonet).

aggregates of all these minerals, which plunges to the SW. Garnet forms fractured porphyroclasts that deflect the foliation and lack parallel inclusion trails, suggesting that they predate the foliation. They possess slightly asymmetric sigma-type tails of staurolite, biotite partly replaced by sillimanite, unstrained (i.e. wholly recrystallized) quartz, and garnet fragments. Their asymmetry, and that of small plagioclase porphyroclasts, mica fish and shear bands, indicates a top-to-the SW extensional ductile fabric (Fig. 3b, Table 1). Further south along the valley, the schists are dominated by green amphibole, K-feldspar and plagioclase. A strong foliation is defined by bands rich in feldspar. Alignment of partly acicular amphiboles forms a SW-plunging mineral lineation (Fig. 3b). There is little evidence of non-coaxial strain in these rocks. However, they were observed 2 km east of the mountain upon which the stream transect north of Pandere, described below, lies. Float in the west-flowing stream is dominated by proto-mylonitic, coarse grained amphibole-feldspar gneisses and schists. It is likely that the higher parts of the mountain are composed of a strongly sheared correlative of the amphibole schist, which are being washed west into the Palu valley.

A stream transect (Fig. 5) up the side of the Palu valley north of Pandere reveals a sequence of deformed and variably mylonitic schists with a broadly north-dipping foliation and NW, north, NE and east-plunging mineral lineations (Fig. 3c). Both metamorphic grade and strain increase eastwards, and some of the most strongly mylonitic rocks in the Palu valley occur near the top (east) of the transect. These rocks are biotite–quartz – garnet schists which also contain muscovite, aegirine augite, calcite and chlorite. Garnets up to 3 mm in diameter form prominent porphyroclasts which commonly retain their euhedral shape, and deflect the surrounding foliation (Fig. 4d). Their asymmetric, sigma-type tails are composed of small quartz grains formed by bulging recrystallization of matrix quartz, and muscovite. The garnets are corroded on faces attached to their tails, and this occasionally emphasizes curved inclusion trails which are continuous with elongation of quartz in the tails, and indicate that the garnets rotated during growth, and are therefore syntectonic (Fig. 4e). Coarse ridges formed by the deflection of foliation over the garnets and their tails form a conspicuous lineation, which plunges gently to the NW

DUCTILE FLOW IN CENTRAL SULAWESI

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Fig. 4. Characteristics of ductile shear from the northern Palu valley. Latitude, longitude, view direction (if in situ) and orientation relative to ductile fabrics is shown above each image. (a) Hornblende fish and (b) feldspar porphyroclast in an amphibole schist. Top-to-the-right shear sense. Plane polarized light (PPL). (c) Asymmetric ductile boudin of quartzofeldspathic material in a large boulder. (d) Garnet porphyroclast in mylonitic garnet-mica schist. Top-to-the-right shear sense. PPL. (e) Curved inclusion trails in a garnet porphyroclast, indicating syn-kinematic origin. Top-to-the-right shear sense. PPL. (f ) Biotite fish in mylonitic garnet-mica schist. Top-to-the-right shear sense. PPL. (g) Quartz-rich asymmetric boudin train. Top-to-the-left shear sense.

(Fig. 5). Kinematic indicators such as asymmetry of sigmoidal garnet inclusion trails, asymmetry of their sigma-type tails, S –C0 fabrics and well developed

biotite fish (Fig. 4f ) all consistently indicate top-to-the SE ductile thrusting parallel to the lineation (Fig. 5, Table 1). In the central part of the

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Table 1. Summary of kinematic indicators and overall shear sense from parts of the Palu valley and neck metamorphic rocks Area

Location

Lineation

Shear-sense options

Kinematic indicators Mica Amphibole Delta Sigma Spiral Rotated fish fish porphyroclast porphyroclast inclusion inclusion trails trails

Northern Palu NW Palu valley valley

40– 058

Bora region

River near Pandere

Central Palu valley

Neck area

10– 213

West

30– 306

Centre

35– 047

East

30– 323

River near Tuwa

12– 221

Momipi river

28– 219

Tawaeli-Toboli Road

28– 260

Alindau river

West

05– 316

Centre

02– 214

East

40– 048

Thrust (top-to-WSW) Extension (top-to-ENE) Thrust (top-to-NE) Extension (top-to-SW) Thrust (top-to-SE) Extension (top-to-NW) Thrust (top-to-SW) Extension (top-to-NE) Thrust (top-to-SE) Extension (top-to-NW) Thrust (top-to-NE) Extension (top-to-SW) Thrust (top-to-NE) Extension (top-to-SW) Thrust (top-to-E) Extension (top-to-W) Thrust (top-to-SE) Extension (top-to-NW) Thrust (top-to-NE) Extension (top-to-SW) Thrust (top-to-SW) Extension (top-to-NE)

5

1

3

32 12

1

15

4

35 5 19

15

6

11

2

1

1

1

3

4

18

1

22

1

1

3 1 1

4

Lineation, representative mineral stretching lineation from each location. Kinematic indicators, observed parallel to the lineation, on

transect, ductile thrusting has a more SSW–SW direction (Fig. 5, Table 1). Very fine grained, dark grey rocks crop out at the bottom (west) of the section. A fine, slatey foliation is pervasive, and a mineral lineation defined by elongated biotite is commonly developed parallel or oblique to the foliation dip direction (Fig. 5). The rock is streaked with foliation-parallel quartz veins, typically stretched into lenticular ductile boudins joined by biotite-rich shear planes (Fig. 4g). Shear plane-foliation angles, boudin asymmetry, S –C0 fabrics, rotated opaque grains and spiral mica cleavage all indicate top-to-the SE ductile thrusting (Fig. 5, Table 1). Small, semiductile thrust faults cut across the foliation and have a similar vergence direction. These may

represent continued, broadly south directed thrusting under retrograde conditions.

Central Palu valley Gneisses and schists exposed in the central Palu valley between Tuwa and Gimpu (Fig. 2) show signs of non-coaxial strain. Migmatitic paragneisses exposed in a west-flowing stream south of Tuwa are strongly foliated by discontinuous melt lenses and alignment of biotite and a brown amphibole in dark melanosomes. Elsewhere, a gneissic banded texture is developed by more systematic segregation of quartzo-feldspathic and mafic minerals. Thicker leucocratic veins with an assemblage identical to that in the melt lenses lie parallel to the foliation. These are only weakly foliated, so were probably

DUCTILE FLOW IN CENTRAL SULAWESI

165

Kinematic indicators Spiral mica cleavage

Rotated amphibole grains

Rotated opaque grains

Biotite strain shadows

Asymmetric quartz segregations

Asymmetric folds

Overall shear-sense 0

Asymmetric Shear S– C Oblique Displaced boudinage band fabric foliation broken in quartz grains 4

1

Dextral/thrust (top-to-WSW)

4

7 8

2

Extension (top-to-SW)

2

4

4

2

5

Thrust (top-to-SE)

1

1

1

4

4

3

1

1

Thrust (top-to-SW)

2 1

Thrust (top-to-SE) 2

1

1

1

8

5

1

3

8

3

1

1

1

4

3

Extension (top-to-SW) 3 1 1

25

1

17

7

19

3

Thrust (top-to-NE) 1

1 Extension (top-to-NW)

7

6 11

Extension (top-to-SW) 2 1

8

Extension (top-to-W)

2

Extension (top-to-NE)

4

foliation-normal surfaces. Numbers represent the number of each kinematic indicator observed at each location.

injected along an existing metamorphic fabric rather than being transposed by post-anatexis shear and flattening. Melt lenses commonly form en-echelon veins. Some larger leucosome veins have been stretched into trains of ductile boudins (Fig. 6a). These also have a stair-stepping geometry. Shear bands between boudins have a shear sense consistent with that of the stair-stepping elements. Highly asymmetric intrafolial folds are common (Fig. 6b). All these criteria show a consistently down-dip, extensional shear sense parallel to a weak biotite aggregate lineation which plunges to the SW (Fig. 3d, Table 1). Straight, sharp sided dykes of unfoliated biotite granite and aplite cut the foliation (Fig. 6c), and larger bodies dominate the lower part of the river.

These post-date the regional metamorphism and ductile fabric, and are interpreted to relate to the Pliocene Dondo Suite intrusives. Garnet peridotite boulders in the stream may be from tectonic slivers within the gneiss, or may have been exhumed by and eroded from the younger intrusives. Metamorphic rocks crop out in the Momipi river near Namo. Gneissic segregation is more advanced than in the Tuwa example, biotite and garnet are much more prevalent, and kinematic indicators are more conspicuous. In addition to well developed examples of the kinematic indicators described above, discrete shear bands and large scale S –C0 fabrics extend and deflect the gneissic fabric (Fig. 6d). An aggregate mineral lineation plunges moderately to the south and SW. In contrast to the Tuwa gneisses, kinematic indicators in the Momipi

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I. M. WATKINSON

Fig. 5. Map showing features along a stream transect up the eastern Palu valley side north of Pandere. Representative structural data and shear sense indicated. For location see Figure 2. Qtz-bt, quartz–biotite; Bt-gt-act, biotite–garnet– actinolite.

river gneisses show a consistent top-to-the north and NE ductile thrust geometry (Fig. 3e, Table 1). A major brittle NE-striking fault zone composed of vertical and steeply-dipping faults cuts through the gneisses and post-metamorphic granite dykes. Oblique slickenside lineations plunge south, and individual strands truncate and juxtapose lithological units, with a small extensional component. The whole system brings large unfoliated biotite granite bodies into tectonic contact with an 8 m wide sliver of the gneisses (Fig. 6e, f). The fault zone’s orientation indicates that it may be a synthetic strand of the Palu-Koro Fault. Large, proximally-derived boulders in the Palu river north of Tuwa provide clear examples of strain style in the gneissic rocks eroded from the central part of the Palu valley. A boulder of well foliated amphibole–plagioclase gneiss contains angular and rounded fragments of amphibolite, wrapped by the dominant foliation. Highly leucocratic material, which may have been melt segregations, forms asymmetric strain shadows around these ‘mega-porphyroclasts’, and includes amphibolite fragments mechanically removed from the core (Fig. 6g). A boulder of migmatitic gneiss, similar to the Tuwa gneisses, shows well developed shear bands that deform the dominant foliation. Leucocratic material lines the shear planes, showing that a period of melt mobilization, and perhaps generation, was syn-kinematic with respect to ductile fabric formation (Fig. 6h). This does not imply that shear zone heating caused melting, simply that shear occurred in rocks which were already melting during regional metamorphism. Amphibolite breccias with an aplitic cement are common. The prevalence of massive, or weakly foliated amphibolite as clasts in both breccias and gneisses

indicates that an amphibolitic terrane was later subjected to further regional metamorphism and noncoaxial strain to produce the strongly foliated amphibole schists and gneisses in which evidence of non-coaxial strain is common.

Neck area Schists and migmatitic paragneisses are exposed in the southern and central part of Sulawesi’s ‘neck’ (Fig. 2). These are included, by van Leeuwen & Muhardjo (2005), in the Palu Metamorphic Complex. Egeler (1947) described in detail the petrography of andalusite, amphibole, augite and garnet schists and amphibolites from the TawaeliToboli road, and from the Boemboe river, close to Toribulu. Recent development of the Tawaeli-Toboli road has revealed many new outcrops, and the Alindau river, a major west-flowing river south of the Boemboe river, also provides excellent exposures. Most rocks possess a strong foliation, defined by one or more of the following fabrics: compositional banding (pelitic/psammitic alternations or schistose and gneissic segregation), mica alignment, alignment of quartzo-feldspathic segregations or melt lenses. Foliations in rocks exposed along the Tawaeli-Toboli road generally dip westwards at moderate to steep angles (Fig. 3f). Strong lineations are common on foliation surfaces, sometimes in the form of fine crenulations, otherwise as polymineralic aggregates. These plunge to the north, west and south (Fig. 3f). En-echelon quartzo-feldspathic segregations, some of which are boudins of stretched veins, mica fish, porphyroclasts and shear bands indicate top-to-the west extension, parallel to west-plunging lineations (Fig. 3f, Table 1).

DUCTILE FLOW IN CENTRAL SULAWESI

167

Fig. 6. Characteristics of ductile shear from the central Palu valley. Latitude, longitude, view direction (if in situ) and orientation relative to ductile fabrics is shown above each image. (a) Sheared felsic boudin in migmatites. Top-to-the-left shear sense. (b) Asymmetric intrafolial fold. Top-to-the-right shear sense. (c) Biotite granite dyke cutting gneissic foliation (top left-bottom right). Note hammer in centre of view for scale. (d) Large-scale shear bands (top right-bottom left) extending older gneissic foliation (top left-bottom right). Top-to-the-left shear sense. (e) Part of brittle fault zone near Namo, juxtaposing unfoliated granite with non-coaxially deformed gneiss. (f ) Interpretation of Figure 6e showing the nature and orientation of contacts. (g) Amphibolite porphyroclast in a sheared feldspar-amphibole matrix. Unoriented boulder. (h) Felsic melt within shear planes (top left-bottom right) in a migmatitic gneiss. Unoriented boulder.

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Extensive exposures along the Alindau river reveal similar, dark, fine grained rocks, often with a phyllitic, schistose, or locally gneissic appearance. A wide range of foliation orientations include moderate westerly dips, moderate to steep northerly, and moderate to shallow northwesterly dips (Figs 3g & 7). The foliation is folded by a penetrative folding fabric typified by small, open similar folds. Recumbent, tight folds with wavelengths in the order of 10 m appear to fold these fabrics. Fold axial planes lie close to the bulk foliation orientation. Additionally, parallel, locally disharmonic and ptygmatic folds are common. Strongly attenuated shear bands have developed in the hinge regions of some of these folds. Macroscopic shear sense indicators are well developed, including asymmetric folds, asymmetric quartzo-feldspathic segregations, asymmetric boudinage, and shear bands (Fig. 8a, b). In thin section, S–C0 fabrics and an oblique foliation in recrystallized quartz segregations indicate a shear sense consistent with the macroscopic structures. Overall, kinematic indicators show top-to-the NW ductile extension in the lower (west) parts of the river, top-to-the SW ductile extension in the middle part of the river, and top-to-the NE ductile

extension in the upper (east) parts of the river (Figs 3g & 7, Table 1). However, in all cases there are a significant number of kinematic indicators showing an opposite (thrust) shear sense. This may be due in part to re-orientation caused by post-simple shear penetrative folding. All shear bands which extend the folds have a consistent shear sense. Quartz grains in quartzo-feldspathic segregations and boudin trains generally have a strain free appearance, indicating complete recrystallization and annealing (Fig. 8c). Grain boundaries are irregular and lobate, due to grain boundary migration recrystallization (Fig. 8d), a process characteristic of temperatures above 500 8C. Locally, chessboard sub-grains are developed, indicative of temperatures exceeding c. 650 8C (Stipp et al. 2002). A float sample from the Alindau river contains sheared quartzo-feldspathic segregations within a schistose foliation, both of which have been cut at an angle of c. 208 by a thin granitic dyke. This has itself been stretched and slightly sheared by continued deformation, indicating that it was intruded synor inter-kinematically (Fig. 8e, f ). All metamorphic fabrics have been cut by younger unfoliated granitic dykes (Fig. 8g).

Fig. 7. Map showing features along a stream transect up the Alindau river, on the western side of the neck. Representative structural data and shear sense indicated. For location see Figure 2. Bt, biotite.

DUCTILE FLOW IN CENTRAL SULAWESI

Constraints on the causes and timing of non-coaxial deformation There are two broad scenarios for the formation of ductile shear fabrics in the Palu area. In the first, fabrics formed before, and are unrelated to Neogene deformation and uplift, and were subsequently exhumed, undergoing static retrograde metamorphism that passively overprinted their deformation fabrics. Alternatively, fabrics formed during their Neogene exhumation. No data directly date the timing of metamorphism or ductile deformation in the Palu valley. However, the tectono-magmatic evolution of the area is becoming increasingly well known. A summary is provided here in order to place constraints on the causes and timing of ductile fabric development.

Origin of the western metamorphic rocks Many of the metamorphic rocks of Sulawesi’s basement are of Mesozoic, Palaeozoic, and probably also Precambrian origin. A fragment of Gondwanan continental crust probably lies beneath western Sulawesi (Elburg et al. 2003; van Leeuwen et al. 2007). This fragment, the Argo block, which may underlie much of east Java as well as west Sulawesi, bears zircons with Archaean –Cambrian age peaks (Bergman et al. 1996; Smyth et al. 2007; van Leeuwen et al. 2007) which correspond to periods of orogenic growth characteristic of the Gondwana margin. Potassic calc-alkaline to ultrapotassic intrusions within the metamorphic rocks of west Sulawesi yield Sr, Nd and Pb isotopic ratios which indicate that they have incorporated crust of Australian origin, presumably the metamorphic basement (Priadi et al. 1993, 1994; Bergman et al. 1996; Polve´ et al. 1997, 2001; Elburg & Foden 1999; Elburg et al. 2003). This continental fragment rifted from the northern margin of Australia during the late Jurassic (Powell et al. 1988; Hall et al. 2009), and, alongside others in the East Java Sea, SE Kalimantan and the southern Makassar Straits, accreted to the SE Sundaland margin during the Middle Cretaceous (Hamilton 1979; Manur & Barraclough 1994; Parkinson et al. 1998; Hall et al. 2009). The Palu, Malino and Karossa metamorphic complexes of western Sulawesi are composed of pieces of this continental fragment (van Leeuwen & Muhardjo 2005; van Leeuwen et al. 2007). They form a distinctive belt dominated by staurolite and sillimanite þ andalusite þ cordierite-bearing amphibolites (Egeler 1947). The age of the metamorphic rocks (Archaean–Cambrian) provides an upper age limit for ductile deformation in these rocks. It is possible that deformation observed in

169

the Palu metamorphic rocks occurred during an intracontinental tectonic event prior to the breakup of Gondwana. Significant folding and local faulting during younger tectonic events could explain the diverse orientations of observed ductile fabrics (Figs 3 & 9).

The Central Sulawesi Metamorphic Belt Me´lange, ultramafic rocks, radiolarian cherts and high pressure metamorphic rocks, including blueschists and eclogites, make up a dismembered accretionary complex within, and around the eastern margin of the metamorphic complexes. These rocks are exposed in the Bantimala Me´lange Complex, the Latimojong mountains, and the Pompangeo Schist Complex of the Central Sulawesi Metamorphic Belt (Sukamto 1975; Parkinson 1991, 1998; Bergman et al. 1996; Wakita et al. 1996). As glaucophane-bearing rocks associated with massive ultramafic bodies (Brouwer et al. 1947) and high shear strains (Parkinson 1998), these units are characteristic of rocks formed in a subduction zone. They yield Aptian –Albian white mica and whole rock K –Ar ages, indicating that metamorphism associated with subduction occurred during the Middle Cretaceous (Parkinson 1998). An east-verging ductile shear sense has been described from these rocks (Parkinson 1998), and it is possible that ductile fabrics in the Palu metamorphic rocks represent contemporaneous deformation in the middle crust of the over-riding plate. In particular, west to south-verging thrusting in the metamorphic rocks of the Palu valley (Fig. 9) could represent back-thrusting in such a setting. Similarly directed extensional fabrics, particularly along the neck and near Bora and Tuwa, may have originated in the same thrust setting, but were later rotated into their present orientation. Mesozoic shallow marine or continental margin sedimentary rocks inter-thrust with the high pressure metamorphic belt represent the parental material of the complex, which may be a microcontinental fragment that was incompletely subducted (Wakita et al. 1996; Parkinson 1998; Parkinson et al. 1998). Continental arc magmatism above this Cretaceous subduction system is probably represented by granitoids of central Kalimantan, which yield Barremian to Cenomanian ages (Pieters & Supratna 1990), and not the magmatic arc of western Sulawesi, which is of Neogene age (see below). A deep marine fore-arc basin west of this NW-dipping subduction zone was filled with clastic turbidites of the Latimojong Formation during the Late Cretaceous (van Leeuwen 1981; Hasan 1991; Sukamto & Simandjuntak 1983; Bergman et al. 1996; Calvert 2000). Both the Bantimala Me´lange Complex and the Pompangeo Schist

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Fig. 8. Characteristics of ductile shear from the neck metamorphic rocks. Latitude, longitude, view direction (if in situ) and orientation relative to ductile fabrics is shown above each image. (a) Biotite schist boulder showing sheared, lenticular quartz segregations, some with recrystallized tails. Weak shear bands dip to the right. Top-to-the-right shear sense. (b) Rounded, quartzofeldspathic ductile boudin with thin, recrystallized tail, in amphibole–biotite schist. (c) Quartz segregation in thin section, showing recrystallization of quartz. Crossed polars. Top-to-the-right shear sense.

DUCTILE FLOW IN CENTRAL SULAWESI

171

Complex are unconformably overlain by Albian to Cenomanian radiolarian cherts (Silver et al. 1983a; Wakita et al. 1996; Parkinson 1998). This provides an upper age constraint on the back-thrust hypothesis for the Palu metamorphic rocks described above.

Palaeogene events

Fig. 9. Summary map showing orientation, type, and minimum age constraints for shear in the Palu metamorphic rocks. Ages refer to dykes which cross-cut the sheared fabric. See text for details and references. See Figure 1 for location.

Late Cretaceous and Paleocene to Eocene deposits in western Sulawesi are separated by an unconformity, indicating that deformation and/or uplift occurred at the end of the Cretaceous (van Leeuwen & Muhardjo 2005). The Makassar Straits opened during the Middle Eocene, separating Sulawesi from Borneo, and the Mesozoic subduction complex from its magmatic arc (Hamilton 1979; Situmorang 1982; Cloke et al. 1999). Half graben trending NE–SW formed in NW Sulawesi from the Middle to Late Eocene (Calvert & Hall 2003). Localized subductionrelated volcanism in western Sulawesi had ceased by the Oligocene or early Miocene (Elburg et al. 2003; van Leeuwen & Muhardjo 2005; van Leeuwen et al. 2007), as the continental margin changed from active subduction to dominantly strike-slip (Rangin et al. 1990; Hall 1996, 2002). Muscovite from a two-mica granite dyke near Tompe, on the western side of Sulawesi’s neck, yielded K –Ar ages of 33.4 + 0.2 and 33.7 + 0.7 Ma (Elburg et al. 2003). These samples originated about 15 km north of the Alindau river, where similar dykes have been observed to cut the metamorphic fabric (Fig. 8g), and to be late synkinematic with respect to ductile shear (Fig. 8e, f ). Assuming the samples dated by Elburg et al. (2003) are the same age as the Alindau river dykes, and have the same relationship to the metamorphics, then the variably-oriented extensional shear observed in Alindau river schists must have occurred at, or shortly before c. 33.7 Ma (Fig. 9). Hornblende from a quartz diorite, probably collected from the same Tawaeli-Toboli road across the neck north of Palu described above, yielded a K –Ar age of 44.0 + 1.0 Ma (Elburg et al. 2003). This is a minimum age for top-to-the-west extensional shear in schists exposed along this road (Fig. 9). Isotopic ages from the Malino Metamorphic Complex indicate that it cooled through the K –Ar closure temperature for hornblende (c. 500 + 50 8C)

Fig. 8. (Continued) (d) Grain boundary migration between two quartz grains. Crossed polars. (e) Float sample showing sheared quartz-biotite schist (left-right foliation) cut by a weakly sheared biotite granite vein (top left-bottom right), interpreted to be late syn-kinematic. (f) Interpretation of (e), showing strongly sheared quartz segregations (blue) and weakly sheared granite vein (pink). (g) Granite dyke (top to bottom) cutting sheared gneiss (left-right foliation).

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at 33.0 + 1.2 Ma to 9.30 + 0.6 Ma; and for muscovite (c. 350 + 50 8C) at 19.7 + 0.2 Ma to 14.1 + 0.2 Ma (van Leeuwen et al. 2007). van Leeuwen et al. (2007) interpret these ages as recording uplift of the metamorphic complex. Since the granitoid dykes of the neck were intruded during the late Eocene, it is also possible that the neck metamorphic rocks were uplifted at a similar time. A Middle Eocene –early Oligocene, possibly northwardyounging sequence of mid-crustal, low-angle extension, decompression melting and dyke-intrusion, and subsequent exhumation is therefore possible for both the Malino Metamorphic Complex, and the northern part of the Palu Metamorphic Complex in the neck.

Neogene collisions A large ophiolite body was obducted onto eastern Sulawesi during the Neogene. It may have originated in the Banda Sea (e.g. Katili 1978; Hamilton 1979), in the Gorontalo Basin (Silver et al. 1983a), or near the northern margin of Australia (Mubroto et al. 1994). The age of the oceanic crust comprising the ophiolite is not well known (e.g. Mubroto et al. 1994; Parkinson 1998). Biostratigraphy indicates an Early Cretaceous age (Simandjuntak 1992), while K – Ar ages from the mafic rocks range from Middle Cretaceous to Miocene (Binsil & Batusimpang, reported in Simandjuntak 1986; Mubroto et al. 1994). It is likely that the ophiolite is composite, partly explaining the wide age range (Parkinson 1998). K –Ar isotopic dating of hornblende from amphibolites of the metamorphic sole below the ophiolite indicate that it was thrust onto the Pompangeo Schist Complex during the Oligocene (Parkinson 1996, 1998). Continental fragments to the east of Sulawesi, including the Banggai-Sula and Buton-Tukang Besi microcontinents, were sliced from the Australian northern continental margin in New Guinea, and travelled westwards between 1300 and 2500 km (e.g. Visser & Hermes 1962; Hamilton 1979; Silver & Smith 1983; Pigram et al. 1985; Garrard et al. 1988). These fragments collided with, and were thrust below, the obducted ophiolite (Hamilton 1979; Silver et al. 1983a). The timing of this collision is debated, and may have begun in the Middle Miocene or earlier (Wilson & Moss 1999), during the Middle Miocene (Ku¨ndig 1956; Sukamto & Simandjuntak 1983; Simandjuntak 1986), Middle Miocene to Pliocene (Garrard et al. 1988), or during the Late to latest Miocene (Hamilton 1979; Davies 1990; Longley 1997; Kadarusman et al. 2004). These collisions caused local deformation in the east arm of Sulawesi. Additionally, numerous events across Sulawesi and beyond have been attributed to the collisions. These

include overthrusting of the Central Sulawesi Metamorphic Belt onto west Sulawesi (Simandjuntak & Barber 1996), magmatism in West Sulawesi and the opening of the Bone Gulf (Bergman et al. 1996), inversion in east Kalimantan, including in the Kutai Basin (e.g. van de Weerd & Armin 1992; Cloke et al. 1997; Longley 1997; McClay et al. 2000). If it is accepted that the collision caused such far-field effects, it could be argued that non-coaxial strain in the Palu metamorphic rocks was, in part, an early result of the collision. Although this model would predict west-directed thrusting in the upper crust, deformation in the middle to lower crust might be absent or coaxial, inconsistent with observed ductile fabrics. Thermal subsidence in NW Sulawesi continued through the Oligocene and Miocene, with no angular unconformities or major breaks in sedimentation (Calvert & Hall 2003). The absence of Miocene syn-orogenic sediments in the area indicate that non-deformation, or extension was dominant in the west at this time (Calvert & Hall 2003). Conversely, at the western end of the north arm, an angular unconformity above an intensely folded and thrusted Palaeogene succession is synchronous with the 23 –11 Ma uplift of the Malino Metamorphic Complex, indicating a major localized early Miocene tectonic event in this area (van Leeuwen et al. 2007). This event may also have affected the metamorphic rocks of the neck, and low-angle extensional fabrics described herein from the neck region north of Palu (Fig. 9) may be related to metamorphic core complex development proposed by van Leeuwen et al. (2007). However, extension in the Malino Metamorphic Complex must have been directed north or south, that is, normal to its east– west elongated dome. This is inconsistent with kinematic observations in the neck showing broadly west-directed extension.

Neogene intrusive rocks and exhumation Neogene high potassium calc-alkaline magmatism, starting at c. 13– 14 Ma, was widespread in western Sulawesi (van Leeuwen 1981; Polve´ et al. 1997; Harahap et al. 1999; Elburg et al. 2003). These rocks probably formed in an extensional setting (Yuwono et al. 1988; Priadi et al. 1994; Polve´ et al. 1997; Macpherson & Hall 1999; Elburg et al. 2003), and were sourced from midcrustal continental rocks of Australian affinity (Bergman et al. 1996; Elburg & Foden 1999; Elburg et al. 2003), not from the lower crustal granulites, which show little evidence of anatexis (Helmers et al. 1990). The Neogene magmatism included widespread potassic to ultra-potassic volcanic and volcaniclastic sediments, minor felsic volcanic rocks,

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granodiorites and granitic plutons, stocks and veins. No intermediate or mafic bodies have been described (Polve´ et al. 2001). In west Sulawesi, two potassic suites can be identified (van Leeuwen & Muhardjo 2005): a Miocene high K suite, and a Pliocene potassic calc-alkaline suite. All potassic calc-alkaline granitoids of the Palu area were intruded between 5 and 3.4 Ma, and geochemically similar rhyolites erupted at 1.9 Ma (Polve´ et al. 1997). Intrusions into the metamorphic rocks formed a post-regional metamorphism low-pressure overprint in many of the metamorphic complexes in the Palu area, and intrusions and metamorphic rocks were both affected by late low-grade metamorphism (Egeler 1947; Helmers et al. 1990). Thin, straightsided dykes of biotite granite and two mica granite cut across the ductile foliation of metamorphic rocks in the central Palu valley (Fig. 6c), and include xenoliths of the sheared rocks. The dykes show very limited evidence of ductile deformation, and are clearly post-tectonic with respect to the ductile fabric. An aplite dyke dated by Polve´ et al. (1997) from Kulawi, north of Gimpu in the Palu valley, yielded a whole rock K–Ar age of 3.49 + 0.10 Ma. This is a minimum age for the top-to-the-north and NE ductile thrusting in the metamorphic rocks nearby (Fig. 9). Granitic intrusions at Tuwa and Gimpu, nearby, yielded a whole rock K–Ar age of 5.08 + 0.11 Ma and a biotite K–Ar age of 3.95 + 0.19 Ma respectively (Polve´ et al. 1997). A minimum age for the top to the SW ductile extension at Tuwa is therefore c. 5 Ma (Fig. 9). The Palu-Koro, Lawanopo and Matano faults of central Sulawesi are considered to form the small circle around a rotation pole about which north and east Sulawesi rotates clockwise (Hamilton 1979; Silver et al. 1983b; Surmont et al. 1994). Based on geological reconstructions, about 250 km of sinistral slip along these structures may have occurred since their initiation at about 5 Ma (Silver et al. 1983b). Since non-coaxial strain in the Palu metamorphic rocks certainly occurred before c. 3.5 Ma, and probably before c. 5 Ma, it is unlikely that it is the result of strike-slip movement along the Palu-Koro Fault. The absence of steeply-dipping ductile strike-slip fabrics, or a consistent obliquity in low-angle fabrics, also suggests that the observed ductile strain is not due to the PaluKoro Fault. The main expression of this structure at the surface is an array of steeply-dipping brittle fault zones, indicating that its mid- or lower-crustal roots have not yet been exhumed. Thick-skinned folding and thrusting inverted the Palaeogene basins of NW Sulawesi towards the end of the early Pliocene (Calvert & Hall 2003). This was the result of a tectonic event which continued through the Pleistocene, formed

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the present-day 3 km high mountains in Sulawesi’s central and neck regions, and led to the deposition of widespread, thick syn-orogenic sediments above a regional unconformity (Bergman et al. 1996; Hall 2002; Calvert & Hall 2003; Fraser et al. 2003). Although Calvert & Hall (2003) show that thrusting involved the basement, crystalline rocks were by that time already at a relatively shallow crustal level (see their Fig. 4), and it is likely that deformation was limited to the upper crust, incompatible with the high temperature ductile fabrics observed in the Palu metamorphic rocks. It is therefore considered that this thrusting event was not responsible for ductile non-coaxial deformation in the metamorphic rocks. Fission track ages from granitoids in central Sulawesi indicate rapid uplift (200 –700 m/Ma) between 7–5 and 2 Ma (Bergman et al. 1996; Bellier et al. 1998). Unpublished data by van Leeuwen et al. suggest cooling of the Palu metamorphic rocks at the same time. Exhumation of the Palu metamorphic rocks by the later stages of metamorphic core complex development, alongside voluminous granite magmatism, may have contributed to the rapid elevation of western and central Sulawesi during the late Neogene (van Leeuwen & Muhardjo 2005).

Conclusions Metamorphic rocks along the Palu-Koro Fault show abundant evidence of non-coaxial ductile strain. Low-angle extension, directed mostly towards the west, occurred before c. 44 –33.7 Ma in the Palu Metamorphic Complex of the neck area. This is before the Middle Miocene onset of cooling in the Malino Metamorphic Complex of the north arm (van Leeuwen et al. 2007). However, it is possible that it represents early, possibly aborted, westdirected extension in the middle crust as a precursor to this later event. Top-to-the SW to top-to-the SE ductile thrusting and extension in the metamorphic complexes of the northern Palu valley, and top-to-the SW extension and top-to-the NE thrusting in the central Palu valley occurred before 5– 3.5 Ma. This may be due to the same event, or may be an older inherited fabric. It is more likely that west-verging thrusting in this area was related to deformation in the over-riding plate of the Cretaceous subduction zone which formed the Central Sulawesi Metamorphic Belt to the east. However, the possibility remains that any or all of the deformation occurred during intracontinental deformation within Gondwana during the Mesozoic, Palaeozoic or Precambrian. What is clear is that none of the observed non-coaxial ductile strain fabrics can be correlated to the Palu-Koro Fault,

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and it is likely that its deep crustal roots, if they exist at all, have not yet been exhumed. I am grateful to A. M. Surya Nugraha and Benjamin Sapiie at Institut Teknologi Bandung, Indonesia, and to Mr Darwin Sumang of the Government Office of Culture and Tourism in Palu, for assistance with field work, and to Theo van Leeuwen, Mike Cottam and Robert Hall for insightful discussions and comments on the text. Steve Calvert and Nick Timms are thanked for their constructive reviews which considerably improved the manuscript. This work was funded by the SE Asia Research Group.

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Basement character and basin formation in Gorontalo Bay, Sulawesi, Indonesia: new observations from the Togian Islands M. A. COTTAM1*, R. HALL1, M. A. FORSTER2 & M. K. BOUDAGHER-FADEL3 1

SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey, UK

2

Department of Earth Sciences, Australian National University, Canberra, ACT, 0200, Australia 3

Department of Earth Sciences, University College London, London, UK *Corresponding author (e-mail: [email protected])

Abstract: We present a new stratigraphy for the Togian Islands, Sulawesi, and interpret the age, character and evolution of Gorontalo Bay. At its western end the bay is underlain by continental crust. The central part is underlain by Eocene to Miocene oceanic and arc rocks, although the area south of the Togian Islands could have continental crust of the Banggai-Sula microcontinent thrust beneath this and the East Arm ophiolite. Gorontalo Bay was not a significant deep bathymetric feature before the Miocene. Field relationships indicate a latest Miocene to Pliocene age for inception of the basin. Medium-K to shoshonitic volcanism in the Togian Islands is not due to subduction but reflects crustal thinning and extension in the Pliocene and Pleistocene, causing the underlying mantle to rise, decompress and melt. Extension is continuing today and is probably the cause of volcanism at Una-Una. Volcanic activity migrated west with time and volcanic products have been offset by dextral strike-slip displacement along the Balantak Fault. Extension and subsidence was driven by rollback of the subduction hinge at the North Sulawesi Trench with a possible contribution due to flow of the lower crust.

Gorontalo Bay is one of the most enigmatic basins in East Indonesia. It is relatively deep with water depths up to 2000 m, and Hamilton (1979) showed up to five kilometres of sediment in its western depocentre. It is surrounded by land on three sides and receives large volumes of sediment from nearby mountains up to three kilometres high. Miocene carbonates are widespread in these areas (van Leeuwen & Muhardjo 2005) and suggest that the deep basin formed since their deposition but the timing and mechanism of basin inception remain unclear. The nature and age of the crust beneath Gorontalo Bay is also unknown. To the north, the North Arm of Sulawesi is interpreted as a volcanic arc built on Eocene oceanic crust (Taylor & van Leeuwen 1980; Elburg et al. 2003; van Leeuwen & Muhardjo 2005). In contrast, at the western end of Gorontalo Bay, there are two kilometre high mountains with young metamorphic ages and evidence of continental crust, Miocene extension and core complex formation (Sukamto 1973; Elburg et al. 2003; van Leeuwen et al. 2007). To the south, the East Arm of Sulawesi comprises ophiolitic rocks of the East Sulawesi Ophiolite (Simandjuntak 1986; Monnier et al. 1995; Bergman et al. 1996; Parkinson 1998; Kadarusman et al. 2004). Silver et al. (1983b) suggested that Gorontalo Bay was a fore-arc basin,

underlain by ophiolitic crust equivalent to the East Arm ophiolite, situated in front of the North Arm volcanic arc that has been thrust south onto the Banggai-Sula microcontinent. The Togian Islands, situated in the centre of Gorontalo Bay (Fig. 1), offer a unique opportunity to investigate aspects of the basin’s origin and evolution. The archipelago forms a broadly WSW–ENE trending ridge that continues to the west as a submarine feature. Geological maps of the islands show igneous rocks and contrasting interpretations of them. Ku¨ndig (1956) reported andesitic intrusive rocks in the central islands, and older ophiolitic rocks in the eastern islands – suggesting a possible link to the East Sulawesi Ophiolite. In contrast, Rusmana et al. (1982) reported widespread tuffs and sedimentary formations of Mio-Pliocene age. The volcanic rocks could therefore be part of the ophiolite, could form part of the North Arm volcanic arc, or could be subduction-related products that predate the collision (Garrard et al. 1988; Davies 1990) of the Banggai-Sula microcontinent with the East Arm. The Togian Islands are also close to the isolated active volcano of Una-Una, just NW of the Togian archipelago, which has a K-rich chemistry and erupted violently in 1983 (Katili et al. 1963; Katili & Sudradjat 1984). It is not a typical subduction

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 177–202. DOI: 10.1144/SP355.9 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Tectonostratigraphic provinces of Sulawesi. Modified after Hall & Wilson (2000), Calvert (2000) and van Leeuwen & Muhardjo (2005).

volcano in position (about 200 km above the Benioff zone) and, if related to this subduction, is unusual in being the only volcano. We present a stratigraphy for the Togian Islands based on new field observations and dating. In many cases dating was restricted by the intense tropical weathering typical of SE Asia, and/or a lack of datable material. We combine these new data with earlier studies and observations of the physiography, bathymetry and seismicity of the northern Sulawesi region, to elucidate the Cenozoic history of Gorontalo Bay.

Tectonic setting Sulawesi comprises a complex association of magmatic arcs, metamorphic rocks (varying in grade from low to high), ophiolites and microcontinental fragments that have been variously assembled and deformed during the Late Mesozoic and Cenozoic (e.g. Audley-Charles 1974; Hamilton 1979; Hall 2002). It has been subdivided into four tectonostratigraphic terranes separated by major faults (e.g. Hamilton 1979). The composition of the terranes surrounding the study area is described below. Following recent studies (e.g. Calvert 2000; van Leeuwen & Muhardjo 2005; van Leeuwen et al. 2007) we do not use the term Western Sulawesi Plutono-Volcanic Arc Terrane. Instead, we follow

previous authors in separating this ‘terrane’ in to two different entities based on the recognition of significant differences in age and character of rocks (e.g. Taylor & van Leeuwen 1980; Calvert 2000; Elburg et al. 2003). We adopt the terms Western Sulawesi Province and Northern Sulawesi Province (e.g. van Leeuwen et al. 2007; see Fig. 1). The position of the boundary between these provinces remains uncertain (Elburg et al. 2003).

Western Sulawesi Province The Western Sulawesi Province (Fig. 1) represents a continental margin segment (van Leeuwen et al. 2007). It has a metamorphic basement that includes the Malino and Palu Metamorphic Complexes, exposed at the NW and SW corners of Gorontalo Bay respectively (Elburg et al. 2003; van Leeuwen et al. 2007). These rocks form part of an arcuate zone of dismembered accretionary complexes (Parkinson 1998) and continental fragments, metamorphosed in the mid-Cretaceous during emplacement along the SE margin of Sundaland by NW directed subduction (Parkinson 1998). The basement is overlain by weakly metamorphosed Upper Cretaceous sedimentary rocks of the Latimojong Formation, which are in turn overlain by a sequence of weakly metamorphosed Palaeogene sedimentary rocks and subordinate volcanic rocks belonging to the ‘Older Series’ of Elburg

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et al. (2003). The exact nature of the contact (depositional or faulted) is not known. Close to the study area these rocks include the Tinombo Formation (Brouwer et al. 1947), fore-arc basin sediments characterized by a transition from syn-rift sedimentation to platform carbonates and deeper marine sedimentation between the Late Eocene and Middle Miocene (Coffield et al. 1993; Wilson & Bosence 1996; Calvert 2000). The contemporaneous Tinombo Formation volcanic rocks (c. 51 to 17 Ma) range from basalt to rhyolite and include dykes, volcanic piles and co-magmatic intrusive stocks (Elburg et al. 2003). Intrusive and extrusive rocks of the ‘Younger Series’ (Elburg et al. 2003) include an acidic high-K calc-alkaline (CAK) suite of plutons (Kavalieris et al. 1992) and comagmatic volcanic rocks (van Leeuwen et al. 1994; Elburg et al. 2003), and a high-K calc-alkaline, shoshonitic and ultra-potassic alkaline (HK) suite of dykes, small stocks and less common extrusive rocks (Elburg et al. 2003).

Northern Sulawesi Province The Northern Sulawesi Province (Fig. 1) comprises a dominantly tholeiitic Tertiary volcanic arc built on Eocene oceanic crust (Taylor & van Leeuwen 1980; Elburg et al. 2003; van Leeuwen & Muhardjo 2005). Volcanism was driven by the northward subduction of Indian Ocean lithosphere beneath the North Arm (e.g. Hall 1996, 2002; Rangin et al. 1997). The Papayato Volcanic rocks are the products of this arc, a bimodal suite of mafic and felsic volcanic rocks cut by co-magmatic stocks of gabbro and diorite (Trail et al. 1974; Kavalieris et al. 1992; van Leeuwen et al. 1994; Elburg et al. 2003) belonging to the ‘Older Series’ of Elburg et al. (2003). Limited isotopic and palaeontological ages suggest a Middle Eocene to earliest Miocene age (van Leeuwen et al. 2007) making them the broad age equivalent of the Tinombo Formation in the Western Sulawesi Province. However, contrasting volcanic– sedimentary proportions suggest that they were formed in different tectonic environments (van Leeuwen et al. 2007). The Papayato Volcanic rocks are overlain by a thick series of Neogene volcanic rocks and volcaniclastics of calc-alkaline composition and cut by co-magmatic intrusives (‘CA Suite’ of Polve´ et al. 1997), which are accompanied by marine sedimentary rocks (Kavalieris et al. 1992) that include wellbedded shallow marine sediments and limestones of Early to Middle Miocene age (e.g. Sukamto 1973; Norvick & Pile 1976; Ratman 1976). All these rocks are cut by Neogene volcanic rocks belonging to the ‘Younger Series’ (Elburg et al. 2003). They include andesitic and dacitic stocks, dykes and

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epiclastic rocks of the calc-alkaline ‘CA Suite’, and associated Early –Middle Miocene marine sediments (Elburg et al. 2003).

East Sulawesi Ophiolite The East Sulawesi Ophiolite (Fig. 1) comprises a sequence of dunite, lherzolites and harzburgites, ultramafic cumulates, layered gabbros, isotropic gabbros, sheeted dykes and basaltic pillows and lavas (e.g. Simandjuntak 1986; Parkinson 1991, 1998). Field mapping (Kadarusman et al. 2004) and geophysical studies (Silver et al. 1978) suggest an abnormally large reconstructed stratigraphic thickness of at least 15 km. The origin of the East Sulawesi Ophiolite has been variously attributed to a typical mid-oceanic ridge (e.g. Soeria-Atmadja et al. 1974; Simandjuntak 1986), supra-subduction zone (Monnier et al. 1995; Bergman et al. 1996; Parkinson 1998) and oceanic plateau settings (Kadarusman et al. 2004). K –Ar dating of the ophiolite ranges in age from Cretaceous to Eocene (Simandjuntak 1986). They are interpreted to reflect Cretaceous, specifically Cenomanian, ocean floor with younger seamounts (Simandjuntak 1986). K –Ar dating (Parkinson 1998) has been interpreted to suggest intra-oceanic thrusting of the ophiolite at c. 30 Ma.

Microcontinental fragments The Banggai-Sula block (Fig. 1) has a basement of Palaeozoic or older metamorphic rocks intruded by Permo-Triassic granites associated with acid volcanic rocks. These rocks are overlain by undated, probably Lower Jurassic, terrestrial sediments and by Jurassic and Cretaceous marine shales and limestones. In the western parts of the islands there are Eocene to Neogene limestones (Garrard et al. 1988; Supandjono & Haryono 1993; Surono & Sukarna 1993). The block is a continental fragment derived from northern Australia (e.g. Audley-Charles et al. 1972; Hamilton 1979; Pigram et al. 1985) which collided with a subduction margin represented by the ophiolites and associated rocks of East Sulawesi. Hamilton (1979) suggested it was sliced from New Guinea and carried westward along a strand of the Sorong Fault system and this view has become widely accepted and incorporated in many tectonic models (e.g. Pigram et al. 1985; Garrard et al. 1988; de Smet 1989; Daly et al. 1991; Smith & Silver 1991; Hall et al. 1995; Hall 1996, 2002). The collision is generally thought to have occurred in the Neogene (Simandjuntak & Barber 1996) but a wide range of ages has been suggested including Late Oligocene or Early Miocene (Milsom et al. 2001), within the Miocene (Hamilton 1979),

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Early to Middle Miocene (Bergman et al. 1996), Middle Miocene (Sukamto & Simandjuntak 1983; Simandjuntak 1986), Middle Miocene to Pliocene (Garrard et al. 1988) and Late Miocene (Silver et al. 1983b; Davies 1990; Smith & Silver 1991; Parkinson 1998). Buton–Tukang Besi has been suggested to be another microcontinental fragment (Hamilton 1979) that collided in the Early or Middle Miocene (Fortuin et al. 1990; Smith & Silver 1991), after strike-slip faulting sliced it from New Guinea. Although these microcontinents are small, their collisions are often interpreted to be responsible for widespread deformation in Sulawesi and Borneo. Westward thrusting of the central Sulawesi metamorphic belt, a foreland fold and thrust belt in west Sulawesi, deformation in the Makassar Straits, deformation in the Meratus Mountains, and inversion in the Kutei basin have been attributed to the collision (e.g. van de Weerd & Armin 1992; Coffield et al. 1993; Simandjuntak & Barber 1996; Pubellier et al. 1999; McClay et al. 2000). Many authors suggest the collision, or collisions, followed westward subduction of ocean lithosphere (e.g. Garrard et al. 1988) interpreted to have produced a magmatic arc in West Sulawesi (e.g. Hamilton 1979; Parkinson 1991) or alternatively postcollisional magmatism (e.g. Bergman et al. 1996; Polve´ et al. 1997; Elburg et al. 2003). The age of collision is difficult to determine and could vary within Sulawesi. It requires dating of terrestrial clastic rocks (‘Celebes Molasse’) that rest unconformably on deformed sedimentary, metamorphic and ophiolitic rocks. In the East Arm Umbgrove (1938) reported a Lower Miocene unconformity, Brouwer et al. (1947) recorded isoclinal folding of Early to Middle Miocene age, and Ku¨ndig (1956) interpreted a Middle Miocene orogenic phase followed by molasse sedimentation and later Pliocene folding. Hamilton (1979) reported that ‘lower Miocene strata are fully involved in the imbrication and upper Miocene clastic rocks were derived from the thrust belt’. Other authors have reported Middle Miocene folding and thrusting (e.g. Audley-Charles et al. 1972; Audley-Charles 1974; Katili 1978; Parkinson 1991). Surono (1995) suggested that conglomerates from the SE Arm are the oldest Lower to Middle Miocene parts of the Langkowala Formation which rests unconformably upon the ophiolite. In Buton, Smith & Silver (1991) interpreted a deformed complex including Upper Eocene or Lower Oligocene pelagic limestones to be overlain by Lower Miocene conglomerates, but because of the lack of ophiolite detritus interpreted the conglomerates to be the product of erosion associated with slicing of the block from New Guinea rather than collision. They suggested that separate microcontinents may

have collided with East Sulawesi or that a single large microcontinent may have been fragmented during oblique collision. Recent work has cast doubt on the existence of a subduction-related volcanic arc in West Sulawesi during most of the Palaeogene and Neogene (Polve´ et al. 1997; Elburg et al. 2003). There is also little evidence for a collision that affected West Sulawesi (Hall & Wilson 2000; Calvert & Hall 2007), and it is now known that the North Banda basin formed by oceanic spreading during the Middle Miocene (Hinschberger et al. 2000). Spakman & Hall (2010) have proposed a tectonic model for the Banda and Sulawesi region that reconciles these and other observations with earlier interpretations, and offers an alternative to the previously accepted idea of slicing of continental slivers from New Guinea. There was an Early Miocene collision of the Sula Spur with the North Arm volcanic arc and East Arm ophiolite, and this continental area was then fragmented during extension caused by subduction rollback into the Banda embayment.

Celebes Molasse Pre-Miocene rocks of the different provinces are unconformably overlain by the Celebes Molasse – a weakly to moderately consolidated association of interbedded sedimentary formations that is widespread across Sulawesi (Sarasin & Sarasin 1901; van Bemmelen 1949). Sediments include conglomerate, quartz sandstone, greywacke and mudstone with subordinate intercalations of breccia, marl and coral limestone (e.g. van Bemmelen 1949; van Leeuwen et al. 2007). They have been interpreted to reflect deposition in a coastal alluvial plain environment situated along the flanks of rapidly uplifting and eroding mountains (Calvert 2000). The Celebes Molasse was originally interpreted to relate to a single Miocene collision (Ku¨ndig 1956). More recently it has been suggested to be diachronous across Sulawesi, representing several tectonic events (Hall & Wilson 2000). Within West Sulawesi and the East Arm it is interpreted to represent latest Miocene to Plio-Pleistocene uplift and erosion (Hall & Wilson 2000).

Stratigraphic observations We present a new stratigraphy (Fig. 2) for the western, central and eastern Togian Islands (Fig. 3). Based on new field observations and laboratory analyses, we define three new units, the Walea Formation, Peladan Formation and Benteng Intrusives, and integrate them with the previously recognized Lamusa Formation (Rusmana et al.

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Fig. 2. Schematic Neogene stratigraphy of the western, central and eastern Togian Islands, incorporating the age ranges derived in this study. Age (Ma) from Gradstein et al. (2004); PZ, Planktonic Foraminiferal biozones from BouDagher-Fadel (2008); LS, Far East Letter Stages from BouDagher-Fadel (2008). Note that the timescale is not linear.

1982, 1993), Bongka Formation (Rusmana et al. 1993) of the Celebes Molasse (Sarasin & Sarasin 1901; van Bemmelen 1949), Lonsio Formation

(Rusmana et al. 1982, 1993) and Luwuk Formation (Garrard et al. 1988). Our new stratigraphy ranges in age from possible Mesozoic basement rocks

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Fig. 3. Simplified geological map of the Togian Archipelago, modified from Rusmana et al. (1982, 1993) based on new field observations. Island names in bold italics; population centres in regular. Open circles and/or underlined dip measurements indicate locations examined in this study. Other structural information from Rusmana et al. (1982, 1993). Arrows and bold numbers (all prefixed RTG-) highlight the location of samples explicitly discussed in the text, for which GPS locations (decimal degrees) are listed in the inset table.

TOGIAN ISLANDS AND GORONTALO BAY

through to Quaternary deposits (Fig. 2). The most complete section is seen in the eastern islands (Walea Kodi and Walea Bahi; Fig. 3) where basement rocks, possibly of Eocene to Oligocene age, are overlain by Middle Miocene, Pliocene and Quaternary strata. The central and western islands expose more restricted sections dominated by Pliocene volcaniclastics and clastics respectively.

Lamusa formation Indurated sedimentary rocks of different types occur in several small exposures at the southern end of the channel between the islands of Batu Daka and Togian (Fig. 3). The rocks are weakly bedded and dip to the north. Lithologies include calcareous sandstones, interbedded with non-calcareous sandstones and dark mudstones, and dark, fine-grained recrystallized limestones. They are heavily brecciated and crushed. All lithologies are cut by small extensional faults. No fossils or sedimentary structures were identified. The formation has a minimum thickness of 3 m, but neither the top, nor the base was seen. Following Rusmana et al. (1993) we assign these rocks to the Mesozoic Lamusa Formation. Their highly indurated and veined character is consistent with the Mesozoic age suggested by Rusmana et al. (1993) and suggests that they may form part of the basement of the Togian Islands.

Walea formation Arc-related volcanic and volcaniclastic rocks are observed in exposures along the western coast of Walea Bahi and eastern coast of Walea Kodi. They include volcanic breccias, pillow lavas and arc-derived volcanogenic sediments. Well-bedded volcanogenic sedimentary rocks are exposed as a large, possibly fallen, block on the west coast of the southern peninsula of Walea Bahi. Mediumgrained, feldspar-rich, grey-brown beds are interbedded with green and blue-grey units with a finegrained green matrix on a scale of c. 5 cm. All show internal stratification and possible grading. Further north, just south of a large coastal embayment, a larger outcrop exposes an in-situ section of gently dipping (19–298 to the east) volcanogenic sediments (Fig. 4a) including interbedded sands and silts, some of which are calcareous. Mostly beds are laterally persistent with normal grading, parallel and cross-lamination and ripple cross bedding. Bedding parallel bioturbation and water escape structures are evident in the more sandy layers. Finer-grained siltstones dominate the upper part of the exposed sequence. The volcanogenic sedimentary rocks are interpreted to have been deposited as turbidites and debrites in a deepwater arc-related setting.

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Volcanic rocks occur along the west coast of the southern peninsula of Walea Bahi and in coastal outcrops along the channel between Walea Bahi and Walea Kodi. They include breccias, pillow lavas and more massive and layered lavas of basaltic to andesitic composition. The rocks have a finegrained groundmass of feldspar, pyroxene and altered olivine + phenocrysts of plagioclase feldspar + amygdales (up to 1 cm) of zeolite and/ or calcite. Large blocks (1  0.6 m) of breccia are exposed in the beach along the west coast of the southern peninsula of Walea Bahi. There are large, sub-rounded, clasts of dark grey (c. 10 cm) and green (c. 6 cm) material within a light grey matrix. The clasts have within them feldspar phenocrysts and amygdales of low-grade epidote-rich alteration products. Further north, pillows are exposed in several outcrops along the channel between Walea Bahi and Walea Kodi, often forming small headlands. Pillows are grey greenish in colour, weathering to grey brown. In most places they are heavily weathered and altered with late-stage alteration along fractures. Where relatively fresh, pillows show spectacular teardrop shapes (around 30 cm across), picked out by dark, glassy chilled rims of between 0.5 and 3 cm and fine-grained interpillow material (Fig. 4b), which provide right way-up criteria. Pillows contain abundant zeolite and/or calcite amygdales up to 1 cm in size; chilled rims contain small (5 mm) amygdales and alter to rusty coloured skins where weathered. More massive, layered lavas are also present; individual flows are marked by craggy tops and brecciated areas. Rusmana et al. (1982) reported similar pillow lavas, breccias, conglomerates and sandstones from Poh Head (Fig. 1), at the east end of the East Arm, and within the eastern Togian Islands, assigning them both to the Miocene Malik Formation. Simandjuntak (1986) assigned basaltic rocks from Poh Head to the basalt zone of the Balantak Ophiolite, and suggested a Late Cenomanian to Eocene age based on K –Ar ages. In a later revision, Rusmana et al. (1993) assigned these rocks to the Cretaceous Mafic Complex, whilst those in the eastern Togian Islands were reassigned to the MioPliocene Lonsio Formation (see below). Based on new observations we assign the basaltic lavas and volcanogenic sedimentary rocks of the Togian Islands to the Walea Formation, a new formation named from the type localities on the islands of Walea Bahi and Walea Kodi. Neither the top, nor the base, of the Walea Formation is observed but there is a minimum thickness of 5 m of pillows and 7 m of volcanogenic sediments. The total thickness of the formation is probably much greater. The exact age of the Walea Formation is unknown, but it is the stratigraphically oldest and structurally lowest unit seen in the eastern islands.

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a

b

30 cm

c

20 cm

d

20 cm

~2 m

Fig. 4. Field photographs of the Walea Formation and Lonsio Formation. (a) Arc-related (?) volcaniclastic sediments of the Walea Formation. (b) Basaltic pillows of the Walea Formation exposed on the west coast of Walea Bahi. (c) Well-bedded tuffs of the Lonsio Formation. Coarser tuff units (centre of image beneath pen) show rough stratification, dewatering and cross bedding. Finer tuff units (upper and lower sections of photograph) are more massive, have irregular bases and show an increase in joint density towards the upper boundary (lower section of photograph). (d) Syn-sedimentary folding and faulting within the Lonsio Formation.

Volcanic arc sedimentary rocks have not been reported from the East Arm ophiolite. Their association with basaltic lavas is more similar to the oldest rocks known from the North Arm which formed in an intra-oceanic arc between the Middle Eocene and earliest Miocene.

Peladan formation Hard, indurated limestones occur on (at least two) small islands situated around 250 m off the central west coast of Walea Bahi. Lithologies include micritic wackestones and packstones with

TOGIAN ISLANDS AND GORONTALO BAY

planktonic and benthic foraminifera and finegrained volcanogenic material (Table 1). These outcrops define the type section for the new Peladan Formation. Benthic and planktonic foraminifera indicate shallow inner platform or fore-reef shelf and deeper inner platform environments (Table 1). The sequence is well-bedded on a decimetre scale, up to a maximum of around 1 m (mode c. 30 cm) and dips gently towards the north. The sequence has a minimum observed stratigraphic thickness of around 12 m. The top and base of the sequence is not seen and the true thickness may be much greater. No other structure (folding/faulting) was observed. In places the beds have a rubbly texture interpreted to reflect re-working of components prior to deposition. Some thin (c. 10 cm), finer grained horizons appear not to have been reworked. In places the limestone are partially dolomitized. Early–Middle Miocene limestones of a similar age and character are reported from the North and East Arms of Sulawesi (e.g. Sukamto 1973; Norvick & Pile 1976; Rusmana et al. 1982; Garrard et al. 1988; van Leeuwen & Muhardjo 2005). In the Togian Islands Rusmana et al. (1982) previously assigned these rocks to the Salodik Formation and suggested a Late Paleocene to Early Miocene age. Later, Rusmana et al. (1993) reassigned them to the Lonsio Formation tuffaceous units. Micropalaeontological analyses of larger foraminifera and planktonic foraminifera were performed on five samples of the Peladan Formation (Table 1). Nannofossil dating was not attempted. We correlate the standard Planktonic Foraminiferal biozones (PZ) with the ‘Letter Stages’ (LS) of the Far East (as defined by BouDagher-Fadel 2008), relative to the geological timescale of Gradstein et al. (2004). Analyses indicate a late Middle Miocene age (PZ: Late N12 – Early N17; LS of the Far East: Late Tf2 – Early Tg). Based on their lithology and age we assign these rocks to the new Peladan Formation, named for one of the two islands on which they were observed.

Bongka formation (Celebes Molasse) Weakly to moderately consolidated interbedded sediments with characteristic lithic-rich horizons occur in heavily weathered outcrops along the channel between the islands of Batu Daka and Togian. They are sub-rounded, green-brown, mediumgrained sandstones with bands of coarser, angular lithic fragments, medium-grained sandstones with a slabby, bedded character, and brecciated material with possible ultrabasic content. Petrographical analyses reveal a matrix of serpentinite-rich material. The sequence dips moderately (c. 308) to the north. Sediments with a coarser grain size, but of comparable composition and structure (moderate

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c. 308 dips to the NE/NNE) are seen in a more extensive cliff outcrop on a small island east of the village of Katupat. Lithologies at this location include laminated siltstones and sandstones, and pebble conglomerates containing well-rounded pebbles (up to 2 cm) dominated by ophiolitic material (basalts, dolerites, gabbros and serpentinite) with some chert and limestones. The silts and sands contain abundant, highly oxidized, plant material. The formation has a minimum thickness, as observed in outcrop, of 15 m but neither the top nor the base of the unit is seen. Similar deposits, but coarser still in grain size, were observed in roadside outcrops on the northern coast of the East Arm of Sulawesi, west of the town of Bunta (Rusmana et al. 1982, 1993; this study). Here they comprise coarse, massive, sandstones with pebble-rich horizons that include large clasts (up to 3 cm) of red chert and cobbles (up to 15 cm) of basalts, dolerites, gabbros, metagabbros and serpentinite with some limestones. Again, the sequence dips north at moderate angles of c. 308. We observed a minimum stratigraphic thickness of around 20 m, although neither the top nor the base of the unit was seen. Based on strong lithological and compositional similarities between these rocks and those within the Togian Islands, we follow Rusmana et al. (1993) in assigning all of these rocks to the Bongka Formation of the Celebes Molasse. In northern Sulawesi palaeontological dating of the Celebes Molasse suggests a Late Early Pliocene to Mid Pleistocene age (Norvick & Pile 1976; Ratman 1976; Hadiwijoyo et al. 1993; Chamberlain & Seago 1995). Late Miocene – Pliocene ages have been reported for the East Arm (Surono & Sukarna 1996). The Celebes Molasse has been interpreted as alluvial fan and coastal fan delta deposits that reflect the deposition of locally sourced sediment in alluvial plain environments with a marginal marine influence (Calvert 2000). In contrast, the ophiolitic material observed in the Togian Islands has no local source, and such material can only have been derived from the East Arm Ophiolite. Based on the relative grain size and shared structural characteristics (gentle north dip), we suggest that outcrops of the Bongka Formation within the East Arm and the Togian Islands represent proximal (coarser) and distal (finer) alluvial fan deposits respectively, both having been transported north from the interior of the East Arm.

Lonsio formation Volcaniclastic rocks are extensively exposed in coastal outcrops on the northern peninsula of Tala Teoh, the north coast of Togian and the west coast of Walea Kodi. They are grain-supported rocks

Depositional environment

Microfacies

RTG 18 A

Shallow inner platform/ fore-reef shelf

Micritic packstone of planktonic and benthic foraminifera. Micritic patches reworked into the matrix.

RTG 18 B

Shallow inner platform/ fore-reef shelf

Micritic packstone of larger benthic foraminifera

RTG 18 C

Shallow inner platform/ fore-reef shelf

Micritic packstone of recrystallized algae and benthic foraminifera. Micritic patches reworked into the matrix.

RTG 18 D

Shallow inner platform/ fore-reef shelf

Micritic packstone of foraminifera and algae.

RTG 18E

Relatively deeper inner platform

Micritic wackestone of foraminifera. Reworked patches of micrite are also present.

Components Benthic foraminifera: Cycloclypeus indopacific, Katacycloclypeus martini, Amphistegina spp., Cycloclypeus pillaris, Cycloclypeus spp., Sphaerogypsina spp., Lepidocyclina spp., Lepidocyclina (Nephrolepidina) spp., L. (Nephrolepidina) angulosa Planktonic foraminifera: Sphaeroidinellopsis spp., Globorotalia praemenardii, Globigerinoides spp., Globorotalia peripheroacuta, Globorotalia praefohsi, Globoquadrina altispira, Planorbulinella solida Globoquadrina spp., Globoquadrina dehiscens, Echinoid spp., fragments of rodophyte algae. Benthic foraminifera: Cycloclypeus spp., Cycloclypeus pillaria, Cycloclypeus carpenteri, Amphistegina spp., Discogypsina discus. Textularia spp., Carpenteria spp., Katacycloclypeus annulatus, Planorbulinella spp. Planktonic foraminifera: Dentoglobigerina altispira, Globigerinoides primordius, Globigerina spp., Globigerinoides quadrilobatus, Orbulina suturalis, Globorotalia praemenardii, Echinoid spp., fragments of rodophyte algae and corals, Gastropods, fragments of bryozoa. Benthic foraminifera: Cycloclypeus spp., Amphistegina spp., Textularia spp., Miliolid spp., Sphaerogypsina spp. Planktonic foraminifera: Globigerinoides quadrilobatus, Orbulina spp., Globigerinoides spp., Globorotalia menardii, fragments of rodophyte algae, Lithophyllum spp., Lithothamnium spp., Gastropods, Echinoid spp., rare fragments of bryozoa. Benthic foraminifera: Cycloclypeus pillaria, Planorbulinella solida, Gypsina spp., Sphaerogypsina spp., Elphidium spp., Nodosaria spp. Planktonic foraminifera: Globoquadrina spp., Globigerinoides trilobus, Globigerinoides spp., Orbulina suturalis, Globorotalia conoidea, Globorotalia menardii, Globorotalia scitula, Gastropod spp., fragments of bryozoa, fragments of coral. Benthic Foraminifera: Lepidocyclina spp., Carpenteria spp., Cycloclypeus spp., Cycloclypeus pillaria, Operculina spp., Heterostegina spp., Gypsina spp., Planorbulinella larvata, Lagena spp., Textularia spp. Planktonic foraminifera: Globoquadrina altispira, Globorotalia spp., Globorotalia scitula, Globoquadrina dehiscens, Globorotalia menardii, Globoquadrina dehiscens, Globorotalia fohsi, Ostracod spp., Gastropod spp.

Age (PZ/LS)* (based on first appearance) Late N12/Late Tf2

Late N12/Late Tf2

N12 and younger/Tf2 and younger

Late N12 – Early N17/Tf3 – Early Tg

Late N12 – Early N13/Late Tf2 – Early Tf3

*We correlate the standard Planktonic Foraminiferal biozones (PZ) with the ‘Letter Stages’ (LS) of the Far East (as defined by BouDagher-Fadel 2008), relative to the biostratigraphical timescale (as defined by Gradstein et al. 2004)

M. A. COTTAM ET AL.

Sample ID

186

Table 1. Biostratigraphical age, facies and palaeoenvironmental analyses of the Peladan formation

TOGIAN ISLANDS AND GORONTALO BAY

with a carbonate (dominantly sparry) matrix. Microfossils and algal fragments are also embedded within the matrix; their abundance varies between units (Table 2). The volcaniclastics are well-bedded, commonly at the decimetre scale, with a maximum bed thickness of around 3 m. Two main bed types alternate at a range of scales. Individual beds appear laterally persistent at the outcrop scale. Stratified beds are typically around 10 to 30 cm in thickness and show parallel lamination of fine to coarse sand. They contain rare horizons of small (up to fine pebbles) angular lithic fragments (Fig. 4c). In places the beds show spectacular dewatering structures, and may be wholly or partly cross-bedded, producing an irregular upper surface. The base of the beds is almost universally planar. Stratified units are overlain by fine-grained cream coloured material, which show little variation in grain size or internal structure (Fig. 4c). Beds range in thickness from cm scale to a maximum of 3 m. Their bases are commonly irregular, reflecting the topography of the stratified layer below, and they display an increase in joint density towards the upper boundary, which is characteristically planar. In places the finer beds may be very thin, or entirely absent from the sequence. Overall, the sequence dips gently in various directions. Locally, the rocks dip steeply and show intense syn-sedimentary folding and faulting (Fig. 4d), interpreted to reflect soft sediment deformation. The sequence has a minimum stratigraphic thickness of around 20 m, however, the top and base of the sequence is not seen and the true thickness is probably much more. Comparable volcaniclastic rocks are observed on Poh Head, where they include thick sequences of coarse stratified units (this study). Rusmana et al. (1982, 1993) described these rocks as tuffaceous sediments and assigned them to the Lonsio Formation. Simandjuntak (1986) interpreted similar volcanogenic sediments from the East Arm as megacyclic turbidites, and assigned them to the Lonsuit Turbidites of the Batui Group. We interpret these rocks as tuffaceous sediments that reflect rapid aqueous reworking of primary volcaniclastic material during deposition in a shallow marine environment soon after eruption. Microfossil observations suggest depths less than 200 m. Stratification reflects crude sorting of coarse ash during settling through the water column; cross-bedding may reflect turbidity currents formed by ash initially held in suspension. Finer-grained ash settled more slowly through the water column, draping topography in the underlying coarse units. Pumice is largely absent and may have been floated off and not preserved (e.g. Freundt 2003). The repeated sequence of coarse and fine tuff may reflect pulses within a

187

single eruption or input from several eruptions. Dewatering structures within coarse units suggest rapid loading by the subsequent fine units. Jointing present near the upper boundary of the finer units may be syn- or post-depositional. Based on their striking similarity to tuffaceous units observed on Poh Head (Simandjuntak 1986; A. J. Barber, pers. comm. 2009) we assign these rocks to the Lonsio Formation of Rusmana et al. (1982, 1993). Micropalaeontological analyses of larger foraminifera and planktonic foraminifera were performed on five tuff samples from the Lonsio Formation (Table 2). Nannofossil dating was not attempted. Foraminiferal assemblages range from N4 and younger (PZ) and Te and younger (LS), and constrain a Late Miocene to Early Pliocene age (PZ: N19; LS: Early Th).

Benteng Intrusives Intrusive rocks of intermediate composition are exposed in isolated outcrops, along the northern and southern coasts of Togian Island. They occur as small intrusions, often forming topographic highs and small islands. We infer the presence of additional intrusive bodies within the interior of Togian Island based on the presence of isolated steep-sided topographic highs visible from the coast as shown on the map of Ku¨ndig (1956). The rocks have a fine to medium grained light-grey groundmass with phenocrysts of phlogopite mica (up to 7 mm) + feldspar (6–7 mm) + hornblende (1– 3 mm) + mafic xenoliths (up to 2 cm). In places feldspar phenocrysts are concentrated into ‘trails’ up to 20 cm long. Orthogonal sub-horizontal and sub-vertical joints spaced at around 20 to 50 cm, and resulting in a characteristic blocky appearance, suggest intrusion at shallow depths. In places the rocks are cut by east –west trending brittle faults, producing breccia zones around 1 m wide. These rocks are classified (Fig. 5; Table 3) as trachydacites and trachyandesites on the total alkalis v. silica (TAS) diagram of Le Maitre (1989) (they are syenites on TAS diagrams adapted for plutonic rocks (e.g. Wilson 1989)) and belong to the alkaline magma series (Kuno 1966; Irvine & Baragar 1971). They have an extremely K-rich chemistry and plot within the shoshonitic field of Rickwood (1989) on a K2O v. SiO2 diagram. Intrusive intermediate rocks were first recognized on Togian Island by Ku¨ndig (1956), who identified rocks of andesitic composition. These were subsequently misidentified as basaltic (Rusmana et al. 1982) or volcaniclastic (Rusmana et al. 1993) in character. We assign these intrusive rocks to the new Benteng Intrusives, named for the village of the same name in south central Togian Island (Fig. 3).

188

Table 2. Biostratigraphical age, facies and palaeoenvironmental analyses of the Lonsio formation Sample ID

Microfacies

RTG 25

Inner neritic, planktonic & shelf benthic foraminifera drifted/ reworked into volcanic deposits.

Sparitic packstone of volcanic sediments rich in embedded planktonic foraminifera and rare larger benthic and algae fragments

RTG 26

Inner neritic

Sparitic packstone of volcanic sediments with rare embedded planktonic foraminifera Sparitic packstone of volcanic sediments with rare embedded planktonic foraminifera Sparitic packstone of volcanic sediments rich in embedded planktonic foraminifera and rare larger benthic and algae fragments

RTG 27 RTG 30

Inner neritic, planktonic & shelf benthic foraminifera drifted/ reworked into volcanic deposits.

RTG 36

Inner neritic

Sparitic packstone of volcanic sediments with rare embedded planktonic foraminifera

Components

Age (PZ/LS)* (based on first appearance)

Globoquadrina altispira, Globoquadrina spp., Orbulina spp., Globorotalia margaritae, Globorotalia scitula, Sphaeroidinellopsis subdehiscens, Globigerinoides trilobus, Globigerinoides quadrilobatus, Globorotalia acostaensis, Fragments of rodophyte algae, Elphidium spp. Globigerinoides spp.

N19/Early Th

Globigerinoides spp.

N4 and younger/Upper Te and younger

Catapsydrax spp., Orbulina universa, Globoquadrina dehiscens, Pulleniatina primalis, Globoquadrina altispira, Globorotalia globosa, Globorotalia humerosa, Globorotalia mayeri, Globorotalia scitula, Globigerinoides sacculifer, Globigerinoides quadrilobatus, Elphidium spp., Amphistegina spp., Heterostegina spp., Asterigerina spp. Orbulina universa, Globigerinoides spp., Globigerinoides quadrilobatus, Globoquadrina spp.

N19/Early Th

N4 and younger/Upper Te and younger

N4-N19/Upper Te – Early Th

*We correlate the standard Planktonic Foraminiferal biozones (PZ) with the ‘Letter Stages’ (LS) of the Far East (as defined by BouDagher-Fadel 2008), relative to the biostratigraphical timescale (as defined by Gradstein et al. 2004)

M. A. COTTAM ET AL.

Depositional environment

TOGIAN ISLANDS AND GORONTALO BAY

Fig. 5. Major element classification diagrams for the volcanic rock samples analysed in this study. (a) Total alkalis (K2O þ Na2O) v. silica (SiO2) diagram. Field boundaries are those of Le Maitre (1989): 1, andesite; 2, dacite; 3, trachyandesite; 4, trachydacite. Subdivision into alkaline and sub-alkaline series: dashed curved line – Irvine & Baragar (1971); solid curved line – Kuno (1966). (b) K2O v. SiO2 diagram. Series boundaries and nomenclature: dashed lines and bold italics, Le Maitre (1989); solid lines and nomenclature in parentheses, after Rickwood (1989).

Table 3. Major element data (weight %) for samples of the Benteng Intrusives analysed in this study Sample ID

RTG08

RTG09

RTG12

RTG31

SiO2 Al2O3 Fe2O3 MgO CaO Na2O K2O TiO2 P2O5 MnO

63.36 15.52 3.92 2.25 3.67 4.41 5.40 0.34 0.42 0.09

63.14 15.43 3.77 2.37 3.77 4.62 5.34 0.33 0.41 0.09

58.97 15.85 4.09 2.97 4.38 3.88 6.51 0.76 0.45 0.06

61.39 14.54 5.82 2.24 2.21 3.44 6.46 0.69 0.43 0.12

Total

99.38

99.26

97.92

97.34

189

Five high-purity mica separates from four samples of the Benteng Intrusives were dated using 40 Ar/39Ar techniques. Samples were crushed, graded using disposable nylon cloth sieves in a brass collar and separated using conventional electromagnetic techniques. High-purity mineral separates were handpicked from the 63– 250 mm fraction, and for RTG-12 from the .250 mm fraction, thus any contamination in the analyses is assumed to be due to intra-grain alteration and/or contaminants. All analyses were undertaken in the Argon Laboratory of the Research School of Earth Sciences, The Australian National University, using the furnace step-heating technique (Table 4). Samples were irradiated at the McMaster Nuclear Reactor, McMaster University, Canada using Sanidine 92– 176 from Fish Canyon Tuff, Colorado (K/Ar reference age 28.10 + 0.04 Ma) as the Fluence Monitor (Spell & McDougall 2003). Ages were calculated using the 40K abundances and decay constants of Steiger & Ja¨ger (1977). Uncertainties in isotopic ratios and ages are quoted at the 1s level. For all samples plots of 36Ar/40Ar v. 39Ar/40Ar demonstrate the presence of one main gas population, with varying amounts of contaminants (such as excess argon), and a large atmospheric argon component – particularly in the coarser grained samples (Fig. 6). The oldest ages are preserved in the high-temperature heating steps of coarsegrained (.250 mm) biotite from samples RTG12 (2.40 + 0.01 Ma; MSWD (mean square of weighted deviation) 1.58) and RTG31 (2.02 + 0.01 Ma; MSWD 0.01) (Fig. 6). However, significant atmospheric argon contents, and evidence of argon loss and possible younger events render the meaning of these ages ambiguous. Analysis of fine-grained (250 –63 mm) biotite from sample RTG12 contains significantly less atmospheric argon than the coarser-grained biotite and produced a reliable, consistently flat spectrum of 1.80 + 0.01 Ma (MSWD 3.95) (Fig. 6). This analysis provides the best age for this sample and the most robust age for the Benteng Intrusives. Analyses of fine-grained mica from two other samples gave robust Pleistocene ages. Despite disturbance during the initial heating steps (linked to variation in Ca), over 50% of the gas emitted from RTG08 produced a strong plateau with an age of 1.52 + 0.02 Ma (MSWD 0.3) (Fig. 6). Except for several contaminated intervening steps, analysis of fine-grained biotite from RTG09 would have produced a similar plateau, giving an age of 1.68 + 0.09 Ma (MSWD 3.9) with a younger age of 1.37 + 0.02 Ma evident (Fig. 6). Based on our new field observations and laboratory analyses we interpret these rocks as shallow level stocks and dykes of Late Pliocene to Early Pleistocene age. The observed and inferred intrusive

190

Table 4. Temp (8C)

40

Ar/39Ar step heating analyses

Ar36 (mol)

err (%)

Ar37 (mol)

Lambda K40 ¼ 5.5430E-10

err (%)

Ar39 (mol)

err (%)

Ar40 (mol)

err (%)

Ar40* (%)

Ar40*/ Ar39(K)

Cumulative Ar39(%)

Calculated age Ma + 1s.d.

Ca/K

3.90 6.16 3.02 8.62 2.87 3.48 2.45 3.38 0.88 3.74 1.57 0.64 1.06 5.07

1.79E-16 1.40E-16 2.12E-16 3.92E-16 7.11E-16 8.54E-16 1.66E-15 2.09E-15 2.93E-15 2.42E-15 1.67E-15 9.95E-16 4.91E-17 1.27E-17 1.43E-14

3.38 3.73 1.75 0.29 1.25 0.21 0.91 0.54 0.63 0.47 0.64 0.72 0.80 20.02

6.52E-15 6.08E-15 1.28E-14 2.52E-14 4.07E-14 4.97E-14 9.11E-14 1.27E-13 1.99E-13 1.66E-13 1.21E-13 7.28E-14 2.00E-15 4.88E-16 9.20E-13

0.24 0.56 0.20 0.11 0.09 0.09 0.11 0.11 0.12 0.09 0.13 0.08 0.50 0.61

1.30E-13 4.77E-14 8.16E-14 1.57E-13 3.24E-13 4.00E-13 9.10E-13 9.24E-13 8.37E-13 6.55E-13 3.11E-13 1.19E-13 1.25E-14 1.04E-14 4.92E-12

0.25 0.64 0.21 0.16 0.11 0.11 0.13 0.18 0.17 0.11 0.20 0.14 0.53 0.70

4.60 8.90 10.00 5.50 6.60 5.30 5.80 6.60 11.60 12.10 15.50 28.10 45.10 4.20

0.92 0.70 0.64 0.34 0.53 0.43 0.58 0.48 0.49 0.48 0.40 0.46 2.94 0.92 0.49

0.71 1.37 2.76 5.49 9.92 15.32 25.22 39.02 60.61 78.64 91.83 99.74 99.95 100.00

2.87 + 0.78 2.20 + 0.36 2.00 + 0.14 1.08 + 0.16 1.66 + 0.09 1.35 + 0.13 1.83 + 0.12 1.51 + 0.14 1.53 + 0.06 1.50 + 0.08 1.25 + 0.06 1.44 + 0.06 9.20 + 0.90 2.89 + 4.44 1.53 + 0.10

0.59 0.59 0.29 0.16 0.14 0.18 0.15 0.14 0.07 0.06 0.11 1.09 103.00 74.20

0.99 0.60 0.66 1.27 0.69 0.26 0.59 0.91 0.33 0.40 0.35 1.33 1.33

7.96E-15 1.06E-14 2.20E-14 4.12E-14 1.52E-13 1.28E-13 2.38E-13 3.39E-13 3.07E-13 7.00E-13 4.27E-13 4.59E-14 4.55E-14 2.46E-12

0.28 0.21 0.13 0.38 0.09 0.12 0.22 0.10 0.11 0.06 0.09 0.21 0.09

2.78E-13 1.78E-13 2.08E-13 3.45E-13 1.48E-12 1.26E-12 1.23E-12 7.54E-13 5.74E-13 1.16E-12 8.00E-13 2.19E-13 1.53E-13 8.64E-12

0.31 0.26 0.16 0.43 0.11 0.13 0.34 0.19 0.13 0.11 0.12 0.32 0.14

2.10 2.90 4.30 6.70 5.40 5.10 11.10 23.60 26.70 26.40 25.60 19.80 18.40

0.72 0.49 0.41 0.56 0.52 0.50 0.58 0.53 0.50 0.44 0.48 0.95 0.62 0.50

0.32 0.75 1.65 3.32 9.50 14.70 24.37 38.13 50.58 78.98 96.31 98.16 100.00

2.26 + 0.94 1.53 + 0.35 1.27 + 0.15 1.77 + 0.23 1.64 + 0.13 1.56 + 0.18 1.80 + 0.11 1.65 + 0.03 1.57 + 0.03 1.37 + 0.02 1.50 + 0.04 2.99 + 0.20 1.95 + 0.08 1.57 + 0.06

0.54 0.36 0.34 0.21 0.17 0.15 0.11 0.06 0.07 0.11 1.17 20.20 4.47

J ¼ 1.7413E-3 +0.413

Sample RTG-09 (R2) Biotite 600 9.22E-16 0.55 2.25E-15 650 5.83E-16 0.63 1.99E-15 700 6.71E-16 0.45 3.90E-15 750 1.09E-15 0.72 4.56E-15 800 4.73E-15 0.36 1.40E-14 840 4.03E-15 0.37 1.01E-14 930 3.69E-15 0.56 1.43E-14 970 1.92E-15 0.39 1.14E-14 1020 1.40E-15 0.60 1.07E-14 1070 2.84E-15 0.40 4.04E-14 1140 2.07E-15 0.80 2.64E-13 1200 7.59E-16 1.23 4.84E-13 1350 4.57E-16 0.84 1.07E-13 Total 2.52E-14 9.67E-13 Lambda K40 ¼ 5.5430E-10

Ar38 (mol)

5.24 2.74 2.72 2.40 4.79 1.43 1.65 2.18 2.75 1.27 0.22 0.36 0.60

3.05E-16 2.54E-16 4.28E-16 7.13E-16 2.73E-15 2.31E-15 3.58E-15 4.60E-15 4.10E-15 9.22E-15 5.79E-15 9.02E-16 6.62E-16 3.56E-14

J ¼ 1.7378E-3 +0.413

M. A. COTTAM ET AL.

Sample RTG-08 (R1) Biotite 600 4.19E-16 0.87 2.04E-15 650 1.47E-16 1.46 1.89E-15 700 2.48E-16 0.73 1.96E-15 750 4.99E-16 0.48 2.07E-15 800 1.02E-15 0.24 2.88E-15 840 1.28E-15 0.38 4.70E-15 890 2.90E-15 0.26 7.08E-15 930 2.91E-15 0.41 9.44E-15 970 2.49E-15 0.34 7.38E-15 1020 1.94E-15 0.48 5.66E-15 1070 8.81E-16 0.67 6.81E-15 1140 2.98E-16 0.73 4.17E-14 1200 5.94E-17 2.41 1.04E-13 1350 4.00E-17 3.20 1.85E-14 Total 1.51E-14 2.16E-13

err (%)

Lambda K40 ¼ 5.5430E-10

1.48 0.35 0.88 1.91 0.83 0.28 0.45 0.98 0.51 0.53 0.76 0.26 1.77

2.39E-15 4.43E-15 1.20E-14 2.49E-14 5.73E-14 9.55E-14 2.24E-13 5.49E-13 4.40E-13 3.52E-13 3.82E-13 7.21E-13 5.27E-14 2.92E-12

0.25 0.32 0.08 0.30 0.11 0.17 0.35 0.17 0.08 0.19 0.08 0.06 0.53

6.94E-13 5.60E-13 7.64E-13 1.41E-12 3.43E-12 2.50E-12 3.61E-12 4.79E-12 2.22E-12 1.28E-12 1.14E-12 1.38E-12 9.48E-14 2.39E-11

0.26 0.33 0.10 0.33 0.12 0.21 0.45 0.24 0.14 0.23 0.16 0.08 0.64

-0.20 -0.10 0.60 1.20 1.20 1.10 2.10 6.70 12.80 18.30 25.90 39.70 38.50

0.00 0.00 0.40 0.68 0.70 0.30 0.34 0.58 0.65 0.67 0.77 0.76 0.69 0.65

0.08 0.23 0.65 1.50 3.46 6.74 14.40 33.22 48.30 60.37 73.47 98.19 100.00

0.00 + 4.90 0.00 + 2.69 1.27 + 1.05 2.15 + 1.08 2.20 + 0.58 0.94 + 0.36 1.08 + 0.38 1.83 + 0.13 2.04 + 0.05 2.09 + 0.12 2.42 + 0.07 2.40 + 0.02 2.17 + 0.09 2.03 + 0.13

0.35 0.19 0.34 0.42 0.21 0.11 0.02 0.01 0.01 0.01 0.01 0.05 0.09

2.11 2.01 0.98 2.31 0.45 0.54 0.78 0.29 0.44 0.34 0.25 0.18 0.35 4.55

8.05E-15 1.62E-14 2.74E-14 1.15E-13 7.21E-14 1.21E-13 1.85E-13 2.72E-13 3.50E-13 3.51E-13 3.15E-13 5.33E-13 1.29E-13 5.50E-15 2.50E-12

0.25 0.18 0.20 0.17 0.39 0.10 0.16 0.12 0.34 0.08 0.07 0.13 0.09 0.41

1.18E-13 1.23E-13 1.14E-13 2.92E-13 1.26E-13 1.72E-13 2.03E-13 2.37E-13 3.17E-13 3.71E-13 3.58E-13 5.91E-13 1.61E-13 1.88E-14 3.20E-12

0.29 0.22 0.25 0.19 0.45 0.14 0.23 0.17 0.38 0.10 0.10 0.15 0.14 0.44

7.40 7.90 15.60 24.20 32.60 41.70 52.90 66.30 63.50 54.20 48.90 52.40 48.90 17.00

1.09 0.60 0.65 0.62 0.57 0.59 0.58 0.58 0.58 0.57 0.56 0.58 0.61 0.58 0.58

0.32 0.97 2.07 6.65 9.53 14.37 21.75 32.62 46.61 60.66 73.27 94.61 99.78 100.00

3.39 + 0.51 1.87 + 0.19 2.02 + 0.13 1.92 + 0.10 1.78 + 0.04 1.85 + 0.04 1.81 + 0.02 1.80 + 0.01 1.80 + 0.01 1.79 + 0.01 1.74 + 0.01 1.81 + 0.01 1.90 + 0.02 1.81 + 0.24 1.81 + 0.02

0.54 0.53 0.47 0.36 0.20 0.12 0.05 0.02 0.02 0.02 0.04 0.16 1.23 2.30

J ¼ 1.7442E-3 +0.426

Sample RTG-12 (fine-grained) (R4) Biotite 600 3.70E-16 1.23 2.29E-15 12.90 650 3.83E-16 0.85 4.55E-15 1.32 700 3.24E-16 1.14 6.82E-15 3.16 750 7.46E-16 1.34 2.17E-14 4.21 800 2.83E-16 0.87 7.56E-15 1.05 850 3.30E-16 1.68 7.82E-15 3.81 890 3.08E-16 1.10 5.27E-15 4.20 930 2.46E-16 0.90 3.26E-15 4.00 970 3.62E-16 1.10 2.77E-15 13.41 1020 5.43E-16 0.74 3.14E-15 0.94 1070 5.92E-16 0.72 6.44E-15 2.47 1140 9.20E-16 0.59 4.58E-14 0.58 1200 2.96E-16 0.73 8.36E-14 0.72 1350 5.45E-17 2.59 6.65E-15 1.25 Total 5.76E-15 2.08E-13 Lambda K40 ¼ 5.5430E-10

4.87E-16 4.18E-16 6.49E-16 1.24E-15 2.93E-15 2.80E-15 5.12E-15 9.69E-15 6.72E-15 5.04E-15 5.28E-15 9.57E-15 6.83E-16 5.06E-14

1.88E-16 2.74E-16 4.00E-16 1.59E-15 9.40E-16 1.60E-15 2.46E-15 3.56E-15 4.57E-15 4.57E-15 4.14E-15 6.96E-15 1.71E-15 8.01E-17 3.30E-14

TOGIAN ISLANDS AND GORONTALO BAY

Sample RTG-12 (coarse-grained) (R3) Biotite 600 2.35E-15 0.47 4.36E-16 16.96 650 1.90E-15 0.63 4.40E-16 22.82 700 2.57E-15 0.45 2.14E-15 5.57 750 4.70E-15 0.45 5.49E-15 1.01 800 1.15E-14 0.28 6.25E-15 2.98 840 8.35E-15 0.38 5.59E-15 5.75 900 1.19E-14 0.50 2.25E-15 5.51 980 1.51E-14 0.43 3.31E-15 10.36 1020 6.52E-15 0.29 2.40E-15 11.44 1060 3.49E-15 1.03 2.18E-15 11.95 1100 2.81E-15 0.98 2.34E-15 4.76 1200 2.76E-15 0.43 1.71E-14 0.79 1350 1.94E-16 2.36 2.49E-15 5.92 Total 7.41E-14 5.25E-14

J ¼ 1.7305E-3 +0.356 (Continued)

191

192

Table 4. Continued Temp (8C)

Ar36 (mol)

err (%)

Ar37 (mol)

err (%)

Lambda K40 ¼ 5.5430E-10

err (%)

Ar39 (mol)

err (%)

Ar40 (mol)

err (%)

Ar40* (%)

Ar40*/ Ar39(K)

Cumulative Ar39(%)

Calculated age Ma + 1s.d.

2.80E-17 1.50E-16 3.39E-16 9.63E-16 1.99E-15 3.23E-15 3.87E-15 2.68E-15 3.48E-15 2.19E-15 3.00E-15 2.32E-15 3.17E-15 1.74E-15 6.73E-16 4.96E-17 2.99E-14

4.16 4.28 0.32 0.59 0.56 0.47 0.58 0.79 1.11 0.73 1.10 0.38 1.20 2.32 1.57 1.36

1.17E-15 7.16E-15 1.65E-14 5.15E-14 1.12E-13 1.91E-13 2.26E-13 1.60E-13 1.99E-13 1.35E-13 1.73E-13 1.32E-13 1.69E-13 1.01E-13 3.66E-14 1.26E-15 1.71E-12

0.50 0.20 0.21 0.18 0.11 0.05 0.07 0.66 0.15 0.10 0.26 0.17 0.21 0.14 0.15 1.19

1.82E-14 9.00E-14 2.08E-13 5.10E-13 8.98E-13 1.40E-12 1.66E-12 1.23E-12 1.56E-12 8.55E-13 1.28E-12 1.14E-12 1.59E-12 8.18E-13 1.79E-13 1.75E-14 1.35E-11

0.51 0.24 0.23 0.21 0.14 0.10 0.12 0.79 0.16 0.16 0.29 0.21 0.31 0.22 0.19 1.21

-0.30 0.70 -0.40 2.30 2.50 5.30 6.40 6.90 7.40 8.10 8.80 7.50 6.10 8.90 16.30 22.50

0.00 0.08 0.00 0.22 0.20 0.39 0.47 0.53 0.59 0.52 0.65 0.65 0.58 0.72 0.80 3.14 0.52

0.07 0.49 1.45 4.46 10.98 22.14 35.36 44.72 56.32 64.20 74.31 82.01 91.90 97.79 99.93 100.00

0.00 + 1.83 0.26 + 0.50 0.00 + 0.36 0.70 + 0.22 0.63 + 0.10 1.20 + 0.09 1.47 + 0.08 1.65 + 0.28 1.83 + 0.17 1.61 + 0.17 2.03 + 0.16 2.01 + 0.19 1.80 + 0.13 2.25 + 0.18 2.49 + 0.21 9.76 + 1.91 1.62 + 0.16

J ¼ 1.7286E-3 +0.426

Ca/K

0.17 0.08 0.04 0.04 0.04 0.04 0.04 0.04 0.04 0.04 0.04 0.05 0.05 0.11 3.13 1.93

M. A. COTTAM ET AL.

Sample RTG-31 (coarse-grained) (R6) Biotite 470 6.17E-17 4.37 1.04E-16 24.22 510 3.02E-16 1.19 3.13E-16 15.37 550 7.04E-16 0.89 3.85E-16 2.84 600 1.68E-15 0.67 1.12E-15 3.27 650 2.96E-15 0.36 2.58E-15 0.45 700 4.48E-15 0.35 4.28E-15 2.73 750 5.25E-15 0.28 4.91E-15 5.91 790 3.87E-15 0.81 3.03E-15 2.15 840 4.88E-15 0.73 3.76E-15 5.11 890 2.65E-15 0.91 2.99E-15 4.52 950 3.94E-15 0.57 3.62E-15 7.25 1000 3.55E-15 0.73 3.19E-15 3.05 1050 5.05E-15 0.54 4.65E-15 5.84 1100 2.52E-15 0.77 6.00E-15 2.63 1200 5.26E-16 1.56 6.02E-14 0.42 1350 4.62E-17 13.90 1.28E-15 8.90 Total 4.25E-14 1.02E-13

Ar38 (mol)

TOGIAN ISLANDS AND GORONTALO BAY

193

5.0

5.0

RTG-09

RTG-08 4.0

3.0

3.0

Age (Ma)

4.0

2.0

2.0

1.0

1.0

0.0

0.0 0

0

20 40 60 80 100 Cumulative % 39Ar released

20 40 60 80 100 Cumulative % 39Ar released

5.0

5.0

5.0 RTG-31

RTG-12 C

RTG-12 F 4.0

4.0

3.0

3.0

3.0

2.0

2.0

2.0

1.0

1.0

1.0

Age (Ma)

4.0

0.0

0.0 0

20 40 60 80 100 Cumulative % 39Ar released

0.0 0

20 40 60 80 100 Cumulative % 39Ar released

0

20 40 60 80 100 Cumulative % 39Ar released

Fig. 6. 40Ar/39Ar age spectra plots for biotite step-heating analyses performed on four samples from the Benteng Intrusives. For sample RTG12 separate analyses were undertaken on coarse (.250 mm; RTG-12 C) and fine (63–250 mm; RTG-12 F) mica.

bodies follow a broadly north–south trend through the centre of Togian Island, supporting the spatial observations of Ku¨ndig (1956), and indicating a possible structural control on their intrusion.

Quaternary age and assign them to the Luwuk Formation (Garrard et al. 1988).

Discussion Luwuk formation Reefal limestones are found throughout the archipelago, and dominate outcrop in the western islands (e.g. Batu Daka). They occur as high cliffs and raised terraces of poorly bedded, rubbly limestones containing broken coral fragments. The limestones have been uplifted to heights of around 200 m within the archipelago and to more than 300 m on the East Arm (Garrard et al. 1988). Following Rusmana et al. (1982) we allocate these rocks a

The Togian Islands offer a unique opportunity to investigate Gorontalo Bay. Our new stratigraphy offers insight into several aspects of the basin including the nature of its basement rocks, its age and its mode of formation.

Basement rocks beneath Gorontalo Bay Based on geophysical evidence, Silver et al. (1983b) suggested that much of Gorontalo Bay is underlain

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by basement rocks belonging to the East Sulawesi Ophiolite (East Sulawesi Ophiolite). Beneath these may be continental basement rocks belonging to the leading edge of the Banggai-Sula microcontinental block (Silver et al. 1983a; Hall & Wilson 2000). Other hypotheses are that the bay is underlain by oceanic crust of the Northern Sulawesi Province (Monnier et al. 1995) or that the basement of Gorontalo Bay comprises a complex amalgamation of at least two tectonostratigraphic provinces. The position of the Togian archipelago in the middle of Gorontalo Bay provides an opportunity to test these hypotheses. Our new field observations suggest that the central part of Gorontalo Bay, including the Togian Islands, is underlain by oceanic and arc basement of the Northern Sulawesi Province rather than the continental basement of the BanggaiSula Block. The Walea Formation represents the basement within the Togian Islands. Its age is not known but it is inferred to be older than the Middle Miocene limestones against which it is faulted. The formation comprises an association of volcanic rocks and subordinate volcanogenic sediments that we suggest represent the products of a submerged volcanic arc rather than an ophiolite, as previously interpreted (Ku¨ndig 1956). A similar association of volcanic rocks and subordinate volcaniclastics is reported within the Papayato Volcanic rocks of the North Sulawesi Province (Elburg et al. 2003; van Leeuwen et al. 2007), and is consistent with the suggestion that the basement of this province continues southwards beneath the archipelago. Volcanic rocks (breccias, pillows and lavas) similar to those of the lower parts of the Walea Formation are also reported from the Cretaceous Balantak Ophiolite of East Sulawesi (Simandjuntak 1986; A. J. Barber, pers. comm. 2009), but they do not show the same association with contemporaneous volcaniclastic sediments. Geochemical and/or geochronological analyses of the Walea Formation, and comparison with the (Middle Eocene to Early Miocene) Papayato Volcanic rocks (North Sulawesi Province of Elburg et al. 2003) and the (Cretaceous) Balantak Ophiolite (East Sulawesi Ophiolite) would help to resolve this issue but the rocks are so deeply weathered that obtaining suitable material has not so far been possible. Field investigations and geochemical analyses suggest that the western end of the bay is underlain

by continental crust (Elburg et al. 2003; van Leeuwen & Muhardjo 2005; van Leeuwen et al. 2007) as far east as 1218E (Fig. 1). This material forms the eastern margin of Sundaland and is probably of Australian origin (van Leeuwen & Muhardjo 2005), but was accreted to Sundaland during the mid-Cretaceous (Parkinson 1991; Parkinson et al. 1998; Hall 2009) and is not part of the BanggaiSulu block. Continental crust probably continues north from the Banggai-Sulu microcontinent beneath the Molucca Sea (Silver et al. 1983b; Watkinson et al. 2010). Beneath Gorontalo Bay earthquake hypocentres (Engdahl et al. 1998) define the southern edge of the westward-subducting Molucca Sea plate. This is a very sharp, almost WNW –ESE, line (Fig. 7a) that we interpret as the former continental – oceanic crust boundary between the Molucca Sea and the Banggai-Sula block. The position of the line implies that continental crust continues north from the Banggai-Sula Islands to the centre of the eastern part of Gorontalo Bay. How far west beneath Gorontalo Bay the continental crust continues is uncertain; the oil that seeps through the ophiolite north of the thrust complex in the East Arm (Ku¨ndig 1956) suggests continental basement may extend at least west to about 1228E (Fig. 8).

Miocene carbonate platform Miocene carbonate rocks are widespread in northern Sulawesi. They include the Middle Miocene limestones of the Peladan Formation reported here, carbonates of the Buol Beds in NW Sulawesi (Ratman 1976; van Leeuwen & Muhardjo 2005), the Salodik Formation within the East Arm (Rusmana et al. 1982), and limestones observed around Palu and the western Toli–Toli region (Sukamto 1973; Norvick & Pile 1976; van Leeuwen & Muhardjo 2005). Benthic and planktonic foraminifera indicate deposition within inner platform/fore-reef shallow marine environments during the late Early to Middle Miocene (van Leeuwen & Muhardjo 2005; this study). Jablonski et al. (2007) report submerged carbonate reefs in Gorontalo Bay based on seismic observations which they interpreted as Oligocene to Middle Miocene in age. The distribution of Miocene carbonate rocks suggests that Gorontalo Bay was an area of extensive carbonate platform deposition during the Miocene. It was probably

Fig. 7. Earthquake hypocentres in Eastern Indonesia based on the dataset of Engdahl et al. (1998). (a) Black crosses denote all hypocentres, those assigned to the westward subducting Molucca Sea Plate are highlighted with blue dots, those assigned to subduction at the North Sulawesi Trench are highlighted in green. Hypocentres associated with volcanism at the Una-Una volcano are shown in purple. Red box denotes the line of section illustrated in (b). (b) North–south cross section though Gorontalo Bay and the Una-Una volcano. Hypocentres associated with volcanism at the Una-Una volcano (purple dots) are notably shallower than those related to the downgoing slab.

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Fig. 7 (Continued) (c) Earthquake hypocentres assigned to the Molucca Sea Plate coloured based on depth. To aid clarity, hypocentres less than 75 km depth are not shown. Colouration shows that the slab dips gently to the NW but is sharply terminated along its southern edge in a steep upturned lip. Black crosses denote hypocentres at depths greater than 75 km elsewhere in the region. They are almost entirely absent in the Banggai-Sula plate.

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Fig. 8. Detailed bathymetry of Gorontalo Bay, modified from Jablonski et al. (2007). Topography based on SRTM (Shuttle Radar Topographic Mission) data (courtesy of NASA, NGA & USGS).

characterized by contiguous shallow marine platforms, but was certainly not a continuous deep bathymetric feature at this time. In west Sulawesi carbonate deposition terminated by the end of the Middle Miocene (van Leeuwen & Muhardjo 2005).

Rapid Pliocene uplift The clastic sediments of the Bongka Formation record localized rapid uplift and erosion of the East Arm in the latest Miocene to Pliocene (Surono & Sukarna 1986; Hall & Wilson 2000), instigating the development of the high (in places .3 km) present-day topography. In some cases, the sudden influx of clastic material may have been directly responsible for the reduction of carbonate areas from large platforms to isolated pinnacle reefs (Jablonski et al. 2007). Uplift has previously been attributed to collision between the Banggai-Sula microcontinent and the East Arm (e.g. Garrard et al. 1988; Davies 1990; Hall 1996; Calvert 2000; Hall & Wilson 2000; van Leeuwen & Muhardjo 2005). This interpretation followed Hamilton’s (1979) proposal of slivers of continental crust moving west from the Bird’s

Head, with westward subduction implied in front of them. However, it has also been suggested that collision between microcontinental blocks and the East Arm began earlier, between the latest Oligocene and Late Miocene (e.g. Audley-Charles 1974; Sukamto & Simandjuntak 1983; Daly et al. 1991; Parkinson 1991; Smith & Silver 1991; Bergman et al. 1996; Milsom 2000; Hall 2002; van Leeuwen et al. 2007; Spakman & Hall 2010). If so, collision significantly predated the rapid uplift and erosion of the East Arm during the latest Miocene to Pliocene, which must have a different cause.

Basin subsidence Seismic surveys (Jablonski et al. 2007) and multibeam surveys of Gorontalo Bay show present-day water depths up to 2000 m in the western part of the basin and .2700 m in the eastern part (Fig. 8). Sediment thicknesses within these areas may be as great as 10 km (Jablonski et al. 2007). There is a bathymetric high area that links the East Arm and the Togian Archipelago, with water depths of between 500 and 1200 m (Fig. 8), which may continue across the entire bay to the North Arm.

TOGIAN ISLANDS AND GORONTALO BAY

This feature appears to have a broadly NW – SE trend. Seismic data have been used to suggest that the basin formed in a predominantly extensional tectonic environment dominated by east –west trending extensional faults (Jablonski et al. 2007). It was interpreted to have formed in the Eocene as a failed rift arm (Jablonski et al. 2007). We infer a much younger, Pliocene age of formation of the deep basin. We interpret deposits of the Bongka Formation (Celebes Molasse) observed in the Togian Islands and the East Arm as distal and proximal equivalents of a Pliocene alluvial fan building out from the East Arm. Seismic data reveal thick (up to 2 seconds TWT (two-way travel time)) north–south trending lobes of sediment that we infer to be submerged parts of this fan (Fig. 9). Prograding fan delta deposits of similar age are also interpreted from elsewhere in the basin (Jablonski et al. 2007). These observations imply that basin subsidence (from close to sea level to present-day water depths of 500 to 1500 m) occurred after deposition of the fan. The age of the Celebes Molasse in the East Arm therefore provides a maximum, latest Miocene to Pliocene age (e.g. Surono & Sukarna

197

1996; Hall & Wilson 2000) for inception of the basin.

Cause of subsidence and uplift The broadly contemporaneous nature of basin subsidence and uplift and erosion at the flanks suggests that these two processes are inherently linked. Together, the rapid latest Miocene to Pliocene uplift (c. 3 km) and subsidence (.2 km) in and around Gorontalo Bay has produced an exceptional total elevation contrast of more than 5 km in less than 6 Ma. The thickness of sediment in the central part of the bay (up to 10 km) suggest much greater differential movements. The North Sulawesi subduction zone probably developed in the last 5 Ma (Silver et al. 1983a; Surmont et al. 1994). We suggest that palaeomagnetic data (Surmont et al. 1994), seismic data (Silver et al. 1983a; Jablonski et al. 2007) and plate tectonic modelling (Silver et al. 1983b; Hall 1996, 2002) indicate that the region has been in extension since the Early Pliocene, with the North Arm moving away from the East Arm. We interpret Global Positioning System (GPS) measurements of present-day motions (Walpersdorf

Fig. 9. Thickness of the sedimentary fill in Gorontalo Bay, modified from Jablonski et al. (2007). Thickness is based on two-way travel-time in seconds (TWT s) between water bottom and basement isochron (Jablonski et al. 2007). Topography based on SRTM (Shuttle Radar Topographic Mission) data (courtesy of NASA, NGA & USGS).

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et al. 1998; Vigny et al. 2002; Socquet et al. 2006) to indicate that this extension continues today. Therefore one possible cause of subsidence is extension of the upper plate that was driven by rollback of the subduction hinge at the North Sulawesi Trench. However, the extremely rapid rates and large amounts of uplift and subsidence in the region suggest that significant flow of lower crust, from beneath the basin towards topographically elevated areas, may also have contributed (Hall 2010).

Young volcanism The rocks of the Lonsio Formation and Benteng Intrusives record young volcanism in Gorontalo Bay during the late Neogene. Although the Togian Islands are in the right position for a subductionrelated volcanic arc ahead of a westward-moving Banggai-Sula microcontinent, volcanism does not appear to be subduction related. Such volcanic activity should have preceded the East Sulawesi– Banggai-Sula microcontinent collision. The youngest age suggested for this is end Miocene (c. 5 Ma) but the dates we have for the Togian Islands and Poh Head volcanic activity are Pliocene or younger. The composition of the volcanic rocks is not typical of most subduction-related volcanism. The Benteng Intrusives are extremely rich in potassium (Fig. 5; Table 3), they are shoshonites using the scheme of Rickwood (1989). Earthquake hypocentres beneath Una-Una volcano (Fig. 7b) show that volcanism is unrelated to subduction beneath the North Arm, being much further west and much shallower than hypocentres related to the downgoing slab. High-K compositions are characteristic of small degrees of partial melting of anomalous (metasomatized or enriched) material in the upper mantle (e.g. Wilson 1989). We infer a similar origin for the Benteng Intrusives and suggest that rapid extensional thinning of the crust beneath Gorontalo Bay caused the upper mantle to rise, decompress and melt. The resulting K-rich melts were intruded into the crust as a series of shallow level stocks and dykes. Present-day high-K volcanism at UnaUna suggests that volcanism has evolved to a relatively less K-rich chemistry, possibly reflecting increased amounts of partial melting, and has moved WNW over time. The tuffaceous rocks of the Lonsio Formation represent the products of extrusive volcanism, reworked during deposition in a shallow marine environment during the latest Miocene and Early Pliocene (N19). They are significantly older (as much as 3 million years) than the Benteng Intrusives, and appear to be derived from a different – or unknown volcanic centre.

Post-Pliocene tectonics Tuffaceous rocks of the Lonsio Formation are also known from the East Arm (Rusmana et al. 1982, 1993; this study), around 150 km SE of the Togian Islands. Following Simandjuntak (1986), we suggest that Poh Head has been offset to the SE along the Balantak Fault. Based on satellite images, field observations and seismic data, we interpret this structure as a steeply dipping, rightlateral, strike-slip fault that can be traced offshore to the east, where it terminates in a zone of dextral transpression (Watkinson et al. 2011). To the west of Poh Head the position of the fault is not known, but it may bend to the north, possibly linking to the fault that we infer between the islands of Walea Kodi and Walea Bahi. The distribution and ages of the volcanic rocks in the Togian Islands and Poh Head could therefore be explained by postdepositional dextral faulting, or by westward migration of the volcanic centre with time.

Conclusions We interpret Gorontalo Bay to be underlain by a composite basement comprising several different tectonostratigraphic provinces. The western end of Gorontalo Bay is underlain by continental crust added to the eastern margin of Sundaland in the mid Cretaceous. The central part of the bay, including the Togian Islands, is underlain by oceanic basement of the Northern Sulawesi Province. It is possible that the area south of the Togian Islands has continental crust at depth, with a thrust contact beneath the Northern Sulawesi volcanic basement and East Arm ophiolite, as suggested by oil seeps through the ophiolite on land. In the Miocene, Gorontalo Bay was an area of extensive carbonate deposition, characterized by contiguous shallow marine carbonate platforms. It was not a significant, continuous, deep bathymetric feature in the Miocene. Instead, broadly contemporaneous flank uplift and basin subsidence give a maximum latest Miocene to Pliocene age for the inception of the deep basin. Volcanism in the Togian Islands is unrelated to subduction that preceded collision of the Banggai-Sula microcontinent. Instead, it records rapid extension of the crust in the Pliocene and Plio-Pleistocene, causing the underlying mantle to rise, decompress and melt. We interpret GPS observations (Socquet et al. 2006) to indicate extension is continuing today and is probably the cause of volcanism at Una-Una. Volcanic activity has migrated west towards Una-Una during the Pleistocene and deposits of the Pliocene volcanic episode may have been offset by dextral strike-slip displacement along the Balantak Fault.

TOGIAN ISLANDS AND GORONTALO BAY

Rapid subsidence associated with crustal thinning was driven by rollback of the subduction hinge at the North Sulawesi Trench. The unusual character of volcanism in the Togian Islands is not due to subduction but reflects crustal thinning and extension. The extreme rates of uplift and subsidence observed in and around Gorontalo Bay (producing an elevation contrast of .5 km) suggest flow of lower crust may also have contributed. The industrial member companies of the SE Asia Research Group Consortium provided financial support for our work. The authors thank Benjamin Sapiie and Alfend Rudyawan (Institute Teknologi Bandung) for facilitating our work in Indonesia. M.A. Forster acknowledges the support of an Australian Research Fellowship provided by the Australian Research Council (ARC) associated with the Discovery grants DP0877274, and additional support from the Research School of Earth Sciences at The Australian National University. eArgon software written by Gordon Lister. We thank Theo van Leeuwen and Moyra Wilson for their reviews of the manuscript.

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Sulawesi. BP Petroleum Development of Indonesia Ltd, unpublished report no. JKT/EXP/0071. Parkinson, C. 1991. The petrology, structure and geological history of the metamorphic rocks of central Sulawesi, Indonesia. PhD thesis, University of London. Parkinson, C. 1998. An outline of the petrology, structure and age of the Pompangeo Schist complex of central Sulawesi, Indonesia. Island Arc, 7, 231– 245. Parkinson, C. D., Miyazaki, K., Wakita, K., Barber, A. J. & Carswell, D. A. 1998. An overview and tectonic synthesis of the pre-Tertiary very- highpressure metamorphic and associated rocks of Java, Sulawesi and Kalimantan, Indonesia. Island Arc, 7, 184–200. Pigram, C. J., Surono & Supandjono, J. B. 1985. Origin of the Sula Platform, eastern Indonesia. Geology, 13, 246–248. Polve´, M., Maury, R. C. et al. 1997. Magmatic evolution of Sulawesi (Indonesia): constraints on the Cenozoic geodynamic history of the Sundaland active margin. Tectonophysics, 272, 69– 92. Pubellier, M., Girardeau, J. & Tjashuri, I. 1999. Accretion history of Borneo inferred from the polyphase structural features in the Meratus Mountains. In: Metcalfe, I. (ed.) Gondwana Dispersion and Asian Accretion. A.A. Balkema, Rotterdam, 141–160. Rangin, C., Maury, R. C. et al. 1997. Eocene to Miocene back-arc basin basalts and associated island arc tholeiites from northern Sulawesi (Indonesia): implications for the geodynamic evolution of the Celebes basin. Bulletin de la Socie´te´ Ge´ologique de France, 168, 627 –635. Ratman, N. 1976. Geological Map of the Tolitoli Quadrangle, North Sulawesi (Quadrangle 2016– 2116– 2117) – Scale 1:250,000. Geological Survey of Indonesia, Directorate of Mineral Resources, Geological Research and Development Centre, Bandung. Rickwood, P. C. 1989. Boundary lines within petrologic diagrams which use oxides of major and minor elements. Lithos, 22, 247–263. Rusmana, E., Koswara, A. & Simandjuntak, T. O. 1982. Preliminary Geological map of the Luwuk Quadrangle, Sulawesi (scale 1:250,000). Geological Survey of Indonesia, Directorate of Mineral Resources, Geological Research and Development Centre, Bandung. Rusmana, E., Koswara, A. & Simandjuntak, T. O. 1993. Geology of the Luwuk Sheet, Sulawesi (Quadrangles 2115, 2215, 2315) – scale 1:250,000. Geological Survey of Indonesia, Directorate of Mineral Resources, Geological Research and Development Centre, Bandung. Sarasin, P. & Sarasin, S. 1901. Entwurf einer geografisch – geologischen beschreibung der Insel Celebes. Kreidel’s Verlag, Wiesbaden, Germany. Silver, E. A., McCaffrey, R. & Joyodiwiryo, Y. 1978. Gravity results and emplacement geometry of the Sulawesi ultramafic belt, Indonesia. Geology, 6, 527–531. Silver, E. A., McCaffrey, R. & Smith, R. B. 1983a. Collision, rotation, and the initiation of subduction in

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in SE Asia triple junction: Sulawesi, Indonesia. Journal of Geophysical Research, 107, ETG7 1 –11. Walpersdorf, A., Rangin, C. & Vichy, C. 1998. GPS compared to long-term geologic motion of the north arm of Sulawesi. Earth and Planetary Science Letters, 159, 47–55. Watkinson, I. M., Hall, R. & Ferdian, F. 2011. Tectonic re-interpretation of the Banggai-Sula – Molucca Sea margin, Indonesia. In: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia–Asia Collision.

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Tectonic re-interpretation of the Banggai-Sula –Molucca Sea margin, Indonesia IAN M. WATKINSON*, ROBERT HALL & FARID FERDIAN SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK *Corresponding author (e-mail: [email protected]) Abstract: High resolution multibeam bathymetric and seismic data from the area north of the Banggai-Sula Islands, Indonesia, provide a new insight into the geological history of the boundary between the East Sulawesi ophiolite, the Banggai-Sula microcontinent and the Molucca Sea collision zone. Major continuous faults such as the Sula Thrust and the North Sula–Sorong Fault, previously interpreted to bound and pass through the area are not seen. The south-verging Batui Thrust previously interpreted offshore to the east of Poh Head cannot be identified. In the areas where the thrust was interpreted there is a north-vergent thrust and fold zone overlain by almost undeformed sediments. Gently dipping strata of the Banggai-Sula microcontinent margin can be traced northwards beneath younger rocks. In the east, rocks of the Molucca Sea collision complex are deformed by multigenerational folds, thrusts and strike-slip faults. There is a series of small thrusts between the leading edge of the collision complex and the foot of the slope. In the west a zone of transpression close to the East Arm of Sulawesi is the termination of the dextral strike-slip Balantak Fault extending east from Poh Head.

The Banggai-Sula microcontinent or block (Fig. 1) lies to the east of the East Arm of Sulawesi in eastern Indonesia within the complex triple junction between the Pacific, Australian and Eurasian plates. Stratigraphic similarities between the microcontinent and the Australian continent led to the idea that it originated from western Irian Jaya (e.g. Audley-Charles et al. 1972; Hamilton 1979) or further east in Papua New Guinea (Pigram et al. 1985; Garrard et al. 1988). It has become a well established concept that the microcontinent was sliced from the Australian northern continental margin in New Guinea and travelled westwards (e.g. Visser & Hermes 1962; Hamilton 1979; Silver & Smith 1983; Pigram et al. 1985; Garrard et al. 1988) along the Sorong Fault, possibly coupled to the Philippine Sea Plate (Ali & Hall 1995; Hall et al. 1995; Hall 1996). The Banggai-Sula microcontinent’s westward movement was arrested by collision with the East Arm of Sulawesi but the timing is debated. It is generally thought to have occurred in the Neogene (Simandjuntak & Barber 1996) but a range of ages has been suggested including Miocene (Hamilton 1979), Early to Middle Miocene (Bergman et al. 1996), Middle Miocene (Sukamto & Simandjuntak 1983; Simandjuntak 1986), Middle Miocene to Pliocene (Garrard et al. 1988), and Late Miocene (Silver et al. 1983; Davies 1990; Smith & Silver 1991; Parkinson 1998).

During the collision, ophiolites were obducted and thrust eastwards over the microcontinent, to form an imbricate collision zone at the east end of the East Arm of Sulawesi (e.g. Ku¨ndig 1956; Silver et al. 1983; Simandjuntak 1986; Davies 1990; Simandjuntak & Barber 1996). Compressional deformation of the Banggai-Sula microcontinent itself, including reactivation of Mesozoic structures on land in the Sula Islands, has also been interpreted by Garrard et al. (1988) to have resulted from the NW-directed (Hamilton 1979; Silver et al. 1983) collision. Although the microcontinent is small, the results of its collision are often considered to extend significantly beyond the immediate zone of orogenesis. Westward thrusting of the central Sulawesi metamorphic belt, a foreland fold and thrust belt in west Sulawesi, magmatism in west Sulawesi, and deformation in the Makassar Strait and Borneo have all been attributed to the collision (e.g. Coffield et al. 1993; Bergman et al. 1996; Simandjuntak & Barber 1996; Pubellier et al. 1999a; Calvert 2000; McClay et al. 2000). Today there is a mountain range over 3000 m high immediately west of the collision zone exposing the ophiolite which remains difficult to explore and as little studied as it was when described by Brouwer (1925) and Rutten (1927). Most studies have been aimed at understanding the tectonic development of the collisional orogen

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 203–224. DOI: 10.1144/SP355.10 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Map showing the location of the study area, the Banggai-Sula microcontinent (pink) and principal tectonic features identified in previous studies. Faults modified after Silver et al. (1983); Garrard et al. (1988). Red lines are seismic lines, shown in bold for sections illustrated in subsequent figures; the blue area is the extent of multibeam data. Note that structures in this map are quite different from those we now propose in Figure 15.

on Sulawesi and at the western end of the BanggaiSula microcontinent, but little attention has been paid to the northern margin of the Banggai-Sula microcontinent. Silver et al. (1983) produced maps based on a range of geophysical and field observations, acquired particularly from marine cruises, showing the position of faults, and these, together with the work of Hamilton (1979), have formed the basis for much subsequent work (Fig. 1). However, since the early 1980s there have been relatively few new observations made in the offshore region around the Banggai-Sula microcontinent. Recently, some new seismic and multibeam bathymetric data have been acquired. This paper presents a structural interpretation based largely on these new data from the area north of the Banggai-Sula microcontinent and the southern Molucca Sea, and some new field observations on land, which provide the basis for a better understanding of the significance of the structures and their regional context.

Tectonic setting East Sulawesi and the Banggai-Sula microcontinent are presently sutured along the Batui Thrust zone (Fig. 1), which accommodated much of the shortening between the two regions during Neogene collision (e.g. Ku¨ndig 1956; Hamilton 1979; McCaffrey et al. 1981; Silver et al. 1983; Simandjuntak 1986; Beaudouin et al. 2003). It is widely considered to be bounded by strands of the

Sorong Fault system. A southern strand of the Sorong Fault, called the South Sula –Sorong Fault was interpreted by Hamilton (1979) to follow the break in slope south of Taliabu and pass between Mangole and Sanana (Fig. 1). The dramatic increase in water depth south of this line suggests that the area is floored by oceanic crust, supported by dredging and marine geophysical observations in the North Banda basin (e.g. Hinschberger et al. 2000), and therefore the fault marks the southern margin of the Banggai-Sula microcontinent. The South Sula –Sorong Fault (Fig. 1) is a splay of the strikeslip fault which can be traced east to Irian Jaya, and which is interpreted to have facilitated the left-lateral westward translation of the BanggaiSula microcontinent (e.g. Hamilton 1979; Sukamto & Simandjuntak 1983; Hall 1996; Simandjuntak & Barber 1996; Villeneuve et al. 2002; Beaudouin et al. 2003). It is often shown to link, via the Matano Fault, to the Palu-Koro Fault of central Sulawesi, sinistral structures which ultimately connect to subduction at the North Sulawesi trench, suggesting clockwise rotation of the block north and east of these faults (e.g. Hamilton 1979; Silver et al. 1983; Walpersdorf et al. 1998; Stevens et al. 1999; Socquet et al. 2006). North of the Banggai-Sula microcontinent, other strands of the Sorong Fault have been mapped. The North Sula–Sorong Fault is traced by Hamilton (1979) from the Bird’s Head peninsula, south of Obi, along the north of the Banggai-Sula Islands,

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and towards Poh Head and is shown in slightly different positions by other authors (e.g. Norvick 1979; Silver et al. 1983; Sukamto & Simandjuntak 1983). Australian continental crust is known from SW Obi (Wanner 1913; Ali & Hall 1995) although most of the island is ophiolite suggesting a splay may pass through Obi. However, no geological evidence of significant strike-slip faulting has been recorded on Obi, and there is little geophysical evidence for the fault in this position. Another strand of the Sorong Fault is interpreted to pass north of Obi and south of Bacan, called the Molucca –Sorong Fault by Hamilton (1979) and may continue towards Gorontalo Bay to the west (Fig. 1). Beneath Gorontalo Bay well located hypocentres (Engdahl et al. 1998) show an abrupt termination of the westsubducting Molucca Sea Plate suggesting that a strand of the Sorong Fault may be traced from south of Bacan into Gorontalo Bay along a line trending about 2858 (Cottam et al. 2011). Thrusts observed south of the North Sula– Sorong Fault by Silver et al. (1983) have been named the Sula Thrust and interpreted to form a continuous north-dipping thrust zone. North of the North Sula–Sorong Fault, almost 10 cm/a convergence between the Philippine Sea and Eurasian Plates is largely accommodated by the Sangihe and Halmahera thrusts, which lie above the double subduction zone in the Molucca Sea (e.g. Silver & Moore 1978; McCaffrey et al. 1980; Moore & Silver 1980; Hall 1987, 2002; Rangin et al. 1996; Pubellier et al. 1999b; Beaudouin et al. 2003). Sediments within this zone are being squeezed southwards, to form the Molucca Sea collision complex (Silver & Moore 1978; Silver et al. 1983). Much of Sulawesi’s present-day seismicity is associated with subduction of the Celebes Sea beneath the North Arm, and Molucca Sea subduction north of the Banggai-Sula microcontinent (Cardwell et al. 1980; Engdahl et al. 1998; Beaudouin et al. 2003). Scattered earthquakes occur outside of these areas throughout Sulawesi and the islands to the east, some of which may be associated with the major Sorong and Matano Faults, and GPS vectors have been interpreted in terms of faultbounded blocks (Socquet et al. 2006). Very little shallow seismicity occurs immediately north of the Banggai-Sula microcontinent (Cardwell et al. 1980; Engdahl et al. 1998; Beaudouin et al. 2003) indicating that there are few active structures in this area, or that deformation is largely aseismic.

Dataset This study is based upon geophysical data acquired from the offshore area immediately north of Pulau Taliabu and Pulau Mangole, part of the Sula Islands east of Sulawesi, Indonesia (Fig. 1).

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Sixteen north– south-trending lines and one east – west-trending seismic line were acquired from January to March 2007 by the M/V Mezen as part of the IndoDeep–Banggai-Sula (BS07) Non Exclusive 2-D survey, on behalf of TGS. The data were recorded by Sercel SEAL instruments using a 3980 cu. in. Sodera G-Gun at 5 m depth with a 25 m shotpoint interval. The seismic data were filtered to remove external noise, de-multipled, and pre-stack time migrated. During March– May 2007 the M/V L’Espoir acquired 40 746 km2 of multibeam data in the same area of which about 22 000 km2 is presented here. This 3D coverage was acquired using a Kongsberg Simrad EM120 Multibeam Echo Sounder using 191 beams at equidistant spacing. Positioning control used a C-Nav Starfire DGPS. During processing, positioning, tidal and calibration corrections were applied, random noise and artefacts were removed, and a terrain model using a 25 m bin size was gridded and exported to ESRI format. Multibeam data were further processed in ERMapper to remove voids and generate digital elevation models (DEMs) in which the azimuth of artificial lighting was rotated in 458 increments through 3608 to illuminate features with different orientations. Images used in this paper are illuminated from the NW and in most a greyscale is used as it more clearly shows structural features. The study area is composed of two distinct zones: in the south, north-dipping seabed (Fig. 2) is underlain by a series of parallel reflections which appear to be continuous with strata onshore in the Sula Islands. This slope area represents the northern margin of the Banggai-Sula microcontinent. From the foot of the slope to the northern extent of the dataset is the second zone, composed of complexly deformed sediments in water depths of 1 km. The relatively shallow, plateau-like, western and eastern parts of the deep area are separated by a deeper central area, and are described separately as the western and NE areas (Fig. 3).

Stratigraphy The Banggai-Sula Islands have a relatively simple stratigraphy (Garrard et al. 1988; Supandjono & Haryono 1993; Surono & Sukarna 1993). The basement is Palaeozoic or older metamorphic rocks intruded by Permo-Triassic granites associated with acid volcanic rocks. These rocks are overlain by undated, probably Lower Jurassic, terrestrial sediments and by Jurassic and Cretaceous marine shales and limestones. In the western parts of the islands are Eocene to Miocene and younger Neogene limestones. On Taliabu the basement rocks form an elevated core to the island with Mesozoic sediments dipping mainly to the north and south.

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Fig. 2. Summary seismic stratigraphy drawn onto seismic line BS07-32. See inset map for location.

Fig. 3. (a) Shaded relief map of the multibeam data. See inset map for location. Illumination from the NW. (b) Interpreted structural map, showing fault kinematics, basin areas, and fields of debris derived from the collapsing slope in the south. Locations of subsequent figures shown.

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Banggai-Sula margin stratigraphy The Banggai-Sula margin and slope is dominated by sub-parallel strata which dip gently north (Fig. 2), and appear to be continuous with Mesozoic sedimentary rocks exposed on Banggai, Taliabu and Mangole, immediately south of the study area. There are no offshore wells, and we have interpreted correlations between seismic packages identified offshore with rocks observed on land to the south. The lowermost seismic package is largely structureless, and is likely to represent crystalline basement. This crops out onshore in the centre and around the southern edge of Taliabu and Banggai. It is composed of folded metasediments, marbles, schists, gneisses and amphibolites of probable Permo-Carboniferous age, intruded by granites and associated acid bodies of Permo-Triassic age (Supandjono & Haryono 1993). Reflections above the basement, which appear to be fault-bounded, are interpreted to be terrestrial conglomerates and sandstones of the Lower Jurassic Bobong Formation within half graben, which rest unconformably on the basement onshore to the south (Garrard et al. 1988; Surono & Sukarna 1993). The strong acoustic contrast between this package and the widespread, onlapping seismic

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package above, are correlated with marine shales of the Middle–Upper Jurassic Buya Formation which overlie the Bobong Formation on land (Supandjono & Haryono 1993). A thick, widespread and weakly reflective seismic package (Fig. 2) above the Buya Formation is interpreted to represent deepwater carbonates of the Cretaceous Tanamu Formation, which lies unconformably above the Buya Formation in the central part of the Sula Islands (Garrard et al. 1988; Supandjono & Haryono 1993; Surono & Sukarna 1993). Platform and reefal carbonates of the Eocene–Miocene Salodik and Pancoran Formations are widespread across eastern Sulawesi and the Banggai-Sula Islands (e.g. Rusmana et al. 1993; Supandjono & Haryono 1993; Surono & Sukarna 1993), and are represented by a package of strong, closely spaced reflections at the top of the slope stratigraphic sequence on the westernmost seismic lines (e.g. Fig. 4a).

Basin floor stratigraphy Most of the material of the basin floor is strongly deformed and lacks coherent seismic reflections. Much of it is almost certainly allochthonous (see discussion in the ‘Basin margin’ section).

Fig. 4. (a) Detail from seismic line BS07-14 showing subsided carbonate platforms (note vertical scale at left). (b) Depth-coloured multibeam DEM showing well developed carbonate platform in plan view, close to the seismic line. See inset map for location.

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Fig. 5. (a) Selection of seismic lines from across the basin margin and slope area from east (top) to west (bottom), showing deformation at the interface between the Banggai-Sula slope sequence (south) and the deformed basin floor stratigraphy (north).

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Fig. 5. (Continued) (b) Interpretations of the seismic lines. Fine dashed lines are reflectors interpreted to be bedding, bold lines are faults. Bold dashed lines are uncertain faults. For seismic line locations see Figure 1.

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However, in the least deformed, northwestern part of the study area, there are submerged carbonate reefs which are similar to those on Peleng and Banggai Islands (Fig. 4b). These are now in water depths of over 1 km implying, if they are part of the same carbonate system as the Peleng examples, that they have subsided substantially since formation. On the basin floor there are a few small basins (Fig. 3b). These are typically bounded by the uplifted hanging walls of thrust faults, and display thin, undisturbed horizontal reflections.

Structural observations The term ‘lineament’ is used throughout for linear or curvilinear features observed on the seabed. Lineaments mapped from the multibeam sea floor imagery (Fig. 3a) have been combined with 3D observations of these features, together with structures seen on the seismic lines, to produce a structural map of the whole study area (Fig. 3b). Discussion of this interpretation is presented below, divided into four structurally discrete regions.

Slope area The slope forming the Banggai-Sula margin area is relatively undeformed (Figs 3b & 5). However, a small number of prominent structures indicate young, gravity-driven deformation. A north-verging (downslope), duplex-like stack of gently folded reflections in seismic line BS07-40 (Fig. 6) occurs in the otherwise structureless strata at the base of the slope. Fine grained, shaly rocks of the Buya Formation acted as the decollement for downslope slip and the duplex-like stack thus represents contraction and overthrusting of mobilized material as it piled up at the base of the slope (Fig. 6). Material derived from the slope has in many cases been reworked by thrusting in the deep part of the basin (e.g. seismic line BS07-26, Fig. 5). A number of NNE-trending lineaments, broadly perpendicular to the slope strike, visible in the multibeam data (Fig. 3a) bound areas of the slope which are topographically lower and smoother than adjacent areas. The lineaments are steep scarps at the edge of areas that have collapsed by slumping (Fig. 7a, b). At the foot of each scarp-bounded smooth area is a debris field extending across the basin floor at the foot of the slope, which includes irregular fragments up to 2 km across (Fig. 7a). These represent bedding-parallel slabs of the northdipping slope that have detached from underlying strata, probably along the same Buya Formation decollement horizon described above, which disintegrated as they moved down the slope to the basin floor. The NNE-trending faults at the margin of the slab therefore had a sinistral strike-slip component

of movement as the hanging wall moved downslope (Fig. 7a, b). The exposed footwall surface of the decollement is smoother than the deeply incised adjacent areas because it has been exposed to marine erosion for a shorter time. Grooves on the slope surface are erosional gullies, not lineations caused by scouring of the footwall by the hanging wall. They are coincident with the slip direction orientation, but are not linked to the collapse, and were probably formed by higher density water (possibly hypersaline) flowing down the newly-formed slope from the shelf. The slope collapse failure surfaces are essentially low-angle normal faults, gravity-driven and probably facilitated by very low friction, possibly within over-pressured shale horizons, and low confining pressure on the north side of the slope. It is notable that major collapse structures in the central part of the slope (Fig. 3b) are associated with large fields of widely dispersed debris, but the eastern collapse shown on seismic line BS07-30 (Fig. 5) is deformed by south-directed thrusts. Reflections within the north-dipping slope represent sedimentary packages which are continuous with rocks exposed on the Sula islands, and are part of the Banggai-Sula microcontinent. Their northward extent can therefore be used to map the extent of the continental fragment. In the west, strongly deformed overlying sediments prevent recognition of the north-dipping reflections much beyond the foot of the slope. However, in the east, the reflections can be traced about 40 km north of the slope foot, to a latitude of 18100 S (e.g. seismic line BS07-38, Fig. 5). Their termination is not observed, and they may continue much further north. This has important implications for the position and role of the Sorong Fault, which has been considered to mark the northern margin of the Banggai-Sula microcontinent (e.g. Norvick 1979; Silver et al. 1983; Sukamto & Simandjuntak 1983) (Fig. 1). This idea is discussed more fully below.

Western area The western area’s bathymetry is dominated by a small number of continuous WNW-trending lineaments that are associated with positive flower structures (e.g. seismic line BS07-22, Fig. 5) and a large number of discontinuous, sinuous, broadly ENE-trending lineaments (Fig. 3a). The three major WNW-trending lineaments have a left-stepping en-echelon geometry, and each is about 50 km long. They are composed of an array of parallel, en-echelon and anastomosing lineaments which are narrow, high amplitude, low wavelength ridges, valleys or steep-sided scarps. The discontinuous, ENE-trending lineaments are concentrated at the eastern ends of the

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Fig. 6. Detail from seismic line BS07-40 showing sigmoidal ‘thrust horses’ verging downslope as a result of slope collapse. See inset map for location. Interpretation below shows faults and slump package-bounding surfaces as bold lines. Fine dashed lines are reflectors interpreted to be bedding.

WNW-trending lineaments (Fig. 8). These mark variable amplitude, but generally long wavelength (up to 5 km), ridges on the seabed. Many are asymmetric, with gentle northern slopes and steep southern faces. In well stratified material these structures have clear expression in seismic lines, and are small, south-verging thrust faults and compressional duplexes (Fig. 9). Most of the bathymetric features are hanging wall anticlines above blind thrusts, but some thrusts propagate to the sea floor, and form the southern fold margins. There is no clear basal detachment but many thrusts curve upwards from bedding surfaces that can be traced through the north-dipping Banggai-Sula margin sequences. Sigmoidal folding of bedding results in a series of thrust ‘horses’, which have total displacements of 100–200 m, measured from offset reflections on seismic lines (Fig. 9). South-dipping back-thrusts occur in the crests of thrust ‘horses’, and elsewhere. In places, undeformed basins filled with horizontally bedded strata lie between the dipping Banggai-Sula margin sequences and the thrust front (e.g. seismic line BS07-18, Fig. 5). Small piggy-back basins occur in valleys between larger thrust-bounded ridges.

In plan view, the anticlines have a lenticular form (Fig. 8), consistent with their formation above short faults whose displacement decreases laterally from central maxima. Some domes are truncated by the WNW-trending structures, whereas the tapered ends of others curve into parallelism with the WNW-trending structures with a dextral asymmetry, leading to their sinuous appearance. There are two possible explanations for the dextral asymmetry and sinuous appearance of the anticlines. Firstly, the folds and underlying thrusts may pre-date a period of dextral slip along the WNW-trending structures. Folds adjacent to the dextral faults were subsequently sheared by locally plastic deformation along the faults, stretching and curving their tapered ends. Alternatively, the folds may have formed above a leading contractional fault array at the termination of dextral faults or as part of a dextral transpressional zone, during a continuous deformation phase. The similarity to an S-C0 deformation fabric, where the folds and thrusts are equivalent to the S-fabric and the dextral shears are equivalent to the C0 -fabric, indicates that the system formed during a continuous deformation phase.

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Fig. 7. (a) Multibeam image showing the footwall of a collapsed part of the slope, its bounding faults, and the resultant debris field. See Figure 3b and inset map for location. (b) Enlarged multibeam view of the margin of the collapsed slope, showing post-collapse extension in the unsupported footwall. See Figure 7a for location.

Regardless of whether the ENE-trending folds are pre- or syn-kinematic with respect to slip along the WNW-trending faults, the faults must have been dextral to produce the observed fold hinge sinuosity. The left-stepover between the two main structures would therefore have a restraining geometry, consistent with uplift and fold intensification which occurs between them. A prominent NNE-trending lineament (Fig. 8) intersects the westernmost dextral fault at an angle of c. 808. It cuts through, and slightly sinistrally displaces a small fold. It is likely that this is an antithetic conjugate structure to the main dextral system. If so, since s1 should bisect the angle between them, it indicates NNWtrending compression during dextral slip (Fig. 8). This direction is perpendicular to the general strike of the fold axes, so folding, thrusting and strike-slip faulting are all kinematically compatible. That folding and thrusting are so strongly associated with kinematically compatible dextral faults suggests that they formed under a dextral transpressive regime. Those folds which are truncated by, or dragged along, the dextral faults may have formed during the early stages of this event, before the

strike-slip faults localized onto discrete strands. Thrusting and dextral slip may be a very young event, as, apart from the piggy-back basins (which may be syn-tectonic), there is no sedimentary drape over these features, and several of the thrust faults pass directly to the seabed. Parts of the Mesozoic and Cenozoic slope sequence can be seen below the deformed material on several seismic lines throughout the area (e.g. Fig. 2 and seismic line BS07-22, Fig. 5).

NE area Bathymetrically, the NE area forms a plateau elevated above the deep part of the basin to the SW. Its southern margins are rounded and lobate, giving it a ‘tongue-shaped’ geometry, and its upper surface is crossed by a dense network of linear features with a variety of orientations (Fig. 3a). The plateau extends to the NE corner of the multibeam data, and represents the southern extent of the Molucca Sea collision complex, which resulted from the Sangihe Arc –Halmahera Arc collision in the central Molucca sea (Hamilton 1979; Silver 1981).

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Fig. 8. Multibeam image showing details of the region of dextral transpression in the west of the study area. See Figure 3b and inset map for location. Antiformal hinge lines marked by black dashed lines, thrusts marked by white dashed lines. Strike-slip faults marked by double half arrows. Maximum horizontal stress orientations for various structures shown in top right.

Sedimentary material caught up in this collision is extruded southwards (Silver & Moore 1978), and the chaotic, highly deformed seismic character of material in the NE of the study area (e.g. the northern end of lines BS07-34, 38 and 42, Fig. 5) is consistent with its having been squeezed out of the collision zone. There are three broad lineament trends on the plateau top: ENE, NE and NNW. Few lineaments can be correlated with structures in the subsurface based on seismic data, because the area is so highly deformed and little stratification remains. However, thrust faults are prominent, together with steeply dipping reverse faults. Many of the latter are arranged into V-shaped arrays with folded strata and seabed within them (Fig. 10), and are geometrically similar to positive flower structures associated with strike-slip faults. Many of the more continuous lineaments on the plateau top are associated with features characteristic of strike-slip faults: arrays of en-echelon Riedel shears; conjugate pairs intersecting at about 608; step-overs associated with small depressions or flat-topped basins or elevated, folded topography;

narrow zones of intense, anatomosing fractures; and terminal splays. The consistency of different features indicates that they are genuine strike-slip faults (Fig. 3b) and the geometry of these features can be used to infer the shear sense, which seems to change with age (Fig. 11). The small size of pullapart basins, the dominance of discontinuous features such as Riedel shears, and the usual absence of a through-going principal displacement zone indicate that strain along these faults is small. There is no evidence in the seismic data that they are the expression in young sediments of more substantial, older faults below. Folds with rounded hinges occur across the plateau, and are mostly cut by the strike-slip faults. They are arranged in a relatively simple curved pattern which follows the plateau margin. Hinge lines on the west and east sides trend north –south, and those on the south side trend east –west. This pattern is repeated within a smaller lobate body near the western side of the plateau (Fig. 11a). Many of the folds around the margins of the plateau, particularly in the south, are large (over 20 km long with 6 km wavelength) and asymmetric, and occur

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Fig. 9. Seismic line BS07-16 and interpretation, showing south-directed imbricate thrusting over the Banggai-Sula slope sediments, and hanging wall anticlines reaching the seabed. See inset map for location.

in the hanging wall of thrust faults. The asymmetry of the folds indicates vergence away from the topographic high. The prominence of an outward-verging fold and thrust belt around the margins of the topographic high indicates that it is propagating outwards, and mainly southwards. Strike-slip faults across the plateau accommodated changes in the rate and direction of flow within the deforming mass. None shows evidence of high strain, and the flow is chiefly taken up by the folds, thrusts, and a possible basal detachment below the plateau. Cross-cutting relationships can be observed between some of the faults on the plateau, which can be used to determine a relative kinematic history (Fig. 11b). The oldest are ENE-trending sinistral and NE-trending dextral strike-slip faults. From these a broadly NE-trending SHmax can be inferred. In the east of the plateau, these are cut by NE-trending sinistral faults (Fig. 12), indicating a

NNE-trending SHmax. A large normal fault near the southern margin of the plateau may be correlated with extension observed in some of the positive flower structures, consistent with a period of relaxation and ESE-trending SHmax. This, and the older strike-slip faults are cut by a set of NNW-trending, mostly sinistral strike-slip faults (Fig. 12). The subordinate lobe at the west of the plateau is bounded by a pair of parallel faults of this trend, but the eastern fault is dextral. Together with a series of thrusts at the southern end of the lobe, these structures indicate that the lobe is moving southwards within the deformed mass. Inferred SHmax orientations for structures bounding this mass radiate outwards from its central axis, mirroring SHmax orientations inferred for folds around the whole plateau (Fig. 11b), which probably formed during all of the kinematic phases outlined above. None of the features cut by faults is displaced by a significant amount, typically less than a few

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Fig. 10. Seismic line BS07-36 and interpretation, showing ‘v’-shaped reverse faults forming a positive flower-structure along a strike-slip fault. See inset map for location.

hundred metres (Fig. 12). This indicates that the faults accommodated minor reorganization within the deforming body, rather than being significant tectonic features.

Basin margin area The basin margin is defined as the area extending from the foot of the slope to about 10 km north across the basin. In the west, it is represented by the region of dextral transpression discussed above. Immediately east of this area, a number of changes occur. There is no evidence of strike-slip faulting or transpression beyond the two major en-echelon faults. Bathymetric lineaments, representing folds and thrusts, curve c. 408 anticlockwise to a more ENE-trend, and may take up the final shortening strain at the ends of the strike-slip faults (Fig. 3b). In the central part of the margin, the basin floor is relatively undeformed. Several small east –westtrending folds associated with south-verging thrusts cut through the debris field formed by collapse of the slope (Fig. 7). In the east, more prominent, but still small, ESEtrending folds are associated with SSW-verging thrusts. These indicate compression from the NNE (Fig. 11b). Their position at the foot of the ESEtrending slope is significant, since the large, south-propagating mass of elevated and intensely

deformed material lies ,30 km to the north. It is likely that there is partial coupling between the deforming mass and the underlying strata of the basin floor. These relatively competent strata were displaced southwards without internal deformation before they buckled and imbricated along small thrusts as they were forced against the foot of the slope. Gently north-dipping strata of the Banggai-Sula margin can be traced below deformed sediments in the deeper part of the basin in all seismic lines (Fig. 5). The gently-dipping contact between these two packages is parallel to bedding in the underlying strata and is clearly tectonic, but its nature is unclear. In the east of the area (e.g. seismic line BS07-42, Fig. 5) south-verging thrusts seem to detach from the contact suggesting it is a basal thrust. In the central part of the area (e.g. seismic line BS07-34, Fig. 5) the contact is a detachment surface for steep south-dipping normal faults. In the west, in the region of the dextral strike-slip system (e.g. seismic line BS07-22, Fig. 5) the Banggai-Sula margin strata are cut by steeply dipping strike-slip faults associated with thrusts and pop-up structures. At the western end (e.g. seismic line BS07-18, Fig. 5) prominent basins filled with undeformed strata lie above the junction between a thrusted sequence to the north (associated with the strike-slip faults) and the Banggai-Sula margin strata to the south.

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Fig. 11. (a) Shaded relief map based on the multibeam data of the collision complex in the NE of the study area. See Figure 3b and inset map for location. (b) Kinematic interpretation of the same area. Inferred faults marked, together with maximum horizontal stress (SHmax) orientations for various structures and deformation generations.

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Fig. 12. Multibeam image showing features of strike-slip faults in the east of the study area, and cross-cutting relationships between different fault generations. Circled numbers represent a relative kinematic chronology, from 1 (oldest) to 4 (youngest). See Figure 3b and inset map for location.

Discussion: implications for major faults in East Sulawesi Previous interpretations of offshore structures were based on widely spaced shallow seismic lines of relatively poor quality, combined with other geophysical data sets such as gravity and magnetic observations, bathymetric data and regional considerations. We have been fortunate in having access to higher quality seismic data, which means that some structures can be more confidently identified, but the distance between lines is still relatively large and in more deformed areas, such as the Molucca Sea collision complex, it is still difficult to interpret to significant depths. The multibeam bathymetric map is a significant aid in correlating between seismic lines and interpreting the structures, and the two new datasets have led to some different interpretations of major structures.

Sula Thrust The Sula Thrust (Fig. 1) is commonly shown as a major structure parallel to the northern margin of the Sula Platform (e.g. Silver 1981; Silver et al. 1983; Garrard et al. 1988). Based on a number of shallow seismic lines, it was described by Silver et al. (1983) as a continuous thrust, to the north of which is an imbricate stack of thrust faults. Silver et al. (1983) considered the Sula Thrust to be the result either of convergence between the

Banggai-Sula block and the area to the north, or compression resulting from the southward extrusion of material from the Molucca Sea collision zone. We do not doubt the presence of the thrusts illustrated by Silver et al. (1983) on two seismic lines, but we believe that the seismic data of TGS, combined with the multibeam data of TGS, show that they are not part of a single major fault zone. Thrust-dominated deformation is present in many places along the northern edge of the Sula platform (Fig. 3b) but the structures have different causes in the east and west. In the west there is a broad zone of discrete, small displacement, south-verging thrusts formed at the southeastern termination of splays of the Balantak Fault in a zone of dextral transpression, discussed further below. In the east there is a zone of SSW-vergent thrusting which is likely to be the result of shortening ahead of the south-propagating Molucca Sea collision complex without a single high strain master fault. As suggested previously (Letouzey et al. 1983; Silver et al. 1983) southwards motion of the highly deformed sediments is due to their expulsion from the region of east –west shortening between converging arcs to the north. We suggest the front of this lobe of material is connected northwards to the Sangihe Thrust as tentatively shown by Silver & Moore (1978). Between these two areas there is little or no thrusting. In the centre of the area, there is a large debris field in a broad depression at the foot of the

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Banggai-Sula slope north of Taliabu. Parts of the Banggai-Sula margin have evidently collapsed into the basin along low-angle normal faults (e.g. Figs 6 & 7). Some of the seismically incoherent material lying above the north-dipping strata could be Banggai-Sula margin sediments but it is not possible to distinguish this from the deformed collision complex sediments. Low elongate ridges curve across the depression and are subparallel to the ENE-trending thrusts in the west and the front of the collision complex to the east. There are small south-vergent thrusts associated with some of these ridges. There must have been a northward dip on the margin before the Banggai-Sula –Sulawesi collision because to the north of it was oceanic crust of the Molucca Sea. Loading by the southwardpropagating collision complex may have contributed to some northward tilting of the parallel strata at the northern edge of the Banggai-Sula microcontinent. In the east there is an elongate WNWtrending trough parallel to the Banggai-Sula margin which deepens eastwards directly south of the collision complex, supporting this idea. However, there is almost no change in dip of the dipping strata from east to west, and in the west where the sediment cover north of the margin is thin, there is evidence of at least 1 km of subsidence of flat-lying carbonates capped by a reef (Fig. 4). Further east in Gorontalo Bay there is evidence of widespread subsidence of carbonate reefs to similar depths based on newly acquired seismic (Jablonski et al. 2007) and multibeam data. This indicates that the subsidence in the west was not caused by collision complex loading. Furthermore, the slope failures and debris fields predate the southward extrusion of the collision complex. Despite the density of deformation in this area, the area immediately north of the Bangaai-Sula margin, and notably the position of the Sula Thrust is almost free of seismicity (Engdahl et al. 1998; Beaudouin et al. 2003). This is a marked contrast to the abundant shallow seismicity associated with the Molucca Sea collisional zone further north. Furthermore, the few earthquakes that have been recorded close to the thrust zone are relatively deep (.20 km) and those for which there are solutions (Beaudouin et al. 2003; Global CMT 2009) are not thrusts. We therefore see no reason to suppose that the Sula Thrust is a continuous through-going fault, that it is a major lithosphere-scale structure, or that there has been north–south tectonic convergence between blocks since the Banggai-Sula–Sulawesi collision.

much of the East Arm. There is an arcuate thrust front (Fig. 1) south of Poh Head and it is usually mapped eastwards towards Balantak. Ku¨ndig (1956) noted the imbricate nature of the thrust SE of the thrust front, and interpreted south to SE-directed vergence. Silver et al. (1983) suggested that the Batui Thrust could be traced 100 km offshore to the east, based on a number of north –south-trending seismic lines, and a steep magnetic anomaly gradient north of the extrapolated thrust. However, the seismic line presented (line 44) by Silver et al. (1983) is of poor quality, and a thrust interpretation is not clear. Seismic line s05-103 from the dataset of TGS (Fig. 13a) is coincident with the position of line 44 of Silver et al. (1983), and shows the same anticlinal feature onlapped to the north by horizontal sediments of Gorontalo Basin. The anticline is formed from a package of highly continuous, parallel reflectors which we interpret to be the same Mesozoic strata that form the north-dipping slope of the north Banggai-Sula margin. Using the data of TGS it is impossible to interpret a north-dipping thrust plane south of, and below, the anticline without cutting continuous strata. Instead, steep, south-dipping normal fault planes deform the southern part of the anticline (Fig. 13b). Two of these faults displace distinctive seismic reflectors and are associated with hanging wall synclines and footwall anticlines consistent with extension. There is no evidence that the Batui Thrust can be traced to this position. Closer to the East Arm, seismic line BS07-20 (Figs 1 & 13c), does show a zone of thrustaccommodated shortening in the position that an extrapolated Batui Thrust might lie. Thrusting is thin-skinned, detaching from horizontal features at a depth of 5 s TWT (two-way travel time). Mound-like features which may be carbonates or fluidized sediments (Ferdian et al. 2010) positioned over the crests of hanging wall anticlines seal the thrusts, and are themselves draped by undeformed layered sediments (Fig. 13d). Unlike the southverging Batui Thrust, faults in seismic line BS07-20 are north vergent, meaning that even if they formed at the same time as the Batui Thrust, they cannot be simply linked to the onshore fault. Silver et al. (1983) argued that a south-vergent Batui Thrust was an active structure which is clearly not the case. The north-vergent structures predate the overlying sediments whose age is unknown, but we speculate that they may be related to the microcontinent-ophiolite collision.

Batui Thrust

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On land the Batui Thrust separates the Banggai-Sula microcontinent from the ophiolite which occupies

In the East Arm the Batui Thrust is often show to curve through Poh Head (e.g. Hamilton 1979;

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Fig. 13. (a) Detail from seismic line s05-103 (see Fig. 1 for location). (b) Interpretation of seismic line s05-103. Fine dashed lines are reflectors, bold lines are faults. Pale blue dashed line shows the thrust fault inferred by Silver et al. (1983). (c) Detail from seismic line BS07-20 (see Fig. 1 for location). (d) Interpretation of seismic line BS07-20. Fine dashed lines are reflectors interpreted to be bedding, bold lines are faults.

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Fig. 14. Digital elevation model created from Shuttle Radar Topography Mission data, of the Poh Head peninsula, eastern Sulawesi. Black arrows point to the ends of a major fault, interpreted to be a strike-slip structure. The Batui Thrust zone occupies the area north of Luwuk. See inset map for location.

Silver et al. 1983). Although Silver et al. (1983) described the fault in this region, there was no structural data to support a thrust fault interpretation. They described low temperature deformation fabrics such as undulose extinction, grain bending, grain boundary slip and outcrop-scale faulting associated with gouge, which increase in intensity toward the fault. Modern imagery shows a fault that crosses Poh Head marked by a single, steep-sided topographic lineament which is notably straight (Fig. 14).

Topography to its north is rugged and mountainous, while relief south of the fault is smoother. If the fault in this area was a thrust, as often shown, it would be expected to be composed of fault traces which curve around the topography due to their low dip, much like the thrusts of the Batui area. The observed straight fault trace is much more indicative of a steeply-dipping or vertical structure, such as a single major strike-slip fault. We interpret it as a right-lateral strike-slip fault, supported by field observations on land (Simandjuntak 1986). This structure trends WNW across Poh Head, and is directly along strike from the WNW-trending offshore faults described above, for which there is abundant evidence for dextral transpressive shear. We suggest it is a dextral strike-slip fault, in keeping with its geomorphic expression (Fig. 14). The structures seen offshore (Fig. 8) would be an expression of dissipation of shear at the end of the fault zone (Fig. 15). Sigmoidal thrusts and folds which link the strike-slip strands observed on the seabed are those previously interpreted (e.g. Silver et al. 1983) to be part of the Sula thrust zone. There are few earthquakes in this area and almost all hypocentres are deeper than 30 km. Fault plane solutions are ambiguous because the fault plane orientation is not known. For the only two shallow earthquakes beneath Poh Head in the CMT catalogue (Global CMT 2009), assuming a fault surface parallel to the Balantak Fault, one (14 km depth)

Fig. 15. Map of the same area as Figure 1, and drawn largely after the same sources, but modified in the light of the present study. Revised faults are shown in red. Principal differences include the absence of a through-going Sula Thrust, the Sorong Fault as a plate boundary which does not reach the surface, and connection of the Poh Head fault to the region of dextral transpression in the west of the study area. Sources of deformation in the region are indicated by regions of colour. See legend and text for details.

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suggests right-lateral movement, whereas a second (12 km depth) indicates a thrust with a right-lateral component. Focal mechanisms reported by Beaudouin et al. (2003) in the area between Poh Head and Taliabu, where the inferred WNW-trending dextral system lies, show sinistral slip on NW-trending structures, or dextral slip on NE-trending planes, which are incompatible with structures observed on the seabed. This indicates that the seabed structures are inactive or that they represent aseismic deformation at shallow levels detached from, and unrelated to, seismic deformation in the underlying crust.

Sorong Fault The Sorong fault is a major east– west-trending sinistral strike-slip fault which cuts across the Bird’s Head of New Guinea, and splits north and south of the Banggai-Sula Islands (Fig. 1). Many authors link the southern strand to the sinistral Matano and Palu-Koro faults of Sulawesi in the west (e.g. Sukamto & Simandjuntak 1983; Simandjuntak & Barber 1996; Villeneuve et al. 2002; Beaudouin et al. 2003). It is thought to form the southern margin of the Banggai-Sula microcontinent. The northern strand of the Sorong Fault has been inferred to lie along the northern edge of the Sula islands, also bounding the microcontinent (e.g. Norvick 1979; Silver et al. 1983; Sukamto & Simandjuntak 1983). Silver et al. (1983) observed structures parallel to their Sula Thrust zone north of the Sula platform which they considered to be the Sorong fault. Interpretation of the new seismic and multibeam data shows numerous strike-slip faults in the elevated pile of deformed sediments north of Mangole, of which many are sinistral and trend east –west to NE–SW (Fig. 3b). These are often associated with ‘pop-up’ structures observed in seismic lines (Fig. 10). The faults lie close to the position of the North Sula–Sorong Fault strand (Fig. 15). It is therefore possible that they mark a strand of the lithospheric Sorong Fault at depth covered by a deforming mass of sediments. However, although some can be traced for up to 50 km, they are generally low strain features, with small pull-aparts, clear terminations and well defined Riedel shears, typical of incipient or very low displacement strike-slip faults. They cannot be traced to significant depths, possibly not even to the base of the sediment pile. These characteristics are not what would be expected above a major structure such as the Sorong Fault. Strike-slip faults of similar character can be found throughout the elevated sediment pile north of Mangole, not just directly over the assumed position of the Sorong Fault. These have both dextral and sinistral slip senses, orientations spanning 1108 anticlockwise from east– west, and

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complex overprinting relationships (Fig. 11). All of these structures are more consistent with complex, entirely upper-crustal, deformation within, and due to, the southward motion of the sediment pile away from the Molucca Sea collision zone, rather than with a major sinistral fault. No east– west-trending strike-slip faults can be identified beyond the collision complex on the main part of the basin floor or basin margin north of Taliabu, suggesting that the Sorong Fault, if it has any expression in the uppermost crust, lies north of the area covered by multibeam data. The WNW-trending strike-slip faults in the western area, discussed above, are clearly dextral, and so are incompatible with the sinistral Sorong Fault. The northward continuity of northward-dipping parallel reflections from the Sula Islands is also inconsistent with the presence of a strand of the Sorong Fault along the northern edge of the Banggai-Sula microcontinent. Even if the fault is presently inactive and has no sea floor expression, it would abruptly truncate the north-dipping reflections. Instead, they continue undisrupted northwards to at least 18100 S (Figs 5 & 15) indicating that the Sorong Fault, if it crosses this area, must lie north of this latitude. Regional well located seismicity (Engdahl et al. 1998) indicates the boundary of the west-dipping subducted Molucca Sea slab is further north and outside the study area, suggesting the Sorong Fault could pass beneath the collision complex in the position shown by McCaffrey (1982).

Greyhound strait fault The NW-trending Greyhound Strait fault (Fig. 1) was identified by Silver et al. (1983) on the basis of a scarp on a north–south seismic line, magnetic data, and the topography of the Greyhound Strait between Taliabu and Peleng/Banggai. It was linked to the NW-trending Gorontalo fault of the North Arm (Katili 1973), which lies along strike, making the structure up to 350 km long. This structure would cross the western part of the multibeam data. However, there is no feature on the sea floor or indication on seismic lines of its presence in this area. This would mean that it lies further west than Silver et al. (1983) proposed, that it is an old structure which has no sea floor expression, or that a fault with this orientation does not exist outside the Greyhound Strait.

Conclusions Our new structural interpretation is shown in Figure 15. Key features are: † North-dipping strata of the Banggai-Sula microcontinent continue below deformed sediments of

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the south Molucca Sea. No linear truncation of the northern Banggai-Sula microcontinent margin is observed, suggesting that it is not cut by the Sorong Fault. No strand of the Sorong Fault can be identified in the study area. If the fault does cross the Molucca Sea it must lie further to the north. A zone of WNW-trending dextral transpression is the termination of the Balantak Fault of the Poh Head peninsula. Complex deformation, dominated by thrusting, folding and small strike-slip faults, occurs within the deformed sediments being forced southwards from the Molucca Sea collision complex. Strike-slip deformation is observed in the area of the collision complex and is entirely related to flow within it. Thrusting at the foot of the northern slope of the Banggai-Sula microcontinent is related to compression between the south-moving collision complex and the Banggai-Sula microcontinent in the east, and in the west to dextral transpression related to the Balantak Fault terminating at the foot of the slope. A through-going Sula Thrust, as previously inferred, cannot be observed. The south-verging Batui Thrust previously interpreted offshore to the east of Poh Head cannot be identified. In the areas where the thrust was interpreted there is a north-vergent thrust and fold zone overlain by almost undeformed sediments.

It remains unclear what is the role and age of the dextral strike-slip system in the west of the area, which terminates the Balantak Fault. There are a few earthquakes associated with the Balantak Fault that suggest it is an active structure. Neither transpressive dextral faulting in the west of the area, thrusting along the southern edge and in the east of the area, nor bedding parallel collapse along the northern Banggai-Sula microcontinent margin, is clearly expressed by modern-day seismicity. The sharply defined sea floor geomorphology suggests young deformation and active structures. Therefore the almost complete absence of shallow seismicity, in contrast to the deforming collision complex further north in the Molucca Sea, is surprising. The few events recorded in the study area north of the Banggai-Sula microcontinent are close to or below the base of the crust. This indicates that the structures mapped on the sea floor are either active but aseismic, or inactive. We are grateful to TGS-NOPEC, who provided the 2D seismic and multibeam data. We thank Fugro (FMCS) and Searcher Seismic for permission to reproduce the seismic line shown in Figure 13a. We thank Chris Elders, Mike Cottam, John Decker and Phil Teas for discussions that greatly improved our understanding of the data and

their meaning. Eli Silver and Manuel Pubellier are thanked for their constructive reviews. This work was funded by the SE Asia Research Group.

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Structural and stratigraphic evolution of the Savu Basin, Indonesia JAMES W. D. RIGG* & ROBERT HALL SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London, Egham, Surrey TW20 0EX, UK *Corresponding author (e-mail: [email protected]) Abstract: The Savu Basin is located in the Sunda–Banda fore-arc at the position of change from oceanic subduction to continent –arc collision. It narrows eastward and is bounded to the west by the island of Sumba that obliquely crosses the fore-arc. New seismic data and published geological observations are used to interpret Australia– Sundaland convergence history. We suggest the basin is underlain by continental crust and was close to sea level in the Early Miocene. Normal faulting in the Middle Miocene and rapid subsidence to several kilometres was driven by subduction rollback. Arc-derived volcaniclastic turbidites were transported ESE, parallel to the Sumba Ridge, and then NE. The ridge was elevated as the Australian continental margin arrived at the Banda Trench, causing debris flows and turbidites to flow northwards into the basin which is little deformed except for tilting and slumping. South of the ridge fore-arc sediments and Australian sedimentary cover were incorporated in a large accretionary complex formed as continental crust was thrust beneath the fore-arc. This is bounded to the north by the Savu and Roti Thrusts and to the south by a trough connecting the Java Trench and Timor Trough which formed by south-directed thrusting and loading.

The Savu Basin is situated in the fore-arc of the Sunda–Banda Arc (Fig. 1) at the margin of the Eurasian plate. It is an area of particular interest since it is immediately north of the position where there is a change from subduction of Indian Ocean crust at the Java Trench in the west to collision between the Australian continental margin and the Banda fore-arc to the east. The Savu Basin has an unusual asymmetrical triangular shape and narrows eastwards from c. 200 km at its maximum width to c. 20 km north of Timor and westwards to c. 50 km north of Sumba. The basin is bounded to the west by the island of Sumba and a submarine ridge that crosses the fore-arc obliquely in a NW–SE direction. To the north is the active volcanic arc including the island of Flores which passes east into an extinct sector of the Banda arc between Alor and Wetar. To the south are the smaller islands of Savu and Roti, and to the east the much larger island of Timor. Timor has been the subject of many studies, notably concerned with collision of the Australian continental margin and the Banda volcanic arc. There is now an unusually short distance between the collision complex on Timor and the inactive volcanic islands of Alor and Wetar to the north. There has been considerable controversy about the significance of the Timor Trough which is significantly shallower than the Java Trench, and in particular whether it is a trench or fore-deep. A sinuous bathymetric trough south of Savu and

Roti and the Savu Basin connects the Timor Trough to the Java Trench. We have studied a recently acquired 2D seismic data set from the Savu Basin. Although there are no wells in the basin it is possible to correlate the stratigraphy offshore with that on land. The islands of Sumba, Savu, Roti and Timor all emerged in the Pliocene or Pleistocene. The stratigraphy and deformation of Sumba have been well documented (Effendi & Apandi 1980; Fortuin et al. 1994) and detailed studies of Savu and Roti (Harris et al. 2009) also provide valuable information that aid in interpreting the offshore data. In addition, marine geophysical investigations combined with gravity and tomographic modelling have recently been used to interpret the deep crustal structure along a regional transect crossing the collision zone and the Savu Basin (Shulgin et al. 2009). The combination of information from these studies with the new seismic data set provides the basis for interpreting the development of the fore-arc basin and the collision complex south of the Savu Basin from oceanic subduction to the earliest stages of arc–continent collision in the context of a new model of Banda subduction rollback (Spakman & Hall 2010).

Seismic stratigraphy The seismic data consists of 32 2D seismic lines the longest of which is 535 km, covering the southern

From: Hall, R., Cottam, M. A. & Wilson M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 225–240. DOI: 10.1144/SP355.11 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Geographical features of the Savu Basin and surrounding area. (a) DEM of satellite gravity-derived bathymetry combined with SRTM topography (Sandwell & Smith 2009). (b) Contoured bathymetry from Gebco (2003) with contours at 200, 100, 2000, 4000 and 6000 m.

and western parts of the Savu Basin, with a total area of 50 000 km2. Two surveys were combined for this study. The first was acquired in 2002 and comprises 2 740 km of long-offset 2D data to a depth of 12 s TWT (two-way travel time). The second was acquired in 2007 and comprises 3 000 km of longoffset 2D data to a depth of 8 s TWT (Toothill & Lamb 2009). The stratigraphic column (Fig. 2) shows the strata present at the eastern end of Sumba (Fortuin et al. 1994) and correlation with the seismic sequences identified in this study.

Unit 1 Unit 1 is the deepest unit and is the sequence below a package of bright reflectors that can be mapped throughout the entire area as the lowest continuous feature identifiable in the dataset (Fig. 3). Below this feature there is locally some reflectivity and the sequence can be split into two parts. The upper part contains some localized reflectors which are sub-parallel, and very bright, within a sequence that has almost no reflectivity. In places there is no clear boundary between the upper and lower parts

of Unit 1 whereas elsewhere there is a sharp boundary, and the lower half of Unit 1 is characterized by moderately bright, laterally discontinuous reflectivity which indicates bedding and gives an overall mottled appearance to the sequence. Bedding can be traced for distances of up to 20 km. The faults that cut the top horizon cannot be traced into the sequence below. There are commonly between three and five parallel reflectors at the top of Unit 1 with a maximum thickness of 0.4 s TWT. They are offset by extensional faults, by a maximum of 0.3 s TWT, and are locally rotated. The tops of fault blocks vary in depth from 2 s TWT close to Sumba to a maximum of 7.4 s TWT in the deepest parts of the basin.

Interpretation The oldest rocks reported from SE Sumba are Cretaceous marine siltstones and sandstones which include volcaniclastic interbeds (Fortuin et al. 1997) and represent submarine fan deposits (von der Borch et al. 1983). They are unconformably

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Fig. 2. Seismic stratigraphy of the Savu Basin and correlation with the stratigraphy of east Sumba after Fortuin et al. (1992). The colours shown next to the seismic section for the four units are used on subsequent seismic profiles. The vertical scale of seismic sections is two-way travel time (TWT) in seconds.

overlain by shallow marine to non-marine Palaeogene sandstones, limestones and volcanic agglomerates (von der Borch et al. 1983). We interpret the lower part of Unit 1 to represent the deeper marine Cretaceous sequence and the upper part to be the Palaeogene in age. The seismic character of Unit 1 is consistent with the descriptions of these rocks on land. The bright package at the top of Unit 1 is suggested to be Eocene –Oligocene or possibly Lower Miocene Nummulites limestones (von der Borch et al. 1983; Fortuin et al. 1992, 1997; van der Werff et al. 1994). The field relationships described by Fortuin et al. (1992, 1994) from Sumba with tilted fault blocks capped by discontinuous carbonates, all overlain unconformably by Neogene strata, are well matched by the observations from the seismic data. Fortuin et al. (1994) suggest these relationships correspond to breakup of a carbonate platform and rapid subsidence in the early Middle Miocene. Fortuin et al. (1994) interpret a Late Burdigalian unconformity with a thin sequence of conglomerates that contain

older shallow water limestones passing up into chalks and marls. We suggest these probably correspond to the top of Unit 1.

Unit 2 This unit is up to 1.6 s TWT in thickness and fills the depressions created by the extensional faulting of Unit 1, onlapping the fault surfaces (Fig. 3). Reflectors are most clear where the unit is thinner and are typically bright, subparallel and discontinuous, forming a complex bedding pattern, whereas the thicker parts of Unit 2 are characterized by a more transparent, less distinctive seismic character. In some places a series of three or four very bright parallel reflectors are visible in the middle of the unit and stand out against the more transparent areas. The seismic character of Unit 2 suggest a facies change which could be from thin bedded lithologies such as carbonates to a more uniform lithology. Overall the amplitude of the reflections increases towards the top of this unit.

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Fig. 3. Uninterpreted and interpreted seismic line from the northern edge of the Sumba Ridge illustrating the relationships between Units 1, 2 and 3. Inset map shows location of seismic lines used in this study and the location of this line in red.

The top of this package is truncated by a prominent unconformity; this surface has extremely strong reflectivity and can be mapped throughout the basin (Fig. 3). The unconformity cuts across bedding in the underlying unit at a low angle. It is now an irregular surface but this appears to be the result of younger contractional deformation.

Interpretation Fortuin et al. (1992, 1994) record a significant change above the thin Late Burdigalian sequence to volcaniclastic turbidites, which we suggest corresponds to Unit 2. All these rocks are poorly dated due to reworking and absence of microfossils. Dates from the volcaniclastic turbidites (Fortuin et al. 1994) are from NN5 (14.8– 13.5 Ma) to NN11 (8.3–5.5 Ma). Fortuin et al. (1997) suggest that volcanic input waned during the Tortonian (11.5–7 Ma) and revived for a short time during the Messinian. Deposits commonly contain abundant fragments of pumice, lava and orientated plagioclase, along with broken crystals of volcanic zoned plagioclase, anorthite (40–60%), volcanic quartz, augite, orthopyroxene, amphibole, hematite and minor amounts of hornblende, all of which points

to an association with a typical island arc (Fortuin et al. 1994). The changes in thickness of this volcanogenic sequence on land between central and east Sumba resembles variation seen on the seismic lines. Thickness is up to 1.6 s TWT, and locally well bedded parts of the sequence pass laterally into unreflective packages consistent with rapid deposition of volcaniclastic material as debris flows and turbidites. Fortuin et al. (1997) suggested that formation of a basin slope, facilitated by NE –SW orientated faults, led to subsidence below the carbonate compensation depth in east Sumba during the Middle Miocene and the Late Miocene in central Sumba. Dissolution of carbonate, but the presence of calcareous nannofossils in some samples, indicate depths of 4 to 5 km (Fortuin et al. 1994, 1997). Fortuin et al. (1997) suggest these deposits were derived from the south and form part of a 100 km fan that progrades northward.

Unit 3 This unit is a particularly distinctive at the top of the seismic lines crossing the southern part of the Savu Basin. It is characterized by a series of flat sheet-like

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deposits, picked out by alternations of very bright and lower amplitude reflectors (Fig. 3). The reflectors become less bright towards the top of the section. The unit typically has a thickness of about 1 s TWT (maximum thickness of 1.8 s TWT) and isochron maps suggest it was derived from the south because it thickens north into the deep Savu Basin from the present-day NW–SE trending submarine Sumba Ridge, and then becomes thinner still further north in the basin. Unit 3 appears to downlap onto the underlying unconformity to the north of the Sumba Ridge (Fig. 4). On the Sumba Ridge Units 3 and 2 are broadly conformable, and on the few undeformed sections south of the Sumba Ridge Unit 3 onlaps the unconformity. Onlap between the individual layers of this unit can be seen throughout although there is no apparent pattern. A number of slumped packages are present within Unit 3, some of which are very well imaged by marked contrasts in reflector character. They are commonly between 0.1 and 0.2 s TWT in thickness, at different levels in Unit 3, and are localized features that can often only be correlated within a 25 km radius. The slumped packages are characterized by a more transparent seismic character than the overlying and underlying well bedded sequences with high amplitude reflectors. There are also a number of more discrete slumped units within the

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top part of this unit, which are on average 0.5–1 s TWT in thickness. These are slumps with a well defined base, which detached at a bedding plane, and at their distal end the basal thrust ramps up against underlying sediment. Their lack of internal seismic character makes them stand out from the sheet-like deposits that make up most of Unit 3. This unit is also cut by many small displacement high angle normal faults, with spacing of 8–10 km.

Interpretation Unit 3 is interpreted as Tortonian to Pliocene, with foraminiferal chalks and marls containing varying amounts of hemipelagic nannofossil oozes and volcanogenic muds (von der Borch et al. 1983; Fortuin et al. 1992, 1997). There are also some thickly bedded volcanic mass flow deposits. The increase in foraminiferal chalks up-section is probably the reason why reflectors become less bright towards the top of the unit. We suggest the apparent downlap north of the Sumba Ridge is the result of uplift of the ridge rotating the former onlapping sequence, supported by the relatively steep dip of beds (up to 5– 68) which for these fine grained sediments is unlikely to relate to an original depositional slope. The onlap between the layers within the correlative unit on Sumba

Fig. 4. Uninterpreted and interpreted seismic line from north of Savu showing apparent northward downlap and tilting of sediments of Units 2 and 3 away from the elevated Sumba Ridge.

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was observed by Fortuin et al. (1997) and is suggested to be the result of shifting fan lobes. The increasing abundance of slumps in the upper parts of the unit suggest that uplift of the Sumba Ridge occurred during deposition of Unit 3. The slumped units with downslope ramps resemble features described by Bull et al. (2009) interpreted as due to pre-existing weaknesses in the mechanical properties of the basal shear surface or stresses generated as material moves downslope.

Unit 4 The boundary between Units 3 and 4 is often unclear, but in places onlap between these sequences can be recognized. The strata of Unit 4 are characterized by a weakly reflective seismic background within which numerous erratically distributed and randomly orientated high amplitude reflectors can be seen. Unit 4 is universally present at the base of the slope north of the Sumba Ridge. Isochron maps indicate that material was derived from the south and movement into the basin was facilitated by extensional faulting at the top of the slope, which can be seen clearly throughout Unit 3. Four phases of slump infill can be identified (Fig. 5). These are separated by minor unconformities and backstep towards their source as the basin was progressively infilled. Large blocks are often present within these deposits which in places retain some of the bedded character of the original deposit.

In the northern part of the basin, deposits are better bedded with flat, parallel and bright reflectors which are less prominent at greater depths. These interfinger with the slumps. The well bedded sequences could represent turbidite deposits at the distal ends of debris flows but could also represent material carried into the deeper parts of the basin from a different source. An important point is that basin infill varies significantly from east to west. In the SE, close to Timor and Roti, Unit 3 is thin (up to 0.25 s TWT) and Unit 4 is thick (about 1 s TWT). Close to Sumba the thicknesses of Units 3 and 4 are reversed. The sediment slides associated with Unit 4 increase in importance towards the east.

Interpretation These deposits formed by simultaneous uplift and downslope transport of debris into the northern part of the Savu Basin during the early Pliocene (van der Werff et al. 1994). Ongoing tilting is fundamental to the generation of the 4000 km2 of superimposed slumps seen on Sumba (Fortuin et al. 1992; Roep & Fortuin 1996) and on offshore seismic lines. These deposits are associated with acoustic voids, which is a direct result of water expulsion, caused by a considerable overburden being deposited during a short time period (van Weering et al. 1989a). The thinning of Unit 3 to the east in the Savu Basin reflects a greater distance from the sediment

Fig. 5. Uninterpreted and interpreted seismic line crossing the Savu Basin showing the units identified in the deeper part of the basin. On this section Unit 2 is thin and Unit 3 is missing in parts of the section. In contrast Unit 4 is much thicker and can be subdivided into several sub-units.

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source which was the Sumba Ridge. In contrast, the greater thickness of Unit 4 in the east is likely to be linked to increased uplift and subsequent slumping from Timor. The well bedded sequences in the most northern parts of the basin may have been derived from the Banda Volcanic Arc to the north, or represent finer material carried further into the basin as part of the slumped packages from the rising Sumba Ridge. There is some interfingering of these bedded deposits with Unit 4 and this favours a Banda Arc origin. There are a number of slumped horizons at the very top of the seismic section closest to the island of Flores and these almost certainly relate to downslope deposition from the Banda Arc.

Structures The principal structures mapped from the seismic dataset are shown on Figure 6. These are described from north to south.

Savu Basin Extensional faulting is clearly visible on the northern edge of the Sumba Ridge, and there are numerous normal faults further north into the Savu Basin. Stretching of the Savu Basin, facilitated by rotation on domino style fault blocks, created a series of full and half graben. The main normal faults are commonly 15 km apart, and although the extension direction cannot be precisely determined because of the wide spacing between seismic lines, it is broadly north–south. The faults have not been inverted. Displacement on these high angle faults has contributed to the subsidence of the top of Unit 1 from depths of 2 s on the ridge to 7 s TWT in the deepest parts of the basin. There has been tilting of Units 2 to 4 which in part reflects uplift of the Sumba Ridge but also greater subsidence in the northern part of the basin.

Savu Thrust The Savu Thrust is actually a zone of thrusting at the northern margin of Savu Island (Harris et al. 2009). The faults dip south and at the rear of this zone some have displacements of more than 2 s TWT (Fig. 7). The most important thrust reaches the seabed whereas others are blind. The geometry of the main Savu Thrust is obscured by poor seismic quality in uplifted areas. There are numerous smaller reverse faults both behind and in front of the main thrust, which in some places break through to the surface. Immediately in front of the main thrust there is a footwall syncline in Unit 3 associated with minor thrusts that do not reach the seabed. Folds further north of the main thrust are associated with

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inversion of older extensional faults that have been reactivated as blind thrusts which cut up section to Unit 3. Behind the main thrust, faults and associated folds are progressively steepened, leading to the formation of back thrusts. Fault propagation folds in the form of hanging wall anticlines and footwall synclines can be seen above the tips of all thrusts. No bends in the fault planes can be seen and there are steep front limbs and some forelimb thinning (Fig. 7). The main Savu Thrust is two separate thrust traces offset by roughly 5 km just east of Savu (Fig. 6). The western thrust can be traced on land into north Savu (Harris et al. 2009) and the eastern thrust is offset to the south and can be traced from the eastern side of the island. The displacement is greatest in the middle of each of these faults, and diminishes to the east and west. There are several episodes of deformation associated with the thrust zone, marked by prominent unconformities in deformed Unit 3. The first major phase of movement folded deeper parts of Unit 3 and folds are onlapped by younger reflectors. There are numerous subsequent subtle onlaps of reflectors near the top of Unit 3, associated with syndeformation thickening away from the main fault.

Roti Thrust This thrust zone (Fig. 8) is located to the NW of Roti and has a similar style to the Savu Thrust. It is located offshore 20 km north of Roti, has the same NNE –SSW orientation as the island, and runs parallel to it for its entire 70 km length. The faults dip southward and again displacement is greatest in the centre of the fault zone. The thrust zone is about 25 km south of the eastern strand of the Savu Thrust. The thrust zone is wider but the maximum displacement on the most important thrusts is less than in the Savu thrust zone, at roughly 1 s TWT. The net displacement on both zones appears comparable (Fig. 6). Like the Savu Thrust, several phases of deformation can be identified by mapping of minor unconformities in the deformed zone with onlap and thickening of packages in front of the main thrust. A difference between the two thrust zones is that the Roti thrust is associated with fewer backthrusts and there are more inverted normal faults in front of the main thrust (Fig. 8).

NE – SW thrusting south of Sumba On the south side of the Sumba Ridge, in the area SE of Sumba there are numerous northward dipping thrust faults. They are associated with a number of anticlines which are very well imaged on seismic lines and are 1000–1400 m across. The orientation

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Fig. 6. (a) Location of seismic lines in the Savu Basin area. The parts of seismic lines shown in this paper are marked in green with the corresponding figure number in square brackets. (b) Summary structure map with the principal tectonic features identified in the area of study. The coloured shaded area north of the Java Trench and Timor Trough is, from west to east, the transition from the accretionary fore-arc complex, which passes east into a deformed zone including fore-arc and Australian sedimentary cover forming the Savu–Roti Ridge, into the arc– continent collision zone of Timor. The dashed blue line crossing the Savu– Roti Ridge and Timor is the approximate position of the former Banda Trench. The dotted black line north of Timor is the inferred northern limit of continental basement. The dashed red line below Sumba is the inferred northern and western limits of the subducted Scott Plateau. See text for discussion. The heavy black line marks the position of the section drawn in Figure 10.

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Fig. 7. Seismic section showing the Savu Thrust north of Savu Island. The upper part of Unit 3 could correspond to Unit 4 in the deeper parts of the basin.

of faults and fold axes is difficult to determine because of the wide spacing of seismic lines in this area but they are clearly not parallel to the Sumba Ridge and our best estimate is that they have a trend of about 0308 (Fig. 6). This zone of thrusting is roughly 60 km across and is overlain by an almost undeformed sequence of sediments which dips northwards and is almost 1 s TWT thick. This sequence appears to be the equivalent of Unit 3 in the Savu Basin and sediments were probably derived from the NW, originally thinned to the SE, and have now been tilted by the rise of the Savu– Roti Ridge. The zone of thrusting, which is now a sedimentary basin, widens towards the Lombok basin to the west.

similar to those seen to the south of Sumba. There are two particularly clear zones of deformed material, shown on Figure 6. The southern zone is characterized by small, well imaged northwarddipping thrusts, which become progressively rotated and steepened towards the north. The majority of faults within the northern, slightly broader, zone of deformation closer to Sumba dip north, although some smaller antithetic faults can also be seen. Between these two zones some faults and reflectors can be seen in places but overall the Savu –Roti Ridge appears to be a seismically opaque mass of deformed material.

Discussion NE – SW thrusting within the Savu – Roti Ridge Within the Savu –Roti Ridge are numerous thrust faults which are visible in the upper 1 s TWT of the seismic lines. The trend of these faults is very

From the Eocene to the Early Miocene there had been subduction of oceanic lithosphere at the Java Trench as Australia moved north (Hall 2002). The region around Sumba was situated at the SE

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Fig. 8. Seismic section showing the geometry of the Roti Thrust, 25 km NW of Roti Island. The upper part of Unit 3 could correspond to Unit 4 in the deeper parts of the basin.

corner of Sundaland in the Early Miocene, at the eastern end of the Sunda Arc, when Australian continental crust of the Sula Spur began to collide with the North Sulawesi volcanic arc. South of the Sula Spur was the Banda embayment, an area of Jurassic –Cretaceous ocean crust within the Australian continental margin with Timor on its south side. At about 15 Ma there was a major change with initiation of new subduction in the Banda region (Spakman & Hall 2010). The Java Trench became aligned with southern side of the Sula Spur and oceanic lithosphere of the Banda embayment began to subduct due to its negative buoyancy and the subduction hinge rolled back to the SE, forming the west-plunging lithospheric fold defined today by seismicity. Figure 9 shows our interpretation of the Savu Basin in the context of subduction in the Banda region. Collision of the volcanic arc with the Australian continental margin in Timor began at about 4 Ma (Audley-Charles 1986; Hall 2002). We suggest that the SE corner of Sundaland which includes the region around Sumba and the Savu Basin was underlain by continental crust that had been accreted to Sundaland in the mid

Cretaceous (Hall et al. 2009). This suggestion is supported by crustal thicknesses and densities for this region presented by Shulgin et al. (2009). Volcanic activity between the Paleocene and Eocene marked a brief phase of subduction which ceased during the Oligocene when shallow water carbonates were deposited on Sumba. We interpret the oldest rocks, which make up Unit 1 in the Savu Basin, to represent the Late Cretaceous to Early Miocene interval. We correlate the distinctive bright reflectors at the top of Unit 1 with the Palaeogene or Lower Miocene shallow marine limestones described on Sumba (von der Borch et al. 1983; Fortuin et al. 1994, 1997; van der Werff et al. 1994). This horizon can be traced across the whole of the Savu Basin implying that the whole region was close to sea level in the Early Miocene.

Subsidence and volcanism Burollet & Salle (1982) suggested subsidence on Sumba began in the Early to Middle Miocene. Based on dating by Fortuin et al. (1994) it appears to have begun at about 15 Ma with breakup of a carbonate platform leading to rapid subsidence in the

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Fig. 9. Cartoon showing evolution of the geometry and tectonic features of the Savu Basin and surrounding region from the Middle Miocene to the present-day. Inset maps show reconstructions of the Banda region based on Spakman & Hall (2010). See text for detailed discussion. In the Middle Miocene the formerly shallow marine are subsided as trench rollback began. The dashed blue line continuing east from the Java Trench on the 3 and 0 Ma maps is the inferred position of the former Banda Trench. By 3 Ma continental crust of the Australian margin had arrived at the trench, and today has been subducted beneath the Sumba Ridge.

early Middle Miocene (Fig. 9). The shallow water Oligo-Miocene limestones have equivalents in the east, on Timor, in the Cablac Limestone (see Audley-Charles 2011), implying that most of what is now the Banda fore-arc between Sumba and East Timor was at sea level before the early Middle Miocene. On Sumba, subsidence coincided with a significant change to volcaniclastic turbidites, marking the beginning of volcanic activity (Fortuin et al. 1992, 1994), that we correlate with Unit 2 offshore. Volcanic input is recorded on Sumba from NN5 (14.8– 13.5 Ma) to NN11 (8.3– 5.5 Ma) with a possible decline in the Tortonian (11.5–7 Ma). Fortuin et al. (1997) suggest these deposits were derived from the south and form part of a fan that

prograded northwards across what is now east Sumba and the offshore region to the east, which implies a short-lived volcanic arc to the south of Sumba. We agree that material was derived from the south but there are several problems with the southern arc interpretation. It would have been unusually close to the subduction trench (van der Werff et al. 1994), significantly south of other Middle Miocene volcanic activity along the Sunda Arc from Java to Sumbawa, and would require an exceptionally steep dip on the subducted slab from the trench to the typical 100þ km depth to the Benioff zone for volcanoes between Java and the Banda Arc at the present day (England et al. 2004). The shift of the position of the arc to its current position in Sumbawa–Flores would

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require a dramatic reduction in slab dip to its present angle. An alternative is that volcaniclastic input may have originated in the volcanic islands to the NW of Sumba and was transported by movement of material as turbidite flows SE through the Lombok Basin and then northward into the Savu Basin. We suggest that Sumba, although submerged, was consistently a relative bathymetric high in the fore-arc approximately parallel to the present orientation of the Sumba Ridge from about the Middle Miocene onwards. West Sumba was at very shallow depths through the Neogene (Fortuin et al. 1997) and to the west of Sumba sediment thicknesses are much thinner than further west in the Lombok Basin (van Weering et al. 1989b), implying an east – west to NW –SE orientated ‘proto-Sumba Ridge’ extending from Sumbawa to East Sumba. NE–SW orientated faults in Sumba, which allowed East Sumba to subside to depths of 4 to 5 km (Fortuin et al. 1994, 1997) before the Tortonian, and would have opened a passage into the Savu Basin from the SW (Fig. 9). Because of the orientation and spacing of seismic lines it is difficult to be certain of the orientation of extensional faults. Fortuin et al. (1994, 1997) observed NE–SW faults on Sumba. Offshore to the north of Sumba, based on those that can be correlated between seismic lines, faults appear to be oriented close to north –south. In the Savu Basin extensional faults have a broadly east– west trend, curving from WNW in the west to ENE in the east. The Sumba Ridge is a feature with inverted normal faults on its north and south sides indicating WNW–ESE trending extensional faults continued as far east as the present longitude of Savu (Fig. 9). We suggest the WNW-trending Sumba Ridge is parallel to a deep basement trend that was inherited from Australian continental basement which accreted to the SE corner of Sundaland in the mid Cretaceous (Hall et al. 2009). In contrast, north–south faults close to the volcanic arc were probably formed by along-arc extension, and we interpret the east –west to ENE-extensional faults in the basin to have formed in response to subduction initiation and rollback into the Banda embayment which began at about 15 Ma, based on regional arguments (Hall 2002, 2009, 2011; Spakman & Hall 2010). Fleury et al. (2009) report unpublished K– Ar ages of 16 Ma from Hendaryono (1998) for the oldest volcanic rocks on Flores consistent with this age estimate.

Thrusting The oldest thrusts are those north of the Savu–Roti ridge to the SE of Sumba. These have a ENE –WSW

trend and are now covered by up to 1 s TWT of sediment. It is not possible to correlate these sequences with the Savu Basin as there are no lines that cross from SW to NE of the Sumba Ridge but the seismic character of the thrusted sequences is very similar to Unit 3. We interpret these as structures formed in the accretionary complex north of, but close to, the former Banda Trench. Thrusts with similar strike are found in the Savu– Roti Ridge within narrow zones all within a much broader zone of seismically opaque, apparently highly deformed material, that forms most of the ridge and resembles accretionary complexes close to subduction trenches. The Sumba Ridge has clearly been elevated during the collision process. The uplift of the ridge postdates the thrusts south of Sumba which are overlain by up to 1 s TWT of sediment but predates the Savu and Roti Thrusts. This uplift clearly postdates Unit 3 which now dips northwards from the ridge. Unit 4 in the Savu Basin includes numerous slumped sequences that have moved northwards from the ridge. There is no observable inversion of normal faults, or thrusting associated with the uplift of Sumba Island and the Sumba Ridge which appears rather to have deformed as a broad upwarp. We suggest this marks the first arrival of the Australian margin at the Banda Trench in the Sumba region, probably at about 2–3 Ma, which postdated collision of the volcanic arc and Australian continental crust in East Timor. In the Savu Basin there are slumps in Unit 4 which are directed basinward from Sumba and Timor but it is not possible to identify relative timing from the seismic data set. East Timor uplift began at about 4 Ma (AudleyCharles 1986; de Smet et al. 1990). On Sumba there was uplift from depths of more than 5 km to emergence of up to 1 km above sea level since 4 Ma (Pirazzoli et al. 1993; Fortuin et al. 1997). The Savu and Roti Thrusts are much younger than all other structures and are currently active. Strictly speaking, these are zones of deformation that include multiple thrusts rather than single faults. Both zones include thrusts that emerge at the sea floor and blind thrusts that are associated with deformation of the sequence above and folding of the seabed. The two main thrust fault segments associated with the Savu Thrust identified on the seismic lines can be correlated onshore with fault segments identified by Harris et al. (2009) which generate topography on the island itself. According to Harris et al. (2009) there are up to seven forward propagating limbs, some of which break through, closely resembling what is seen on the seismic lines. Harris et al. (2009) suggest that most structural features are generally oriented ENE –WSW, sub-parallel to the structural grain of the Scott Plateau and Australian continental margin

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to the east. This is true close to Savu but over a larger area the Savu Thrust zone is closer to east– west. North of Roti the second thrust zone is very similar to that north of Savu, but has a different NNE–SSW orientation, broadly parallel to the island. This too is a zone of active thrusting. The Savu and Roti Thrust are separated by about 25 km and do not link up. Both die out to the east and west. Roosmawati & Harris (2009) show that significant and rapid uplift from water depths of more than 2 km to emergence of the islands of Savu and Roti began at about 2 Ma and they are the first parts of the accretionary wedge west of Timor to emerge. This is supported by the seismic evidence which shows that the area south of the two thrust zones is generally seismically opaque with localized zones in which north-vergent thrusts can be recognized. We assume that this accretionary zone includes material from both the Banda fore-arc and the deep Australian continental margin, and this is also suggested by observations on land in Roti and Savu (Harris et al. 2009; Roosmawati & Harris 2009). The seismic lines show clearly that the Savu and Roti Thrusts are young and not lithosphere-scale features as they die out rapidly to the east and west. There are no other thrust zones north of them in the Savu Basin, which rules out the continuous major thrust zone that is often traced from East Timor to the south of Sumba (e.g. Fortuin et al. 1997; Audley-Charles 2004; Shulgin et al. 2009). In the Savu Basin the only contractional deformation seen on the seismic lines are slumps in Unit 4, and there are no significant features on the seabed which is essentially flat or gently sloping. It appears that the accretionary complex that formed north of the Banda Trench during oceanic subduction has been overridden and/or incorporated in the wedge of deformed material south of the Savu and Roti Thrusts, except in the extreme SE where it is seen beneath flat lying sediments south of Sumba. The Benioff zone thrust of the Banda subduction zone has no surface expression and is beneath the Savu– Roti Ridge. Most of the convergence between the Australian continent and the Banda volcanic arc is concentrated in a zone of contractional deformation about 120 km wide within which are Savu and the Savu–Roti Ridge, and further east the island of Roti. The southern limit of this zone is a trough where there is south-directed thrusting which can be traced west to the Java Trench and east to join the Timor Trough and is entirely within the Australian margin. East of Roti the contractional zone becomes wider and more substantial with elevations on Timor of more than 3 km and the zone of thrusting has overridden the former fore-arc to the north and reduced the distance to the former

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volcanic arc. This is the reason for NE-narrowing of the Savu Basin. The shape of the Australian continental margin was the cause of the deformation history and the complex shape of the collision zone. In the region south of Savu and Timor there were rectilinear steps in the continental– ocean boundary similar to those seen today at the Scott and Exmouth Plateaus of the NW Australian shelf. The first volcanic arc – continent collision was north of East Timor and deformation propagated west with time. Harris (1991) and Roosmawati & Harris (2009) suggested that this sector of the collision can be understood in terms of the present convergent rate between SE Asia and the Australian plate, which is broadly true. However, in addition to the shape of the continental margin and the plate convergent vector it is also important to recognize that there was a component of convergence due to the SE-rollback of the Banda Trench into the Banda embayment (Spakman & Hall 2010). After collision in East Timor most oceanic crust had been subducted as far as the west edge of the Scott Plateau, whereas to the east a significant area of oceanic crust remained which has been subducted since 2 Ma.

Cross section The cross section (Fig. 10) illustrates the key features of the convergence between the Scott Plateau and the Savu Basin. The section is located along one of the long north–south seismic lines just east of Savu and is close to the line of section of Shulgin et al. (2009). We have used the crustal thicknesses and densities of their section which are based on seismic reflection, refraction, tomographic and gravity data. We suggest the position of the subducted ocean– continent boundary is now north of the elevated Savu–Roti Ridge and it was the arrival of the edge of this thickened continental plateau which has driven the young thrusting at the Savu and Roti Thrusts. The crust of the Banda fore-arc and the Australian margin have similar thicknesses and densities because both are continental, and both are ultimately Australian. The Scott Plateau was stretched during Late Jurassic rifting but remained part of the Australian continent. In contrast, the continental crust beneath the Savu Basin was stretched during Late Jurassic rifting but then separated from Australia before accretion to the Sundaland margin in the mid Cretaceous (Hall et al. 2009). It was then stretched again during Middle and Late Miocene subduction rollback into the Banda embayment (Spakman & Hall 2010). Thick Middle Miocene to Pliocene sediments (Unit 2 and lower Unit 3) of the Savu Basin dip and thin northwards into the basin from the elevated

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Fig. 10. Cross section along a north–south transect close to Savu Island (see Fig. 6 for location) sub-parallel to the longest seismic line crossing the Savu Basin which illustrates how the Australian plate and the Banda fore-arc interact in the Savu region. Faults are shown schematically in red. Thicknesses of seismic units are slightly exaggerated for clarity. Crustal densities and deep structure were inferred using information from a transect by Shulgin et al. (2009).

Sumba Ridge, whereas Pliocene to Recent sediments (upper Unit 3 and Unit 4) thicken northwards due to slumping from the Sumba Ridge into the basin. On the north side of the basin Unit 4 probably thickens southward because of sediment carried south from the volcanic arc. There is no significant deformation in the basin. The elevated ridge is an accretionary complex composed partly of Banda fore-arc sediments and partly of Australian margin sedimentary cover. To the south the surface of this wedge dips south and is formed entirely of detached Australian margin sedimentary cover.

Conclusions The Savu Basin records the Miocene to Recent history of convergence between Australia and the SE part of Sundaland. We interpret this region to be underlain by continental crust that was added to the Sundaland margin in the mid Cretaceous. Before the Middle Miocene the region including Sumba and the Savu Basin was close to sea level and subsided rapidly in the late Middle Miocene in response to extension induced by subduction rollback at the Banda Trench as the Java Trench propagated east into the Banda Embayment. The extension is marked by widespread normal faulting. A thick succession of volcaniclastic turbidites was deposited in the basin and was derived from the SW. This material is interpreted as derived from the Sunda Arc to the west of Sumba with flow

being influenced by the relatively shallow Sumba Ridge which caused turbidity currents to flow first SE and then NE into the Savu Basin. The Sumba Ridge is likely to be a feature that reflects the deep structure of the Sunda margin. The western part remained a shallow bathymetric feature during the Neogene although the SE part subsided to depths of more than 4 km. The Sumba Ridge was elevated as continental crust of the Australian margin arrived at the Banda Trench and was flexed into a broad upwarp that tilted the volcaniclastic turbidite sequence and later caused debris flows and turbidites to flow northwards into the basin. Slumps seen on seismic lines came from both Sumba and Timor. Fortuin et al. (1992) noted that slumping affects the area between Sumba and Timor more than areas with similar or steeper slopes such as those offshore Flores where there is also abundant seismicity. This may reflect the combination of tectonic steepening and the layered turbidites and chalk interbeds which detached along bedding surfaces as the collision complex was elevated. Apart from tilting and slumping the Savu Basin is little deformed and there are no thrusts within it. South of Sumba a small part of the original accretionary complex is preserved, now buried beneath almost 1 s of sediment. Most of the precollisional accretionary complex has been incorporated in a large accretionary zone which is south of the Savu and Roti Thrusts and includes both Banda fore-arc material and Australian sedimentary

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cover. This deformed complex is bounded to the north and south by north- and south-vergent thrusts; the former trench is now deep beneath this complex and the lithospheric faults do not emerge at the surface. This deformed zone has developed as the Australian continental crust has been thrust beneath the Banda fore-arc—we suggest the northern edge of the Scott Plateau is now beneath the Sumba Ridge, contributing to the young thrusting north of Savu and Roti. These thrusts reflect the shape of the pre-collisional Australian margin and have caused the islands of Savu and Roti to rise from depths of more than 2 km since 2 Ma and emerge very recently. Underthrusting of the Sumba Ridge by the Scott Plateau is probably contributing to the continued elevation of Sumba and extensional collapse seen on the island. The Timor Trough connects to the Java Trench as a bathymetric feature along the southern zone of south-vergent thrusting but it is entirely a shallow feature entirely within the Australian margin. The trough may have formed at the site of an earlier deeper bathymetric feature within the continental margin but its primary cause is south-directed thrusting and loading by the collisional complex now about 120 km wide south of Savu. We thank the consortium of oil companies who support the SE Asia Research Group, Steve Toothill and CGG Veritas for permission to use the seismic data and for helpful discussion, and Chris Elders, Mike Audley-Charles and Wim Spakman for help and discussion.

References Audley-Charles, M. G. 1986. Rates of Neogene and Quaternary tectonic movements in the Southern Banda Arc based on micropalaeontology. Journal of the Geological Society, London, 143, 161–175. Audley-Charles, M. G. 2004. Ocean trench blocked and obliterated by Banda forearc collision with Australian proximal continental slope. Tectonophysics, 389, 65– 79. Audley-Charles, M. G. 2011. Tectonic post-collision processes in Timor. In: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 235– 260. Bull, S., Cartwright, J. & Huuse, M. 2009. A review of kinematic indicators from mass –transport complexes using 3D seismic data. Marine and Petroleum Geology, 26, 1132–1151. Burollet, P. F. & Salle, C. 1982. Histoire geologique de l’ile de Sumba (Indonesie). Bulletin de la Socie´te´ ge´ologique de France, 24, 573 –580. De Smet, M. E. M., Fortuin, A. R. et al. 1990. Detection of collision-related vertical movements in the Outer Banda Arc (Timor, Indonesia), using micropaleontological data. Journal of Southeast Asian Earth Sciences, 4, 337–356.

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Effendi, A. C. & Apandi, C. 1980. Geological map of Sumba quadrangle, Nusa Tenggara. Geological Research & Development Centre, Ministry of Mines and Energy, Bandung, Indonesia. England, P., Engdahl, R. & Thatcher, W. 2004. Systematic variation in the depths of slabs beneath arc volcanoes. Geophysical Journal International, 156, 377– 408. Fleury, J.-M., Pubelier, M. & Urreiztieta, M. 2009. Structural expression of forearc crust uplift due to subducting asperity. Lithos, 113, 318–330. Fortuin, A. R., Roep, T. B., Sumosusastro, P. A., van Weering, T. C. E. & van der Werff, W. 1992. Slumping and sliding in Miocene and Recent developing arc basins, onshore and offshore Sumba (Indonesia). Marine Geology, 108, 345–363. Fortuin, A. R., Roep, T. B. & Sumosusastro, P. A. 1994. The Neogene sediments of East Sumba, Indonesia – products of a lost arc? Journal of Southeast Asian Earth Sciences, 9, 67– 79. Fortuin, A. R., van der Werff, W. & Wensink, G. 1997. Neogene basin history and paleomagnetism of a rifted and inverted forearc region, on- and offshore Sumba, Eastern Indonesia. Journal of Asian Earth Sciences, 15, 61– 88. GEBCO 2003. General Bathymetric Chart of the Oceans. IHO-UNESCO, Digital Edition 2003, http://www. gebco.net/. Hall, R. 2002. Cenozoic geological and plate tectonic evolution of SE Asia and the SW Pacific: computerbased reconstructions, model and animations. Journal of Asian Earth Sciences, 20, 353–434. Hall, R. 2009. Hydrocarbon basins in SE Asia: understanding why they are there. Petroleum Geoscience, 15, 131 –146. Hall, R. 2011. Australia–SE Asia collision: plate tectonics and crustal flow. In: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia–Asia Collision. Geological Society, London, Special Publications, 355, 73–104. Hall, R., Clements, B. & Smyth, H. R. 2009. Sundaland: Basement character, structure and plate tectonic development. In: Proceedings Indonesian Petroleum Association, 33rd Annual Convention, IPA09-G-134, 1– 27. Harris, R., Vorkink, M. W., Prasetyadi, C., Zobell, E., Roosmawati, N. & Apthorpe, M. 2009. Transition from subduction to arc-continent collision: Geologic and neotectonic evolution of Savu Island, Indonesia. Geosphere, 5, 152– 171. Harris, R. A. 1991. Temporal distribution of strain in the active Banda orogen: a reconciliation of rival hypotheses. Journal of Southeast Asian Earth Sciences, 6, 373– 386. Hendaryono, A. 1998. Etude geologique de l’ile de Flores. PhD thesis, Universite de Savoie, Chambery, France. Pirazzoli, P. A., Radtke, U. et al. 1993. A one million-year-long sequence of marine terraces on Sumba island, Indonesia. Marine Geology, 109, 221– 236. Roep, T. B. & Fortuin, A. R. 1996. A submarine slide scar and channel filled with slide blocks and

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Petroleum Association, 33rd Annual Convention, IPA09-G-013, 1– 12. van der Werff, W., Kusnida, D., Prasetyo, H. & van Weering, T. C. E. 1994. Origin of the Sumba forearc basement. Marine and Petroleum Geology, 11, 363–374. van Weering, T. C. E., Kusnida, D., Tjokrosapoetro, S., Lubis, S. & Kridoharto, P. 1989a. Slumping, sliding and the occurrence of acoustic voids in recent and subrecent sediments of the Savu Forearc Basin, Indonesia. Netherlands Journal of Sea Research, 24, 415–430. van Weering, T. C. E., Kusnida, D., Tjokrosapoetro, S., Lubis, S., Kridoharto, P. & Munadi, S. 1989b. The seismic structure of the Lombok and Savu forearc basins, Indonesia. Netherlands Journal of Sea Research, 24, 251– 262. von der Borch, C. C., Grady, A. E., Hardjoprawiro, S., Prasetyo, H. & Hadiwisastra, S. 1983. Mesozoic and late Tertiary submarine fan sequences and their tectonic significance, Sumba, Indonesia. Sedimentary Geology, 37, 113 –132.

Tectonic post-collision processes in Timor M. G. AUDLEY-CHARLES1,2 1

Department of Earth Sciences, University College London, Gower Street, London WC1E 6BT, UK

2

La Serre, 46800 St. Pantale´on, Montcuq, France (e-mail: [email protected]) Abstract: Indian Ocean crust subducted northwards at the Banda Trench from about 12 to 4 Ma. The Australian continental margin collided with the Asian fore-arc at about 4 Ma. Gradually the Banda Trench was transformed into the fold and thrust mountains of Timor Island. Tectonic collision processes developed when all ocean crust had been subducted and Australian continental crust was refused entry to the subduction path below the Asian fore-arc. The Banda Trench was then gradually converted into a Tectonic Collision Zone (TCZ) progressively filled by two highly deformed Australian continental upper crust mega-sequences. Slowing subduction of Australian sub-crustal lithosphere after c. 2.5 Ma led to uplift of the TCZ that raised Timor 3 km above sea level. Asian Banda fore-arc deformation is linked to c. 30 km southeastwards rollback of the subducting Australian mantle lithosphere. Two Asian fore-arc nappes were thrust southwards from the Banda fore-arc onto the older of two highly deformed Australian continental margin upper crust mega-sequences. The Wetar Suture was created as a thrust at the base of Australian partially detached continental lower crust propagated into the Asian fore-arc. Re-interpretation of BIRPS seismic and gravity data for the Timor region supports this collision model.

This paper discusses the key geological processes associated with the Timor tectonic collision. Timor Island has mountains up to 3 km high, is over 475 km long and has a width of 75 –100 km (150 km wide with its submarine southern slope). Without including the volcanic islands, East and West Timor are together larger than the 280  100 km Swiss Alps. The Banda Arcs are 2300 km long, a length equivalent to the European Alps in France, Italy, Switzerland and Austria as well as the Carpathian Mountains. Timor is not an accretionary wedge of a volcanic fore-arc, nor a fore-arc as some geological maps show, although its rocks include parts of the older volcanic fore-arc. It is comprised mainly of the Australian continental margin deformed by collision with the Banda volcanic forearc. After collision, none of the continental crust was subducted although subduction of Australian sub-crustal mantle lithosphere continued. The Banda subduction trench began to develop between about 15 to 12 Ma from the eastern part of the Sunda –Java Trench (Fig. 1) to form the Banda volcanic arc (Hall 2002). This trench has been assumed by many writers to have been located in what is now the Timor Trough south of the rocks that now form Timor Island (e.g. Hamilton 1979; Rangin et al. 1999), although the Timor Trough is now underlain by about 26 km of Australian continental crust (e.g. Richardson & Blundell 1996; Snyder et al. 1996). Audley-Charles (2004) attempted to explain how the Benioff zone must always have lain north of almost all the rocks that

gave rise to Timor and the other islands of the southern part of the Outer Banda Arc (Fig. 1). Audley-Charles (1986b) and others (e.g. Lorenzo et al. 1998; Tandon et al. 2000; Hall 2002; Londono & Lorenzo 2004; Woodcock 2004) recognized that the Timor Trough is a foreland basin and could never have been a Benioff subduction trench. Audley-Charles (2004) showed how the Banda Trench had been destroyed in the tectonic collision but he did not recognize how the subduction passage below the fore-arc had been partially blocked by the inability of Australian continental crust to subduct from about 4 Ma. Tectonic collision of the Australian continental margin with the Asian Banda fore-arc in Timor can be defined as beginning when the Australian continental crust was first unable to enter the existing subduction passage below the volcanic fore-arc. At 4 Ma the Banda Trench was c. 6 km deep and 30 wide (Audley-Charles 2004). Following collision, rollback of the subducting lithosphere continued by c. 30 km with separation of Australian continental lower crust by delamination from the subducting mantle lithosphere. The former trench became filled by thrust stacking of the Australian continental crust, and Asian nappes that were thrust southwards during rollback, and these rocks now form a region about 35 km in thickness and 110– 150 km in length from NW to SE. There were two important decollements in the Australian continental upper crust; this thickening since 4 Ma was crucial to the orogenic processes. This paper discusses the

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 241–266. DOI: 10.1144/SP355.12 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Plate tectonic reconstruction of part of the Indian Ocean, Banda Trench and NW Australian continental shelf and margin at 10 and 5 Ma based on Hall (2002), and Spakman & Hall (2010). Note position of active Banda Trench north of Timor. Tectonic collision of the Australian continental margin with the Banda volcanic fore-arc proposed by Keep & Haigh (2010) between 10.8 and 9.8 Ma is not tenable.

fundamental roles of rollback, continuing subduction of sub-crustal lithosphere, five major decollements, and the movement of the lower crust creating the Wetar Suture, and how these key processes led to the tectonic evolution of Timor Island.

Summary of geological field studies and regional mapping programmes in Timor Before World War 2 (WW2) geological field studies were focused separately in either Dutch (now Indonesian) Timor, being the western part of the island, or Portuguese (now Timor Leste), being the eastern part of Timor. By far the most important were the studies in West Timor under the direction of Brouwer, at Amsterdam University in 1937. This produced four volumes of papers in 1942 that reported in detail on the geology of many of the key areas of West Timor and some of the adjacent islands of the Banda Volcanic Arc. After the end of WW2 geological investigations in Timor from 1954 to 1969 continued to be conducted separately in West (Indonesian) Timor and in East (Portuguese) Timor. In West Timor de Waard made significant contributions focused on and around the metamorphic massifs of the Mutis Metamorphic Complex and other key stratigraphical targets producing very useful papers: de Waard (1954a, b, 1955, 1956, 1957). Work in East Timor by Escher and Grunau (Grunau 1953, 1956, 1957) and Gageonnet & Lemoine (1958) advanced the

regional stratigraphical and structural geology of East Timor, including the first geological sketch map of all East Timor on a scale of 1:500 000. Van Bemmelen (1949) and Hamilton (1979) in their compendious studies of the geology of the whole of Indonesia produced summaries of the geology of Timor. Lacking plate tectonic ideas before the late 1960s limited the understanding of the tectonics to the geosyncline hypothesis. From 1959 until 2009 UK geologists and their research students based in London University (some of whom continued after completion of their PhD when they moved to other universities, notably Harris and students) worked in both East Timor and West Timor. However, from late 1975 to the 1990s access to East Timor was limited to Bachri and Situmorang only, who were allowed to work there in 1994. Hunter (MSc thesis at West Virginia University 1993) and Reed et al. (1996) had worked in East Timor during the 1990s. Between 1959 and 2009 Audley-Charles, Barber, Carter, Charlton, Giani, Harris, Kenyon and Tobing worked at various times in both East and West Timor. These London-based geologists produced over 60 papers dealing the geology of both parts of Timor, and 14 PhD and MPhil theses were published on the geology of both parts of Timor by the University of London. In 1968 Audley-Charles had published a reconnaissance geological map of Portuguese Timor on a scale of 1:250 000. In 1979 the Geological Survey of Indonesia published a geological map of

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Indonesian Timor also on the scale of 1:250 000. This includes the Kupang quadrangle and the islands of Raijua, Savu and Roti as well as the most western part of west Timor. It is based on the geological mapping of Suwitodirjo & Tjokrosapoetro (1996) and Rosidi et al. (1996). These geological maps of Indonesian Timor employ a very similar stratigraphical scheme and nomenclature to that used in the geological map of Portuguese Timor (Audley-Charles 1968). Together these published maps represent a joint account of the reconnaissance geology of all Timor and of the three adjacent islands of the Outer Banda Arc to the west of Timor. Since about 2003 Haig and Keep and their students, including McCartain and Logan Barber, all based at the University of Western Australia, have carried out geological field work in parts of East Timor, and their papers are referred to below.

Brief geological history of Timor Geologically, Timor and all the islands of the nonvolcanic, Outer Banda Arc are part of the Australian continental margin (Figs 1–3). Their oldest exposed sedimentary rocks are Early Permian in age and associated with Triassic and Jurassic strata. These rocks were deposited in a large Gondwana cratonic basin that underlies much of what is now the continental shelf of northern Australia. The northeastern part of eastern Gondwana was rifted from Australia and New Guinea at about 200 Ma (Pigram & Panggabean 1984). This cratonic basin extended below what are now the islands of the Outer Banda Arc, into that part of Eastern Gondwanaland that rifted from NW Australia at 155 Ma (Audley-Charles 1988; Metcalfe 1988; Powell et al. 1988). Thus, the younger sequence of rocks, of what is now Timor Island, began to be deposited after about 155 Ma on the NW Australian rifted continental margin above the Australian Gondwana cratonic basin. What was to become Timor remained a submarine part of that margin until it emerged as an island in the Late Pliocene following the tectonic collision with the fore-arc of the southern Banda volcanic islands in the Mid-Late Pliocene (Hall 2002). Two mega-sequences can be recognized in the para-autochthon of Timor. The older mega-sequence crops out in the northern three-quarters of the island, whereas the younger is exposed almost entirely in the southern quarter of Timor Island, and it contributes numerous exotic blocks and clay matrix to tectonic melanges exposed with the older megasequence and with the Asian Banda Terrane. The older mega-sequence is the pre-rift or Gondwana mega-sequence. This includes the oldest Australian

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continental margin sedimentary rocks exposed in Timor and all other islands of the Outer Banda Arc, ranging from Early Permian to late Middle Jurassic stratified, mainly sedimentary, rocks. They are now highly deformed and all other rock sequences overlie them. The younger mega-sequence is referred to as the post-155 Ma rift mega-sequence. This ranges in age from Late Jurassic to Pliocene. It is strongly deformed with large-scale recumbent folds, thrusts, with some strong to intense imbrication and pressure solution cleavage, all well exposed in the Kolbano region of SW Timor. The evidence of multiple deformation, and the presence of large-scale flat-lying overthrusts distinguish this mega-sequence of younger rocks from the pre-Late Jurassic rocks of the Gondwana pre-rift megasequence. In a structurally high position on Timor are rocks with an Asian affinity forming thrust sheets called the Banda Terrane (Audley-Charles & Harris 1990; Harris 1991). This rests on the highly deformed para-autochthonous part of the Gondwana mega-sequence of Early Permian to Mid-Jurassic age. The basement rocks of the Banda Terrane are known as the Mutis Metamorphic Complex in West Timor and are correlated with the Lolotoi Metamorphic Complex in East Timor. Banda Trench subduction seems to have been initiated at the easternmost limit of the Sunda – Java Trench and propagated eastwards (Fig. 1) so that it perhaps reached as far east as Seram (Hall & Wilson 2000; Hall 2002; Spakman & Hall 2010). Following collision at about 4 Ma Timor Island was created about 2 Ma by uplift, probably at least partly isostatic, by the Neogene collision between the NW Australian continental margin crust and the fore-arc of the Banda volcanic islands of Wetar, Atauro, Pantar, Lomblen, Adonara and the eastern part of Flores. The collision was associated with the inability of the lower density and thicker Australian upper continental crust and the thinner crystalline lower crust to be subducted with the sub-crustal mantle lithosphere. The sub-crustal mantle lithosphere continued to subduct below the former Banda volcanic arc. Rollback continued southwards after 4 Ma by about 30 km creating a large space that opened southeastwards progressively filled by the stacked upper crust. The age of the collision continues to be a source of controversy. This partly reflects the shape of the pre-collisional Australian margin and the different times at which different parts of this margin came into contact with the Asian margin (Fig. 1; see reconstructions in Hall 2002 and Spakman & Hall 2010). In this paper the collision refers to the time at which the fore-arc of the Banda volcanic arc made first contact with the distal parts of the Australian passive continental margin in the Timor region.

244 M. G. AUDLEY-CHARLES Fig. 2. Summary of Timor autochthonous, para-autochthonous and allochthonous stratigraphy. The allochthonous stratigraphy is found only in the Asian Banda Terrane and its cover rocks, and in the amphibolites of the Aileu Complex. Tectonic collision was N20, although locally N20 strata have been reported in the Australian autochthon. This may indicate diachronous deposition and deformation. See text for discussion.

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Fig. 3. The now invisible Banda Trench became filled with Australian continental crust and two Asian nappes that evolved into the TCZ that became Timor Island and its submarine southern slope. Note filling and overriding of the eastern end of the Java Trench by rocks of the Australian continental margin. The NNE– SSW strike-slip faults that help shape the northern margin of the Australian continent in the Timor region (Audley-Charles 2004, fig. 1) are omitted here as their precise location is unknown.

Dating the Neogene tectonic collision In this paper geological events have been described in terms of their biostratigraphical age, such as Late Pliocene or Early Miocene, with numerical ages based on Gradstein et al. (2005), and/or in terms of the foraminiferal zones of Blow (1969), slightly modified by D. J. Carter for his use in the Banda Arc islands; updated where necessary (Fig. 4) by BouDagher-Fadel (2008). The planktonic foraminifera zonation scheme of Blow (1969) was modified by D. J. Carter in Audley-Charles et al. (1979) and in Audley-Charles (1986a, Fig. 4), and in this paper. The key geological sections for dating the Neogene tectonic-stratigraphical event are in West Timor. One reason for this may be the greater uplift in East Timor, with the greater depth of

exhumation that has removed key sections, and the somewhat greater degree of deformation seen in East Timor.

Micropalaeontological dating in the Kolbano region The Kolbano region of SW Timor (Figs 3–5) has the key exposures by which to date the collision. Here the para-autochthonous sequences from Permian to Early Pliocene reveal that these rocks have been notably indurated. A thin layer of the mixed matrix facies of the Bobonaro Melange has intruded the Neogene section. Above this is the Lower Batu Putih Limestone composed of foraminiferal (Fig. 2) calcilutites, vitric tuffs, and

M. G. AUDLEY-CHARLES

Gl. truncatulinoides Zone Up. Gl. oides quadrilobatus fistulosus Zone

N23

Holocene

N22

Pleistocene and Holocene

N21

Late Pliocene

Lr SphaeoridinellopsisSphaeroidinella Zone G. nepenthesSphaeroidinella Zone

Late N19 (? + N20) Early N19

G. nepenthesSphaeroidinellopsis Zone

N18

Early (and Mid-) Pliocene Early Pliocene Late Miocene Early Pliocene

Favocassidulina favus

Not differentiated on planktonics

Important Benthonic species

Age

Burseolina sp. 2

Planktonic zonation (Blow 1969)

Cassidulinoides infiatus

Local zonation (D. J. Carter)

Hyalinea baltica Bolivinita quadrilatera Hoeglundina elegans Valvulineria jauana Ceratobulimina pacifica

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G., Globigerina; Gl., Globorotalia; Gl. oides, Globigerinoides; Up., Upper; Lr, Lower Fig. 4. Biostratigraphic zonation used by D. J. Carter who used the scheme of Blow (1969) as his framework. Carter modified this scheme for the local zonation as outlined in Audley-Charles (1986a).

A

N

East Timor 50 km

West Timor

B Mt. Cablac Range

Viqueque Group

{ Plio-Pleistocene N21-N23

Banda Sea Volcanics Banda Terrane Gondwana Sequence Post <155 Ma Rift Sequence Aileu Complex Kolbano region

Post-155 Ma sequence thrust over Gondwana sequence

Fig. 5. Principal features of Timor geology simplified after Standley & Harris (2009). Line AB is shown on Figure 8. Banda Sea volcanic rocks are part of the Banda Terrane. The Gondwana mega-sequence shown here includes the Bobonaro Melange, its related broken formation facies, matrix mixed facies and a mixed block in clay facies. These contain some the post-rift mega-sequence blocks and some of its shales and mudstones.

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occasional turbiditic arenites (Kenyon 1974). These were dated by D. J. Carter as zones N18 –N19 (in Audley-Charles et al. 1979) which correspond to an age of 5.6 to 3.8 Ma. Microfossils in the Lower Batu Putih Limestone dated by D. J. Carter (in Kenyon 1974; in Audley-Charles 1986a) from more than 50 samples as zones N18 to N19 were: Globorotalia cultrata, cultrata, Gt. menardii, G. tumida tumida (sinistral), Globigerina nepenthes, Pulleniatina primalis, Globigerinoides gomitulus, Gdes des ruber, Gdes conglobatus, Gdes quadrilobatus sp., Gdes obliquus obliquus, Sphaeroidinellopsis seminulina knocki and Sphaeroidinellopsis subdehiscens paenedehiscens; Benthonics: Textularia spp., Cibicides spp., Spiroplectammmina spp., Uvigerina, spp., Miliolacea, Bulimina spp., Nodosariacea, Planularia spp. and Oridorsalis spp. Overlying these rocks with a disconformity is the notably different, soft and friable Upper Pliocene, Fatu Laob Member of the Upper Batu Putih Limestone that is overlain by the Sabaoe Calcarenite. The Fatu Laob Member was dated by D. J. Carter (in Audley-Charles et al. 1979; in Audley-Charles 1986b) as zone N21, which correlates with about 3.0 to 1.81 Ma, by a microfauna including Globoquadrina dehiscens dehiscens, Sphaeroidinellopsis seminula seminula, Globigerina decoraperta and Pulleniatina obliquilocata praecursor. By using the lower Batu Putih sequence from the Kolbano imbricate zone, that is part of the Australian upper crust (para-autochthon) and the post-rift mega-sequence, below which is the Gondwana mega-sequence, one is sampling the youngest rocks in the strongly deformed part of the orogenic pile. The youngest rocks involved in the deformation indicate the closest we can date the onset of collision, but obviously this date is older than the beginning of the deformation owing to the erosion associated with onset of deformation. The oldest rocks that sit here with a disconformity (the zone N20 being locally absent) on the youngest member of the deformed sequence, namely the upper part of the Batu Putih Limestone, date the rocks deposited after the strong deformation phase. This section, younger than N20, is the oldest part of the Australian autochthonous sequence belonging to the Viqueque Group. Together these two dates, below and above the N20 disconformity, bracket the onset of collision interpreted as the strong deformation in the Tectonic Collision Zone that replaced the Banda Trench, where Australian continental crustal rocks were blocked from subducting below the Banda volcanic fore-arc. It is notable that no N20 zone deposit has been found in the para-autochthonous sequence of the Kolbano region in contrast to its reported presence in the softer and friable autochthonous Viqueque

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Group of the Central Basin of West Timor (de Smet et al. 1990) some 40 km or more to the north. This local absence of N20 zone microfossils and the presence of diagnostic foraminifera for N21 suggest that the strong structural processes that deformed the para-autochthonous sequence in Timor had begun about 4 Ma and ceased about 2.5 Ma in the post-155 Ma rift Australian upper crust autochthonous sequence of the Viqueque Group of Kenyon (1974). The microfauna identified by D. J. Carter (in Kenyon 1974; in Audley-Charles et al. 1979; Audley-Charles 1986a) from the Noele Marl turbidite facies that overlie the Upper Batu Putih Limestone range from Late Pliocene to Recent, N21 to N23 zones, on the basis of Globorotalia truncatulinoides and Gt. tosaensis. Benthonics Hyalina baltica, Bolivinita quadrilera, Hoeglundina elegans are also present. This Viqueque turbidite facies in the Kolbano region therefore ranges from about 3.0 to 0.63 Ma. This correlates with the uplift and subaerial erosion of the emerging Timor Island that appear to have begun by about 2 Ma in some places. The main collision deformation could have been over in the Kolbano region of Timor in about 1 million years during the mid-Pliocene, although some mild folding continued into the Late Pliocene and Pleistocene N22 and N23 in places referred to as the Mataian folding phase (Audley-Charles 1968, 1986a; Kenyon 1974).

Relative dating of the collision by the degree of induration and deformation All the stratified para-autochthonous rocks of Timor, ranging in age from Early Permian to Early Pliocene, are indurated, strongly folded, faulted and thrust, and the younger post-155 Ma rift megasequence rocks are locally imbricated. The sedimentary rocks that belong to the autochthon and range from Middle Pliocene to Pleistocene are not indurated. These folded post-Middle Pliocene– Pleistocene (N22 –23) rocks lack strong deformation. These differences clearly distinguish those that were deposited before the Neogene onset of the c. N20 tectonic collision.

Dating the Bobonaro Melange: pre-decollement activation The Sonnebaitserie of Tappenbeck (1940) and other earlier workers in West Timor was mapped in East Timor by Audley-Charles (1965, 1968) as the Bobonaro Scaly Clay, regarded as an olistostrome emplaced in the Middle Miocene on the basis of relative stratigraphic relationships. Later Harris et al. (1998) renamed this the Bobonaro Melange.

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They also noted three end member subdivisions of the Bobonaro Melange which are broken formation facies, matrix mixed facies, and a mixed blockin-clay facies. One can now consider that the Bobonaro Scaly Clay exposures without exotic blocks are matrix mixed facies. Locally this facies, widely exposed in Timor, contains a N22 Pleistocene foraminiferal assemblage identified by Carter as Globorotalia truncatulinoides, Pulleniatina obliquilocata, Sphaeroidinella dehiscens and ‘Globigerina’ subcretacea (Audley-Charles 1986a). This indicates that the Bobonaro Melange matrix mixed facies with Pleistocene foraminifera derived from the soft strata of the Viqueque Group exposed upslope, is locally a younger reworked version of the Bobonaro Melange and is often associated with landslips. In an unpublished report D. J. Carter (1970, in Audley-Charles 1986a) noted that, despite the intensive study of hundreds of samples, no species of Oligocene or Miocene age (except those in exotic fragments of allochthonous Cablac Limestone) has been found in the Bobonaro Melange. However, Middle Eocene planktonic assemblages are abundant: Globorotalia centralis, Globigerina ampliapertura, Globigerina parva, Globigerina venezuelana, Globigerinatheka barri, Catapsydrax echinatus, Hantkenina dumblei, Hantkenina alabamensis, ‘Globigerapsis index’, Globorotaloides suteri, Truncorotaloides topilensis, Hastigerina micra, Globigerina yeguaensis and Chiloguembelina martini. It is likely that the Bobonaro Melange, that is mud-dominated and in many places in a scaly clay facies, as well as the broken formation facies would have been generated early in the collision process where there was some chaotic crushing and shearing of stratigraphic sequences that were trapped in the vice between the volcanic fore-arc basement and the top of the Australian crystalline lower crust. This seems likely to have preceded the activation of the two important Australian upper crust mega-sequence decollements, D1 moving the post-155 Ma rift and the D2 decollement moving the pre-rift Australian mega-sequence (Figs 2 & 5). These Australian continental upper crust rocks were crushed and sheared so that the Jurassic clay shales and the Eocene clay sequences make a notable contribution to the Bobonaro Melange and to the broken formation. The age of the Bobonaro Melange is not easy to determine. Its age of emplacement is not known precisely. D. J. Carter (pers. comm. 1980) commented that the microfauna found in the Bobonaro Melange does not assist us in this matter, as all taxa are derived and reworked and range in age locally from Permian to Quaternary (Audley-Charles 1968; Giani 1971; Kenyon 1974; Hamilton 1979).

Carter et al. (1976) found the youngest Kolbano strata underlying the Bobonaro Melange were N17 and that N18 faunas of 5 Ma age are found in the overlying Viqueque Group. This suggests that the Bobonaro Melange is not older than 5 Ma. Standley & Harris (2009) and Harris et al. (1998) pointed out that the Bobonaro Melange is closely associated with the allochthonous Banda Terrane, and they noted that it is present below the base of the Lolotoi and Mutis Metamorphic Complexes that form the base of this terrane. It seems likely that the Bobonaro Melange may have acted as, or with, the decollement on which the terrane moved, thus facilitating the terrane’s transfer upwards and southwards from the Banda fore-arc before its emplacement on top of the Australian paraautochthon. Harris et al. (1998) interpreted the Bobonaro Melange as ‘having formed along the suture between the Asian-affinity fore-arc thrust sheets above and thrust duplex of Australian continental margin units below’. They also interpreted the maximum age of the injection of Bobonaro Melange into the Kolbano sequence with its Viqueque Group cover to be 5 Ma.

Offshore bathyal continental terrace Haig & McCartain (2007) suggested that ‘a middle bathyal continental terrace setting continued in this region, at least on the southern side of Timor, from Cretaceous –Palaeogene to the Late Miocene and Early Pliocene’. This means that from 140 to 4.5 Ma the Timor region, or at least its southern part, was in a medium to deep water setting. This makes a strong contrast with the shallow marine cover rocks, such as the Eocene Dartollu Limestone thrust onto, and the Upper Oligocene to Lower Miocene Cablac Limestone found unconformable upon, the metamorphic basement of the Banda Terrane (Audley-Charles 1968; Carter et al. 1976).

Volcanic activity When continental margins collide with volcanic fore-arcs the volcanoes cease their activity. To the north of Timor is the longest and most obvious extinct sector of the entire Sunda– Banda volcanic arc (Wheller et al. 1987) in the islands of Alor, Atauro, Wetar and Romang. All these volcanic islands are less than 50 km from the Australian continental margin that has overthrust the Banda fore-arc, and Atauro is only 25 km from the Australian continental margin. The four islands are no longer volcanically active and activity ceased at about 3 Ma (Abbott & Chamalaun 1981). Scotney et al. (2005) studied massive sulphide and baritegold mineralization of Wetar Island triggered by Banda Arc –Australian continental margin collision.

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They obtained Ar –Ar ages of 4.7 + 0.16 Ma from biotite and 4.93 + 0.21 Ma from illite in altered footwall volcanic rocks associated with mineral deposits that formed on the flanks of a volcanic edifice at water depths of c. 2 km. These are covered by post-mineralization cherts, gypsum, Globigerina-bearing limestone, lahars, subaqueous debris flows and pyroclastic rocks. The youngest dated volcanic rock is a dacite of 2.39 + 0.14 Ma. Scotney et al. (2005) interpret the volcanic stratigraphy and dating to indicate that the Wetar edifice formed at around 12 Ma by extensive rifting and associated volcanism within oceanic crust. Bimodal volcanism and a basement of basalts and basaltic andesite most likely formed around 5 Ma, with mineralization between 4.9 and 4.7 Ma. Postmineralization dacite flows indicates that volcanism continued until at least 2.4 Ma. A similar history is recorded on Sumba where breakup of a carbonate platform began in the early Middle Miocene leading to rapid subsidence from about 16 Ma (Fortuin et al. 1994). There was a significant increase in volcanic activity in the Middle Miocene, and a change to volcaniclastic turbidites deposited on Sumba and in the Savu Basin (Fortuin et al. 1992, 1994) from NN5 (14.8– 13.5 Ma) to NN11 (8.3–5.5 Ma). The very young age of collision west of Timor is indicated by the still active, or only recently inactive, volcanic arc in western Flores (Wheller et al. 1987), ongoing deformation at the margins of the Savu Basin (Rigg & Hall 2011), and recent emergence of the islands of Roti and Savu (Roosmawati & Harris 2009).

Uplift and erosion The obvious has too often not been mentioned in the debate about the age of Timor collision, and one important issue is the timing of uplift and erosion expected as the Australian margin collided with the volcanic arc. In East Timor uplift seems to have began at about 4 Ma (Audley-Charles 1986a) with estimated average rates of uplift of 1.5 km/ Ma. For West Timor de Smet et al. (1990) found that significant uplift began at about 2.2 Ma, and that emergence from water depths of 1 km to elevations of 0.5 km above sea level occurred in the last 0.2 Ma at average rates of 0.75 to 1.0 km/ Ma. Haig & McCartain (2007) state that proximal turbidite deposition from the rising island of Timor began at about 3.35 Ma. Quaternary limestones are mapped in West Timor at elevations above 1 km (Suwitodirjo & Tjokrosapoetro 1996). Roosmawati & Harris (2009) show that for the islands of Savu and Roti significant and rapid uplift from water depths of more than 2 km to emergence began at about 2 Ma. On Sumba there was

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uplift from depths of more than 5 km to emergence of up to 1 km above sea level since 4 Ma (Fortuin et al. 1997). The uplift of Timor over the last 2 Ma was at least 4 km in places (Audley-Charles 1986a). Furthermore, the production of clastic sediment from Timor over the last c. 2 Ma has built the sedimentary slope, aided by the young D5 decollement, that extends from the 3 km deep axis of the Timor Trough to many metres above sea level along the south coast of Timor for about 480 km length and over a width of 50 km.

Orogeny-engendered stratigraphic diachronism Some of the Pliocene para-autochthonous and autochthonous strata of Timor and other islands of the Outer Banda Arc are diachronous. Harris et al. (2009) and Roosmawati & Harris (2009) have noted that in East Timor the Batu Putih Formation (lower or upper parts not indicated) is overlain by N19/N20 Noele Marl clastic deposits (Kenyon 1974). In the Central Basin of West Timor the upper sections of Batu Putih Formation (lower or upper parts not indicated) are dated as N20 by de Smet (1990). In Roti (immediately SW of Timor) the youngest Upper Batu Putih Formation is late N23 (Roosmawati & Harris 2009); further west in Savu Island the youngest Upper Batu Putih Formation is N22/23. This represents an age range of 4 to 0.5 Ma for the Upper Batu Putih Formation. Diachronism is further illustrated by the significantly different results obtained for Neogene – Quaternary vertical movements in the Central Basin of Timor (de Smet et al. 1990) and those obtained in West Timor outside the Central Basin by AudleyCharles (1986a). The Central Basin of West Timor, that was at times subsiding and at other times rising, is located between the Kolbano younger megasequence imbricate zone in the south and the Banda Terrane in the north (Audley-Charles 1986a). In the Batu Putih Limestone (lower or upper part not indicated) of the Central Basin Viqueque Group Australian autochthon de Smet et al. (1990) found N20 foraminifera, but c. 40 km south, located above the Kolbano imbricate succession of the post-rift mega-sequence, underlain by the Gondwana megasequence, Carter (in Audley-Charles 1986a, b) reported that the N20 zone was absent below the N21 Upper Batu Putih Limestone. This orogeny-related diachronism could be the result of several factors: the irregular shape of the Australian continental margin, different times during the collision at which different parts of the post-155 Ma rift mega-sequence were deformed in the Banda Trench, and continual stacking of the deforming Australian continental margin

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mega-sequences in the TCZ; equally there is also a range in age of uplift in different parts of the 480 km long, 75 –100 km wide Timor Island.

Discussion of collision ages In the Banda region, notably in Timor, interpretation of ages has led to confusion and also claims of multiple or pre-Pliocene collisions. As pointed out above, some of the variation in ages may be the consequence of diachroneity reflecting the irregular shape of the Australian margin, or an oblique collision (e.g. Harris 1991). Nonetheless, all the ages are consistent with a collision beginning at about 4 Ma or slightly later. Deep-water sedimentation continued into the Pliocene with a succession of Plio-Pleistocene Viqueque Group friable chalks deposited in lower to middle bathyal water depths with a quiet tectonic interval. Proximal turbidite sedimentation in southern Timor was dated by Carter (in Audley-Charles et al. 1979) at 3 Ma, and by Haig & McCartain (2007) at 3.35 Ma. However, some authors have followed Berry & Grady (1981) and Berry & McDougall (1986) in interpreting high grade metamorphosed rocks with cooling ages of 8 Ma as marking initial collision of Australian crust with the Asian margin. Based on deformation, principally fault reactivation, of NW Shelf sequences in the Timor Sea south of the Timor Trough (i.e. more than 100 km south of the southernmost deformation front located at the axis of the Timor Trough, Fig. 3), an 8 Ma collision age has been suggested in the Sumba region by Keep et al. (2002, 2003). An even older age has been suggested by Keep & Haig (2010) for Timor, where they proposed that tectonic collision began between 10.8 to 9.8 Ma. Pre-Pliocene collision ages are very improbable in the light of the evidence summarized above, and ages of initial collision of about 10 Ma are untenable. Much of the Banda volcanic arc is significantly younger than 10 Ma. North of Timor the age of initiation of volcanic activity is uncertain but there are few volcanic rocks anywhere in the arc with ages greater than 12 Ma (Abbott & Chamalaun 1981; Macpherson & Hall 2002). The oldest volcanic rocks on Flores, NW of Timor are 16 Ma (Fleury et al. 2009), based on unpublished K – Ar ages from Hendaryono (1998), and volcanic rocks there fall into two age groups, with ages from 16 to 8.4 Ma and 6.7 to 1.2 Ma. On Wetar volcanic activity began at about 12 Ma and continued until at least 2.4 Ma (Scotney et al. 2005). In the South Banda Sea back-arc spreading occurred between 6.5 and 3.5 Ma (Hinschberger et al. 2001). A continuation of volcanic arc activity for a period after initial contact between the fore-arc and Australian margin is plausible, but advocates of a Miocene 10

to 8 Ma collision fail to explain why arc volcanic activity continued until after 3 Ma in Flores and Wetar. In fact, Keep & Haig (2010) make little attempt to explain events outside Timor and show only an unscaled diagram which displays elongate promontories projecting northwards in the regions of East Timor and Sumba. Even a brief consideration of rates of subduction show that these narrow promontories would have to extend several hundred kilometres northwards to cause collision at 10 or 8 Ma. Reconstructions of the Banda region (Hall 2002; Spakman & Hall 2011) suggest that about 10 Ma the Banda Trench was located c. 600 km north of its position at 4 Ma. There is no evidence for the Australian continental margin having extended significantly north of what is now Timor, even allowing for the overthrusting of the Banda volcanic fore-arc and shortening of the Australian margin, although there probably were offsets in the continental margin (Harris 1991; Hall 2002). All the evidence discussed above rules out (Fig. 1) a tectonic collision between the continental margin of Australia and the Banda volcanic fore-arc at about 10.8 to 9.8 Ma, as claimed by Keep & Haig (2010). Other explanations of the 8 Ma metamorphic event are discussed below.

Composition and origin of the Banda Terrane basement Metamorphic rocks and sedimentary cover sequence with an Asian affinity, derived from the fore-arc of the Banda Volcanic Arc, were emplaced on what was to become Timor during the Neogene, in structurally high positions as thrust sheets called the Banda Terrane (Audley-Charles & Harris 1990; Harris 1991). They are found sitting on the highly deformed para-autochthonous part of the Australian Gondwana mega-sequence of Early Permian to mid-Jurassic age (Figs 2 & 5). The basement rocks of the Banda Terrane are known as the Mutis Metamorphic Complex in West Timor (Earle 1981; Brown & Earle 1983) and are correlated with the Lolotoi Metamorphic Complex in East Timor (Audley-Charles 1968; Barber & Audley-Charles 1976; Standley & Harris 2009).

Allochthonous Asian status of metamorphic complexes The truly allochthonous nature of the Mutis and Lolotoi Metamorphic Complexes has now been further confirmed by Standley & Harris (2009) in their petrological study of the Lolotoi Metamorphic Complex. They presented data indicating that one sample from the Lolotoi has ‘a Late Cretaceous

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sedimentary protolith deposited after 82 Ma’, and the ‘Lolotoi Complex has a metamorphic age of 46 Ma. It has cooling ages of 39– 35 Ma. The detrital zircon grains in the Lolotoi are as young as 82 Ma within the core of the Bebe Susu Nappe of East Timor’. The Eocene metamorphic age for the Lolotoi and Mutis Complexes are incompatible with these complexes having been a part of the Australian continental basement of Timor Island, as suggested by some authors (Grady 1975; Chamalaun & Grady 1978; Charlton 2002). Standley & Harris (2009) also reported that ‘the Lolotoi Complex records all the deformation events reported from the Mutis Complex’. The postulated correlation of the Alieu with Lolotoi metamorphics (Kaneko et al. 2007) is also incompatible with the detailed work of Standley & Harris (2009), for example, no zircon grains younger than Permian are found in the Aileu Complex. Apatite grains in the Aileu are highly annealed yielding exhumation fission track ages of ,2 Ma whereas in the Lolotoi Complex apatite grains have model ages of ,23 Ma. The Lolotoi and Mutis Complexes are also overlain by a Late Oligocene unconformity.

Significance of Aileu Formation metamorphism for the age of tectonic collision Keep & Haig (2010), citing the work of Berry & McDougall (1986) on the 8 Ma cooling ages of the Aileu Formation, claim that this implies ‘that the earliest collision of some part of the Australian margin occurred prior to this time, and that the rocks were subducted, metamorphosed and exhumed by 8 Ma’. Another view (Standley & Harris 2009) is that the metamorphism of the Aileu with cooling ages of 8 to 4 Ma indicate these rocks were present at c. 8 Ma about 35 km below the Banda Trench. That means they were located in the fore-arc basement of the Banda Volcanic Arc before c. 8 Ma and would have been already metamorphosed at that depth. Keep & Haig’s (2010) suggestion of an 8 Ma collision conflicts with the evidence presented above, and it requires, as does the Standley & Harris (2009) model, that the Australian continental crust had been subducted to about 35 km depth below the Banda fore-arc where it was metamorphosed to amphibolite grade between c. 8 to 4 Ma. Tectonic collision is defined here as contemporaneous with the blocking of Australian continental crust in the trench from subducting below the fore-arc. There is no evidence that any Australian continental crust was subducted in the Banda Trench system, quite the contrary, evidence suggests that much of the upper and lower crust were rejected by the

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subduction system. The upper crust was stacked to fill the trench, while the crystalline lower crust seems to have been thrust northwards from the base of the Banda Trench over the downwardflexing, northward-dipping, subducting mantle lithospheric slab, thus blocking the subduction system against any continental crust (Figs 6–8). The suggestion of an 8 Ma tectonic collision anywhere in the southern part of the Banda Arc finds no foundation in field observations or laboratory science. Another way of thinking about the 8 Ma cooling ages and the first and second deformation phases of the Aileu amphibolites in the narrow, Timor north coast strip and adjacent small islands seems to fit all the evidence. These cooling ages and deformation can be regarded as the product of metamorphism of a part of the fore-arc basement associated with thermal conditions in the fore-arc adjacent to the active volcanism in Alor, Atauro and Wetar Islands. Harris & Long (2000) reported that fission tracks in apatite grains in the Aileu Formation are, highly to completely annealed, indicating an exhumation age of ,2 Ma. This Late Pliocene exhumation age would correspond with upward movement of the Aileu amphibolites facilitated by the extension in the Banda fore-arc, associated with its subsidence into the trench, caused by rollback of the subducting sub-crustal lithospheric slab, and result from the upward movements along the Wetar Suture (Fig. 8). These Aileu amphibolites were taken from the fore-arc basement by the moving Wetar Suture and lower crust and carried upwards where they were moved together with the Aileu glossy slates and phyllites (Fig. 8). These slates and phyllites could have formed deep in the TCZ but some distance above the lower crust. From there the decollement, at the base of the Gondwana mega-sequence, moving upwards associated with the lower crust and Wetar Suture, late in the collision process, could have delivered these rocks to where we find them today (Fig. 8). ‘The structure of the Aileu is characterized by a very distinct layer and foliation parallel shortening that forms a layer normal cleavage and foliation’ (Standley & Harris 2009). This is different from the structure of the rest of the Gondwana mega-sequence, and may have resulted from it having occupied a deeper position in the TCZ before its transfer upwards to its present position (Fig. 8). Another explanation for Australian continental basement in the Banda fore-arc is offered by Hall (2002, 2011) and Spakman & Hall (2010). The Banda Terrane in Timor appears to be a complex that includes continental crust, arc and fore-arc that formed part of the Early Cenozoic Asian margin together with their overlying sedimentary rocks. The Banda fore-arc also included Australian

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Fig. 6. Cartoon cross section of 5 Ma pre-collision. Note no Australian crust was subducted with sub-crustal mantle lithosphere. When X2 arrived at X1 (4 Ma) all ocean crust had been subducted and the Australian continental crust was refused entry to the subduction passage. This tectonic collision of the Asian fore-arc with the Australian continental margin began at the Banda Trench which was to gradually become the Tectonic Collision Zone shown in Figure 8.

continental crust from the Sula Spur that collided in the Early Miocene with Asian Neogene arc and was then extended during the Late Miocene during rollback (Spakman & Hall 2010). The first contact of the Australian and Asian margins was collision of the Australian Sula Spur soon after 25 Ma (Spakman & Hall 2010) and recorded by metamorphic rocks in Asian Sulawesi (e.g. Parkinson 1998a, b; van Leeuwen et al. 2007) and dredged samples from the Banda Ridges (Silver et al.

1985). A 24 Ma age is recorded in Timor (Berry & McDougall 1986) although this has not previously been explained. During the Neogene Australianorigin crust added to the Asian margin by collision was extended by rollback in the Banda Embayment (Spakman & Hall 2010). Extension was discontinuous, indicated amongst other evidence by oceanic spreading in the North Banda Sea from 12.5 to 7.2 Ma and South Banda Sea from 6.5 to 3.5 Ma (Hinschberger et al. 2000, 2001), and an episode

Fig. 7. Cartoon cross section of pre-collision 5 Ma repeating the configuration of Figure 6. The pecked line shows bottom of the Banda Trench profile after the X1 –X2 tectonic collision. It shows that there has been no subduction of continental crust. It also shows the subsequent rollback and subduction of sub-crustal mantle lithosphere. Note subsidence of Banda Trench-TCZ after 30 km of southeastwards rollback from 4 to 0 Ma, and accompanying northwards subduction of sub-crustal mantle lithosphere. Note also subsidence of fore-arc (F.S.) linked with the rollback.

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Fig. 8. Cartoon cross section of Timor today, (cf. Richardson & Blundell 1996, their BIRPS figs 3b, 4b & 7; and their fig. 6 gravity model 2 after Woodside et al. 1989; and Snyder et al. 1996 their fig. 6a). Dimensions of the filled 40 km deep present-day Timor Tectonic Collision Zone are based on BIRPS seismic, earthquake seismicity and gravity data all re-interpreted here from Richardson & Blundell (1996) and from Snyder et al. (1996). NB. The Bobonaro Melange, its broken formation and other facies are not indicated, but they are included with the Gondwana mega-sequence. Note defunct Banda Trench, now the Timor TCZ, filled with Australian continental crust and Asian nappes that occupy all space between Wetar Suture and the 2 –3 km deep deformation front north of the axis of the Timor Trough. Note the much younger decollement D5 used exactly the same part of the Jurassic lithology of the Gondwana mega-sequence in the older D1 decollement that produced what appears to be much stronger deformation.

of arc volcanism from 8 to 3 Ma is recorded in the North Banda region (e.g. Honthaas et al. 1998) and the Banda Ridges (Silver et al. 1985). All these rocks record metamorphic cooling ages associated with the Neogene extension. Similar complex Australian crust is known from several of the small islands east of Timor from Leti to Babar, and Bowin et al. (1980) commented on the anomalous pre-collision position of these rocks within the Banda Outer Arc. Hall (2011) suggests the assumption that metamorphic ages mark contractional deformation that accompanied an early collision is wrong – the ages simply record cooling, which in most cases resulted from extension. The ages do not record the time of collision at the place the

rocks are now found, because they have been moved to their present positions by extension of the upper plate above the retreating subduction hinge, and the Neogene metamorphic ages record extension of this complex upper plate (Spakman & Hall 2010; Hall 2011).

Asian cover rock sequences of the Banda Terrane The Banda Terrane (Fig. 5) is now found occupying high ground, usually as gently folded thrust sheets, with much post-emplacement steep faulting. The metamorphic rocks are covered in places by

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several different, mainly sedimentary sequences, and some basic volcanic and volcaniclastic rocks. Tappenbeck (1940) explained the stratigraphy of some of these rocks overlying the Mutis Metamorphic Complex in terms of three distinct sequences, all of which have been found in the equivalent stratigraphical and structural position in East Timor (Audley-Charles & Carter 1972; Carter et al. 1976; Standley & Harris 2009). What makes these stratigraphic sequences, which are unconformable or thrust onto the Mutis and Lolotoi Metamorphic Complexes, so important is that some of these supra-metamorphic sequences are of Asian affinity, and the others seem likely to be Gondwana mega-sequence thrust slices emplaced after the Banda Terrane arrived in Timor (Fig. 2). Here the key features of the cover rock sequences are first summarized, and then some controversies concerning their status and significance are discussed.

found with angular unconformity on the Cretaceous Noni Formation in West Timor (see Tappenbeck’s map). This Mesozoic –Paleocene sequence is discussed by Tappenbeck (1940), Audley-Charles & Carter (1972), Audley-Charles et al. (1974), Rosidi et al. (1996), and Earle (1983).

Middle and Upper Eocene The Eocene Dartollu Limestone (Audley-Charles & Carter 1972; Carter et al. 1976) is found resting with a faulted base on the Lolotoi and Mutis Metamorphic Complexes in both West and East Timor. It is possible that these Eocene rocks are locally unconformable on the metamorphic complexes. One importance of this limestone is the palaeontological evidence of its Asian origin (Lunt 2003), as this supports the allochthonous status of the metamorphic complexes of the Banda Terrane (see Tappenbeck’s map).

Triassic sequences

Oligocene

Mapping on a 1:25 000 scale by Tappenbeck (1940) shows exotic blocks of Upper Triassic Bahamian limestones, mainly on the slopes of Mt Mollo in West Timor, associated with exotic blocks of Permian crinoidal limestones, as part of the Sonnebaitserie (which today would be referred to as Bobonaro Melange). His map strongly suggests that the Bobonaro Melange with various exotic blocks once covered the whole of Mt Mollo which is a massif of the Mutis Metamorphic Complex. Keep et al. (2009) reported similar Bahamian-type limestones of Upper Triassic age in a thrust stack high on Mt Cablac. They might be Australian, or perhaps Asian and truly allochthonous.

The Oligocene Barrique Volcanic Formation that is found widely in East Timor where it is locally found thrust on the Lolotoi Metamorphic Complex. The Barrique Formation is overlain unconformably by the Cablac Limestone of Late Oligocene to Early Miocene age (Audley-Charles 1968; Standley & Harris 2009). These Oligocene volcanic rocks have been shown to be typical of volcanic arcs and are incompatible with Australian continental margin igneous rocks (Standley & Harris 2009).

Upper Jurassic and Cretaceous Palelo Group sequences The lowest sequence of the Palelo Group rocks, referred to as the Metan Formation, is comprised of mainly agglomerates and tuffs. It is thought to be of Late Jurassic age because it is found below Lower Cretaceous cherts, dated as Aptian – Turonian with radiolaria (Haile et al. 1979) of the Cretaceous Noni Formation (Earle 1983). These rocks are commonly in fault contact with the Mutis and Lolotoi Metamorphic Complexes and are often associated with the Bobonaro Melange. This is shown by Tappenbeck (1940) on his 1:25 000 map of Mt Mollo.

Upper Oligocene – Lower Miocene The Upper Oligocene–Lower Miocene Cablac Limestone sits with eroded unconformity on the Mutis Metamorphic Complex of West Timor, in places with a cobble–pebble basal conglomerate, (Tappenbeck 1940, 1:50 000 map of Mt Booi) in West Timor, and is also present on Mt Mata Bia in East Timor (Standley & Harris 2009). Its presence on the lower slopes of Mt Cablac was reported by Audley-Charles (1968) and Carter et al. (1976). Standley & Harris (2009) also reported it sitting unconformably on the Lolotoi Metamorphic Complex on Mt Cablac. In contrast, Haig et al. (2008) claim that ‘No Upper Oligocene to Lower Miocene stratigraphic unit has been found on Cablac Mountain’. This contradiction is discussed below.

Pliocene Manamas Volcanic rocks Paleocene Palelo Group sequences A Paleocene sequence of volcanic rocks and volcaniclastics, known as the Haulasi Formation, is

Standley & Harris (2009) reported that these typical arc volcanic rocks are the youngest members of the Banda Terrane found locally on the north coast

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of Timor, having been derived from the basement of the Banda volcanic fore-arc. This discovery provides further support for Pliocene tectonic collision and late emplacement of this allochthonous terrane.

Stratigraphic and tectonic significance of the Mt Cablac Range, East Timor The Mt Cablac Range (Fig. 5) is 15 km long. Its base is a thrust sheet of the Lolotoi Metamorphic Complex that functions as the base of the Banda Terrane throughout East and West Timor (Fig. 5) where it is usually overlain unconformably by the Cablac Limestone.

Mesozoic sequences Triassic limestones on Mt Cablac, which occur in a Bahamian facies on the upper slopes, were not mapped by Audley-Charles (1968) because Mt Cablac was then considered sacred by local people and investigation of the upper slopes in 1959– 1962 was prohibited. Upper Triassic –Lower Jurassic limestones are now known from the higher levels of Mt Cablac, thanks to the more recent investigations reported by Haig & McCartain (2007), Haig et al. (2008) and Keep et al. (2009). Keep et al. (2009) published a geological sketch map at a scale of 1:100 000 concerned with the geology of part of the Mt Cablac Range. This map is described by Keep et al. (2009, p. 151) as ‘detailed geological mapping’, whereas on page 155 they describe it in their figure 5 caption as a ‘generalized geological map of the Cablac Mountain Range’. It includes some discoveries, including an Upper Triassic Bahamian limestone, that appears to be similar to that mapped on Mt Mollo in West Timor by Tappenbeck (1940). The variety of lithology and age of sedimentary rocks and the igneous cover rocks reported by Haig & McCartain (2007), Haig et al. (2008) and Keep et al. (2009) on the Mt Cablac Range are all very similar to rocks reported from diverse Lolotoi and Mutis metamorphic massifs (where the unconformable Upper Oligocene to Lower Miocene Cablac Limestones are also found) in East and West Timor. The Mesozoic and younger limestones reported by Haig et al. (2008) on the higher slopes of Mt Cablac appear likely to be either: (i) the remnants of a more extensive overthrusting of the Mesozoic and other rocks onto the upper slopes of Mt Cablac in a post-2.5 Ma event, or (ii) as on Mts Mollo, Booi, Mata Bia, and many other massifs in East and West Timor, exotic blocks that are part of the Bobonaro Melange. Earlier, Carter et al.

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(1976), had mapped Triassic limestones and Jurassic lithologies belonging to the para-autochthonous pre-155 Ma Gondwana rift sequence overthrust onto the Oligocene– Lower Miocene Cablac Limestone on the lower slopes of Mt Cablac.

Upper Oligocene – Lower Miocene Cablac Limestones Audley-Charles (1968), as part of his mapping of East Timor at a scale of 1:250 000, described the variety of limestones he found cropping out on the lower slopes of the Mt Cablac Range as hard, massive calcilutites and calcarenites. Samples of these rock types were dated by D. J. Belford (of the Australian BMR) in 1960, reported in AudleyCharles (1968) as Te, Early Miocene. In 1963 D. J. Carter (reported in Audley-Charles 1968) studied samples from these lower slopes collected by Audley-Charles and confirmed their Te status and Early Miocene age. Later, Carter et al. (1976) visited the same lower slopes of Mt Cablac Range, which Carter sampled and confirmed their age as N3 to N8 (zonation of Blow 1969) which is Late Oligocene to Early Miocene. What Audley-Charles (1967) identified on Mt Cablac as a brecciated, polymict intracalcirudite belonging to Upper Oligocene –Lower Miocene Cablac Limestone has since been interpreted by Haig et al. (2008) as a Pleistocene crush breccia in a local fault zone. On the basis of that revision Haig et al. (2008, p. 10) stated ‘No Upper Oligocene to Lower Miocene shallow marine stratigraphic unit has been found on Cablac Mountain’. This claim that no Upper Oligocene to Lower Miocene shallow marine limestones are present unconformably on the Lolotoi Metamorphic Complex contradicts the micropalaeontological and petrological analyses by D. J. Belford in 1960, and those by D. J. Carter in 1963 reported by Audley-Charles (1968), and Carter et al. (1976) when Carter re-confirmed his earlier finding. The claims by Haig et al. and Keep & Haig (2010) cited above, appear to imply that two very experienced micropalaeontologists, Belford and Carter, must have mistaken Triassic or other Mesozoic foraminifera for all the Late Oligocene and Early Miocene foraminifera in every limestone sample they reported on from Mt Cablac. This seems extremely improbable (Marcelle K. BouDagherFadel, pers. comm. 2009). Moreover, more recently, Standley & Harris (2009, p. 85) also reported finding Oligocene to Miocene Cablac Limestones, very similar to those reported by Audley-Charles (1968) and by Carter et al. (1976), on the lower slopes of the Mt Cablac Range where they can be seen sitting unconformably on the Lolotoi Metamorphic Complex.

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Tectonic processes associated with the Australian para-autochthon and Asian allochthon The collision was associated with the rejection of Australian continental crust by the Timor subduction system when it met the Banda volcanic fore-arc at the Banda Trench. Maybe the low density and the increasing thickness in the proximal slope of the submarine terrace or plateau led to its rejection. In contrast, Australian sub-crustal mantle lithosphere appears to have subducted normally. AudleyCharles (1988, fig. 14), Haig et al. (2008) and Roosmawati & Harris (2009) considered this Australian upper crust sequence similar to the Exmouth, Scott and related plateaux. The great differences in structural style between the younger post-155 Ma rift mega-sequence and the older pre-rift mega-sequence can be largely attributed to the significant differences in the sedimentary lithologies that characterize the different Australian mega-sequences. It is not known whether two megasequences, each with their own principal decollement, were activated simultaneously or not.

Tectonics of the Australian continental upper crust para-autochthon and its two mega-sequences The mechanical solution, that nature exploited, while the subcrustal part of the lithosphere continued subducting, was to provoke the continental upper crust, of mainly stratified sedimentary rocks, that remained in the trench to activate decollements at the base of the two mega-sequences: namely D1 at the base of post-rift mega-sequence ranging in age from Late Jurassic to Neogene, and D2 at the base of the Gondwana mega-sequence ranging from Lower Permian to Middle Jurassic (Figs 2 & 8). The two mega-sequences remained attached to each other south of the subducting slab, where the lower mega-sequence remained attached to the Australian lower crust and thus to the continental mantle lithosphere. This caused the Australian upper crust to be pulled northwards into the trench where the compressive and related stresses must have increased significantly, as the Australian crust began to be packed ever more tightly. Eventually the two principal decollements were activated. This allowed the movement of the two mega-sequences in a coherent and discrete manner. In this way, each of them developed its own very different deformation style related to the very different lithologies comprising the two mega-sequences. Another important consequence of the activation of these two decollements was that, after the initial chaotic crushing phase, each mega-sequence then

seems to have been deformed separately in the collision complex. The initial chaotic phase seems likely to have been responsible for the origin of the Bobonaro Melange and its facies.

Australian continental margin upper crust sub-late Jurassic decollement D1 Once the Middle Jurassic part of the Wai Luli Formation (Audley-Charles 1968), comprising considerable thicknesses of shales, acted as a decollement, this must have allowed the younger, Late Jurassic to Neogene post-rift mega-sequence to be strongly deformed with large-scale recumbent folds, thrusts, and with some strong to intense imbrication, as is well exposed in the Kolbano region of SW Timor. Extensive development of pressure solution cleavage and inversion of strata, evidence of multiple deformation, and the presence of large-scale flat-lying overthrusts distinguish this post-155 Ma rift sequence (Barber et al. 1977) from the structures in the older Australian megasequence comprised of the pre-rift part of the para-autochthon. This younger of the two megasequences is exposed almost only in the southern quarter of Timor Island, although it contributes numerous exotic blocks and clay matrix to the Bobonaro Melange, while the older Gondwana mega-sequence crops out in the northern threequarters of the island. The widespread Bobonaro Melange and the associated broken formation facies suggest that without these two great decollements, D1 and D2, most of the two pre-collision sequences would have made Timor, as Fitch & Hamilton (1974) wrongly claimed it is, an island of mainly chaotic melange.

The Australian continental margin upper crust sub-Permian decollement D2 The oldest Australian continental margin sedimentary rocks of the Gondwana mega-sequence are Early Permian to late Middle Jurassic rocks that are highly deformed and exposed throughout Timor where all other rock sequences overlie the Gondwana mega-sequence. This distribution serves to emphasize the effectiveness of the decollements in the TCZ where they separated the megasequences tectonically. The folds of this older mega-sequence range from ‘large scale culminations and depressions up to c. 10 km wavelength, with limbs often broken, and in thinner bedded sequences crumpled into smaller folds of 10 m or less in amplitude with axial planes vertical or steeply dipping southwards. Many overturned folds are towards the south but some are overturned northwards. Hinges are

DEEP OROGENIC TROUGH TECTONIC PROCESSES

mostly broken and all rocks extensively affected by later faulting’ (Barber et al. 1977).

Asian fore-arc basement, sub-allochthonous Banda Terrane decollement D3 The Lolotoi (East Timor) and Mutis (West Timor) metamorphic complexes are found at the base of the Banda Terrane throughout Timor. For the most part, the Banda Terrane is underlain by the Bobonaro Melange wherever the base of these metamorphic thrust sheets can be seen (Standley & Harris 2009). The Bobonaro Melange was produced from the Jurassic shales of the Gondwana megasequence and from the clay-rich zones of the Eocene rocks of the younger post-rift megasequence. They could have contributed to the lubricant for the decollement that facilitated the transport of the Banda Terrane. The structurally high position of this terrane above the deeply eroded Gondwana mega-sequence, indicates it must have been emplaced relatively late in the collision process (Figs 2, 5–8) following the deep erosion of paraautochthon. The Bobonaro Melange is also often found on slopes of the allochthonous Banda Terrane.

The Australian Aileu – Maubisse Gondwana mega-sequence sub-Permian decollement D4 The sub-Permian decollement at the base of the Gondwana mega-sequence must be present at the base of the Aileu Complex, that is only found in north part of East Timor (Figs 2 & 8) and in some of the very small islands north and east of Timor including Kisar, Leti and Sermata (Kaneko et al. 2007). The Aileu Complex lithologies range from amphibolite facies metamorphic rocks to unmetamorphosed sedimentary rocks of Permo-Triassic age. This complex is found in the most southerly part of its crop to interdigitate with the paraautochthonous rocks of Maubisse Formation of Permo-Triassic age (Barber & Audley-Charles 1976; Prasetyadi & Harris 1996). The most northerly, highly metamorphosed part of the Aileu Complex is exposed only along a narrow north coast strip and in some small islands a few kilometres north and east of Timor. This part of the Aileu Complex reaches amphibolite facies grade (Berry & Grady 1981). The metamorphic grade declines southwards in a relatively short distance from the north coast of Timor, with slates and phyllites that interdigitate in the southernmost part of the crop with the Maubisse Formation. It thus appears that, except for the amphibolites, the Aileu Complex is a part of the Australian Gondwana sequence. These amphibolites must have been

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metamorphosed to the north of Timor, and their last provenance must have been at about 35 km depth in the fore-arc basement of the Banda Volcanic Islands. This metamorphism can be attributed to the heat associated in the fore-arc basement when the Banda volcanic islands of Alor, Wetar and possibly Atauro were active. Long after their metamorphism the Aileu amphibolites were carried upwards from this fore-arc lower crust in the Wetar Suture that emplaced them in the northern part of East Timor at about 2 Ma (Standley & Harris 2009; Spakman & Hall 2010). These metamorphosed sedimentary rocks seem likely to have been stacked in a deeper part of the TCZ than the unmetamorphosed Gondwana sequence (Fig. 8).

The Australian sub late – Jurassic post-rift mega-sequence decollement D5 The sub-Late Jurassic D5 decollement, mapped by Hughes et al. (1996) with BIRPS deep reflection seismic, was active in same Middle Jurassic Wai Luli Formation shales that function as the D1 decollement provoked by the collision event of about 4 Ma. However, the D5 decollement was a notably younger event than that involved in the D1 decollement. This D5 decollement is only present and active north of the deformation front located in the deep axis of the Timor Trough. South of the deformation front there is no D5 decollement, but of course the same Middle Jurassic shales continue to be present, where they are part of the undeformed Australian autochthon that occurs south of the bathymetric axis of the Timor Trough (Fig. 12). This D5 decollement thus marks a sharp transition from the Timor foreland, represented in the Timor Trough and Australian Shelf, and the collision zone north of the deformation front (Figs 8 & 12). An additional contrast (cf. Figs 8 with 12) is the lack of any indication of a decollement or deformation below the two undeformed megasequences that Hughes et al. (1996) mapped in the Australian autochthon south of the deformation front in the Timor Trough. This is another piece of evidence that conflicts with the suggestions that the Timor Trough is or has been a Benioff subduction zone.

Tectonic processes associated with the Banda Trench, Wetar Suture and Timor Trough The Wetar Strait is part of the Banda volcanic fore-arc and separates the volcanic islands Alor, Atauro and Wetar from Timor. The strait drops

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precipitously to more than 3 km depth and varies in width between about 25 km (Atauro), 35 –75 km (Alor) and 50 km (Wetar). In these islands most of the volcanic activity ceased in the Pliocene (Abbott & Chamalaun 1981). The Wetar Strait opens westward into a 200 km wide fore-arc basin of the Savu Sea that also descends steeply from the north coast of Timor, and from the submarine margin of the Roti-Savu Ridge (Audley-Charles 2004, Fig. 1). The downward flexing Australian plate comprising only mantle lithospheric slab and ocean crust (Snyder et al. 1996) is subducting below northern Timor, the Wetar Strait and the islands of the Banda Volcanic Arc. This subducted slab is traceable by earthquake seismology (Figs 8 & 9) to depths beyond 600 km (Engdahl et al. 1998).

The Banda Trench and the Wetar Suture The Banda Trench was the morphological c. 6 km deep trench to the east of 1208E that continued the Sunda–Java Trench from about 12 to 4 Ma. It was

located east of 1208E and south of 68S, where the subduction zone refused to subduct the Australian continental crust at c. 4 Ma. This is called the Banda Trench because its subduction of Indian Ocean crust gave rise to the Banda Volcanic Arc from c. 12 Ma. The Wetar Suture has carried northwards the Timor tectonic collision zone (TCZ), with the stacked c. 40 km thickness of Australian continental crust and allochthonous Asian nappes, over the Banda volcanic fore-arc. All those islands of the Banda Volcanic Arc that are now less than 50 km from the Australian continental margin have been dormant for over 2 Ma. This part of the Australian continental margin correlates positively with the North Timor–Wetar aseismic triangle, an area in which there are almost no hypocentres between 75 and 300 km depth (Engdahl et al. 1998). The Wetar Suture (Fig. 8) can be interpreted as a major thrust zone developed early in the collision phase, intimately associated with the lower crust whose response to the deforming forces was very different from the much softer, well stratified Australian upper crust Gondwana mega-sequence. The Australian continental lower crust in the Wetar Suture carried the Gondwana mega-sequence nappes. This is also implied by the BIRPS data (Richardson & Blundell 1996), and the gravity data of Woodside et al. (1989), re-interpreted in this paper. It is indicated by regional geology and the extraordinary amount of impingement of the Australian continental margin on the Banda fore-arc to the north of central and eastern Timor (Audley-Charles 1986a, b; Hall & Wilson 2000). The continental margin post-rift mega-sequence may also have been carried northwards over the Banda volcanic fore-arc, but these younger rocks have been eroded from most of the northern part of Timor.

The Wetar Suture and associated Australian lower crust

Fig. 9. BIRPS profile and earthquake seismology from Snyder et al. (1996, fig. 6a) with annotations by the present author: A is bathymetric deep axis of Timor Trough; B is Banda Trench now defunct and 40 km deep Timor TCZ, cf. Figure 8 of this paper. C is fore-arc basin, note steep, south-dipping slope of Banda Trench, now TCZ, corresponds to Wetar Suture on Figure 8. D is Banda volcanic fore-arc, E is Banda volcanic arc. F indicates that interpreted Australian continental crust extends only 6 km below the Banda Trench in the Snyder et al. interpretation (cf. Fig. 8).

The elevation difference between the bottom of the Wetar Strait and the north Timor coastal zone approaches 4 km in places. The exceptional steepness of the gradient between them suggests faulting, on what was postulated (Audley-Charles 1981; Audley-Charles 1986a, b) as the southward-dipping Wetar Suture (Figs 7 & 8). Richardson & Blundell (1996) mapped multiple south-dipping thrusts from BIRPS seismic data in the Banda fore-arc immediately east of Wetar and Kisar islands (Figs 10 & 11) that correspond closely with the postulated position of the Wetar Suture. Moreover, Masson et al. (1991) and Snyder et al. (1996) found a south-dipping thrust near the north coast of Kisar, but none of these authors noticed any connection with the postulated Wetar Suture. Breen

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Fig. 10. BIRPS seismic profile of ‘Damar Line’ from Richardson & Blundell (1996) with geological re-interpretation by the present author. A is the bathymetric 2– 3 km deep axis of Timor Trough; B is eastward continuation of submarine ridge of Timor Island 130 km east of Timor; C is fore-arc basin between Timor and Wetar east of those islands. D is Banda volcanic fore-arc; E is dormant Banda volcanoes. WS reflector is Wetar Suture and Australian lower crust (cf. Fig. 8).

Fig. 11. Cartoon of BIRPS deep seismic profiles immediately east of Timor Island summarized and interpreted by Richardson & Blundell (1996) and re-interpreted by the present author. A is 2– 3 km bathymetric deep axis of Timor Trough coincident with southern limit of deformation and southern limit of Banda Trench now the Timor TCZ filled with Australian continental crust and Asian nappes. B is the Hughes et al. (1996) BIRPS profile as interpreted by Richardson & Blundell 1996 in their figure 7 (cf. Figs 8 & 12 of this paper). C is Timor Island. D is northern limit of northward overthrust of Australian continental margin (represented by former Banda Trench, now the Timor TCZ) that was carried over 50 km of Banda fore-arc driven by lower crust in Wetar Suture. Note Wetar Suture principal reflector immediately north of Kisar Island. E is reflectors associated with Wetar Suture located in the Banda fore-arc.

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et al. (1989), Harris (1991) and Prasetyadi & Harris (1996) also found south-dipping thrusts close to the north coast of East Timor. Harris et al. (2009) mapping Savu Island (150 km SW of Timor) found a major south-dipping fault zone in North Savu that carries the Australian para-autochthon and the Bobonaro Melange, that is a part of this para-autotochthon, northwards over the Banda fore-arc basement. This Savu Thrust Zone is structurally similar to the Wetar Suture, but the two have no physical connection although they may have the same effect as the Wetar Suture, as shown by Rigg & Hall (2011). The overthrusting of the Banda fore-arc implied by gravity and BIRPS seismic data suggest the Wetar Suture was created by movements of the Australian continental lower crust. This suture (Figs 3 & 8) is found along most of the northern margin of Timor (Audley-Charles 2004) and may continue as far east as Sermata where it would have a strike length of c. 650 km. It has carried the Australian continental margin by as much as 50 km over the Banda fore-arc, so the suture seems likely to have had the Australian continental lower crust as its base, implying that the lower crust had become detached from the northward subducting part of the Australian continental sub-crustal mantle lithosphere by de-lamination at the Moho. The lower crust thrusts upwards and northwards from where this mantle lithospheric slab bends northwards to subduct below the northern parts of the TCZ. The lithosphere subducts below northern Timor, the Wetar Strait and the Banda volcanic fore-arc. BIRPS seismic and gravity modelling across the Inner and Outer Banda Arc east of Timor (Richardson & Blundell 1996) can be construed as supporting this interpretation of the Wetar Suture, as the Australian lower crust moving northwards on a major thrust zone over the sub-crustal lithosphere, and over the Banda volcanic fore-arc (Figs 8 –13). The driving force for this must be the pulling force of the .600 km deep northward dipping Australian lithospheric mantle slab (Engdahl et al. 1998). It is this subducting Australian sub-crustal mantle lithosphere that is pulling all the Australian crust from south of the point where the sub-crustal lithospheric slab bends downward and northward. Thus, the Australian crust above the Wetar Suture continues to ride over the Banda fore-arc as the subduction passage has been unable to accommodate any continental crust (Figs 7 & 8). This supports the suggestion of there being no Australian continental crust subducted from Timor. The crystalline lower crust must be present below what is now the exposed mega-sequence of Permian to Jurassic stratified, highly deformed, para-autochthon, whose base is seen nowhere. The

different physical properties of the crystalline lower crust, have led to it being overridden by what seems to have been the very effective subPermian decollement D2 (Figs 2 & 8). These properties of the lower crust would have facilitated detachment of the well stratified, much softer rocks of the upper crust Gondwana mega-sequence from the crystalline, relatively strong and rigid Australian continental lower crust during the deformation processes in the Banda Trench, and its progressive replacement by the TCZ during fore-arc collision from c. 4 to c. 1 Ma. The deformation of the two mega-sequences suggests that the lower crust retained its discrete response to the forces involved in the 4 Ma collision at the deepest structural levels (Figs 7 & 8).

The Timor Trough The Timor Trough is a depression 700 km long and between 30 and 75 km wide. Its depth is mainly about 2 km but reaches a maximum of 3.2 km. It is located entirely within what is now the Australian continental slope (as defined by bathymetry), its southern boundary being the shelf edge. It is underlain by Australian continental crust and lithosphere (Richardson & Blundell 1996; Snyder et al. 1996). The Timor Trough, partially filled by marine sediment, has the characteristics of a foreland basin (Audley-Charles 1981; Price & Audley-Charles 1987; Lorenzo et al. 1998; Hall 2002; Londono & Lorenzo 2004; Woodcock 2004). It continues as a bathymetric trough north-eastwards through the 1 km deep Tanimbar Trough, and then passes south and east of Kai, from where it follows the arcuate strike of non-volcanic Outer Banda Arc islands, passing through the Seram Sea north of Seram Island where it, like the Timor Trough, has a depth of mostly 2 km, but which, in the Seram Trough, deepens locally to more than 3 km. The Timor Trough, like the Tanimbar and Seram Troughs, are all Bally-type Ampferer subduction zones (defined by Bally 1983) because there is no ocean floor subduction linked to any of the three troughs. When the Banda Trench had an ocean floor that was subducting ocean crust from 12 to 4 Ma, that trench was north of the Australian continental margin into which the northwards-subducting slab sank into the mantle (Fig. 1). Since the blocking of that trench at 4 Ma no Australian crust has been subducted because it has been stacked in the posttrench tectonic collision zone (TCZ) in which some of that Australian continental crust has overridden the subducting sub-crustal lithosphere, and it is still overriding the Banda fore-arc (Figs 7 & 8). The morphological Banda Trench failed as a Benioff zone about 4 Ma when it was progressively converted into a TCZ. That also led to the Timor

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Fig. 12. A is an annotated BIRPS deep seismic profile across deep axis of Timor Trough (see also Figs 8 & 11). Reflectors identified by Hughes et al. (1996) as follows: A is Mid– Late Tertiary; B is Mid-Base Tertiary; C is Mid–Late Jurassic break-up unconformity; D is Top Permian? The basal decollement of the wedge is labelled.

Trough being formed as a perisutural foreland basin, and the underthrusting Australian crust moving under Timor is part of a Bally-type Ampferer subduction zone. South of the bathymetric axis of the Timor Trough is the Australian foreland to the Banda orogen. Hughes et al. (1996) showed by deep seismic reflection profiling that the Jurassic shales of the Wai Luli Formation, that are below the almost entirely submarine accretionary wedge off southern Timor, are there underthrusting northwards. However, south of the deep axis of the Timor Trough, there have been no active decollements because this region is part of the Australian autochthon. The two great mega-sequences whose stratigraphy in a strongly deformed state, are present north of the deep Timor Trough axis belong to the Australian para-autochthon (Figs 2, 8 & 12).

Late orogenic block faulting and uplift: possible relation to extension linked to Wetar Suture The timing of the uplift of Timor from a submarine position to become an island almost 3 km above sea level is indicated by the evidence of the first subaerial erosion provided by the siliciclastic turbidites

of the Viqueque Group (Kenyon 1974), whose age can be determined from the foraminifera. The lowest Noele Marl turbidites, described by D. J. Carter as high Upper Pliocene to Pleistocene N21 –N22 zone, equivalent to about 2.2 to 1.8 Ma, and widely exposed throughout Timor (Kenyon 1974), record the first erosion products produced from the emergence of northern Timor as an island undergoing sub-aerial erosion. This uplift was associated with extension of the Banda forearc produced by the rollback of the subducting lithosphere slab below Timor (Figs 7 & 8). Then the Wetar Suture thrust system broke through to the surface in the Wetar Strait. This could have reduced compression in the Timor tectonic collision zone that confined all the deformed Australian crust and Asian nappes until the Wetar Suture carried them to the north. The Late Pliocene to Pleistocene turbidites of N21 –N22 age are overlain by the Pleistocene fluvial gravels and fringing coral-algal reefs that are found on and around much of Timor Island today (Audley-Charles 1968; Kenyon 1974) indicating that Timor had begun to emerge as an island by the latest Pliocene or early Pleistocene at about 2.5 Ma. This uplift had been accompanied by PlioPleistocene faulting (Price & Audley-Charles 1987). The faulting affects the Quaternary reef limestone

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Fig. 13. Gravity data model of Timor region from Woodside et al. (1989) interpreted by Richardson & Blundell (1996, fig. 6). Densities of polygons are in kg m23. The model was intended to constrain crustal thicknesses in the central part of the Banda orogen. Note the grey area corresponds very closely with the defunct Banda Trench, now Timor TCZ, filled with Australian continental crust in Figure 8 of this paper. The gravity model annotations by present author are: A is bathymetric 2– 3 km deep axis of Timor Trough; B is Timor; C is Banda volcanic fore-arc; D is Banda volcanic Arc; E draws attention to the overthrusting of the fore-arc by the Australian continental margin represented by the fill in Banda Trench, and later TCZ fill, moved in Wetar Suture by Australian lower crust. Note strata south of the Timor Trough axis are all horizontal, being the autochthonous Australian foreland to the non-volcanic part of the Banda Orogen represented here in grey.

plateaus (see the 1:250 000 geological maps of East and West Timor: Audley-Charles 1968; Rosidi et al. 1996; Suwitodirjo & Tjokrosapoetro 1996), some of which have also been gently folded (Kenyon 1974; Audley-Charles 1986a reported N22 and N23 gentle folding and tilting).

Conclusions: collision tectonics of the Banda Orogen – Timor region The early evolution of the Banda Trench led first to the development of the Banda Inner Volcanic Arc from about 12 Ma, and then after the 4 Ma collision the Banda Outer Non-Volcanic Arc developed into the 2300 km long fold and thrust mountain belt reaching 3 km above sea level.

The most important tectonic event in building the Timor orogen was the blocking of the Australian continental crust from subducting at the Banda Trench. This was probably caused by the c. 10 km thick lower crust being driven over the subduction passage at the bottom of the northern limit of the Banda Trench where it met the Banda fore-arc at c. 6 km depth. This blocked any relatively low density continental crust from subducting. Another early important event was continuation of the rollback of the sub-crustal mantle lithosphere below the Banda Trench by 30 km after 4 Ma together with the pull-down effect of the subducting subcrustal mantle lithosphere. Many other key tectonic events followed. The continual pull of the subducting sub-crustal lithosphere was the engine for the movements of the two major decollements below the Gondwana mega-sequence and post-rift megasequence; and for the detachment of the Australian lower crust from the Moho above the zone where the mantle lithospheric slab continued to subduct. This enabled the lower crust to ride over the fore-arc; and was responsible for the 800 km long Wetar Suture thrust zone. The removal of the allochthonous Banda Terrane metamorphic basement with its range of Asian cover rocks including the Cablac Limestone from the fore-arc, and of the Asian allochthonous Aileu amphibolites also being removed from the fore-arc, all followed from the creation of the Wetar Suture. The range of Bobonaro Melange facies were a consequence of the initial chaotic compression and shearing of Jurassic Wai Lui shales and the Eocene clay-rich rocks present in the tectonic collision zone. I am much indebted to discussions and joint fieldwork over many years with Tony Barber, Luigi Giani, Robert Hall, Ron Harris, Chris Kenyon, John Milsom; and most especially with David Carter whose micropalaeontology supported our fieldwork for years, and with whom I benefited from many discussions. Ron Harris has given me many helpful tutorials by email since 2003 and he also helped by reviewing this manuscript. Robert Hall has given invaluable help with the early version of the manuscript, the illustrations and with tutorials on rollback, and his review of the MS was helpful and constructive. His indication, when I began to write this paper, of the probable absence of any Australian continental crust subduction from the Timor region was a turning point, and it is of profound importance. His invitation to write this valedictory paper was a stimulation and encouragement. I have been helped by discussions and correspondence with Marcelle K. BouDagher-Fadel who checked and corrected micropalaeontological identifications and zonation status. Ian Watkinson redrew many of the figures making notable improvements in the process. I remember my indebtedness to the late Sahat Tobing, and to many other local people in the islands where I have explored the geology. My fieldwork in Timor would not have been possible without the dedicated help of Raimondo Soares, Domingo Soares,

DEEP OROGENIC TROUGH TECTONIC PROCESSES Anotonio Salshinha, Valenti de Jesus, Nahac Mauc and Ah Sung. My fieldwork was supported and sustained by Timor Oil Ltd and International Oils Ltd over very many years.

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Physical oceanography of the present day Indonesian Throughflow DEBRA TILLINGER Lamont-Doherty Earth Observatory of Columbia University, 61 Route 9W, Palisades, NY 10964, The United States of America (e-mail: [email protected]) Abstract: The Indonesian Throughflow (ITF) transfers c. 15 Sv (1 Sv ¼ 106 m3s21) of relatively cool, fresh water from the tropical Pacific Ocean to the tropical Indian Ocean. Additionally, the ITF is a key interocean component of the global ocean warm water route, which returns water from the Pacific Ocean to the Atlantic Ocean to close the loop of the thermohaline overturning circulation associated with North Atlantic Deep Water. That flow consequently freshens the Indian Ocean and transports heat between basins. The ITF can also be described by the island rule, which relates the winds over the entire South Pacific Ocean to the magnitude of the ITF. El Nin˜o-Southern Oscillation (ENSO) dominates the regional variability in the Pacific Ocean and exerts a strong control over the variability of ITF transport. The Indian Ocean responds to the ENSO signal as well, but is also influenced by the Indian Ocean Dipole, a climate phenomenon that may act independently of ENSO to affect the ITF.

On a local scale, the surface layer of the ITF is controlled by local winds, which are primarily monsoonal, but at depth the ITF responds to the pressure gradient between the Pacific Ocean and the Indian Ocean. Newly available observational data from within the major straits of the Indonesian Seas allow for an improved resolution of the total ITF and its variability. Makassar Strait, which is the main inflow channel of the ITF, transported an average of 11.8 Sv of North Pacific water. Lifamatola passage, a route for deep South Pacific water, transported 2.7 Sv below 1250 m and returned 0.9–1.3 Sv to the north. The Karimata Strait has not been monitored observationally, but models estimate a contribution of 1.4 Sv, which is likely important for its role in the transport of freshwater. Transport through the outflow passages was 2.6 Sv through Lombok Strait, 4.9 Sv through Ombai Strait, and 7.5 Sv through Timor Passage. Temporal variability within the straits is seen on timescales from the interannual ENSO signal to semi-diurnal tidal signals. The Indonesian Throughflow (ITF) transports c. 15 Sv (1 Sv ¼ 106 m3 s21) from the Pacific to the Indian Ocean. The structure and magnitude of the ITF varies on temporal scales from days to decades and even over geological time. As the water is transported, its hydrological characteristics are altered by heat and freshwater inputs from the Indonesian Seas and by strong vertical mixing. The physical oceanography of the present day ITF therefore cannot be characterized by a mean transport alone and requires an understanding of the surrounding current systems and climatic conditions. The ITF can be understood as both a local phenomenon, affecting the Indonesian Seas and the edges of the western Pacific and eastern Indian basins,

and as a global phenomenon, providing a path for water to circumnavigate the world ocean. The objective of this paper is to understand the significance of the ITF and its variability in both a global and a local sense. The first section places the ITF in the context of global ocean circulation, showing its importance to thermohaline circulation, estimates of volume, heat, and salt transport, and the framework of large-scale wind-driven circulation. The next section narrows in focus from a global to a regional scale, providing the physical oceanography and prevailing climate phenomena that interact with the ITF. The ITF is considered on a local scale in the following two sections. The third section develops the theoretical frameworks for understanding the existence of the ITF, while the fourth provides observational data for understanding the mean flow and temporal variability within individual straits. The study concludes with a discussion of the implications of the ITF and directions for future work.

Global ocean circulation Thermohaline circulation The circulation of the world ocean is often compared to a conveyor belt. In the North Atlantic, cold, salty surface water sinks and spreads southward through the Atlantic Ocean as North Atlantic Deep Water (NADW). That NADW is then carried by the Antarctic Circumpolar Current and deep western boundary currents to the Pacific and Indian Oceans. In order for this transport to occur, there must be a return flow of comparatively warm water balancing the mass transport. Gordon (1986) proposed that this exchange must happen in the

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 267–281. DOI: 10.1144/SP355.13 0305-8719/11/$15.00 # The Geological Society of London 2011.

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oceans’ thermocline, the layer of seawater between the surface mixed layer and the cooler deepwater in which the water temperature changes rapidly with depth. This occurs via the ITF and the subsequent export of those waters from the Indian Ocean into the South Atlantic through the Agulhas Leakage. This concept, now well accepted in principle, requires that the ITF balance the transport of mass, heat, and salt from the Pacific to the Indian Ocean. The metaphor of the ‘conveyor belt’ originated with Broecker & Peng (1982) and referred to a conceptual model of global circulation (for a discussion of the history of the ‘conveyor belt,’ see Richardson 2008). The concept was popularized through a magazine article (Broecker 1987) that showed a simplified version of the circulation described in Gordon 1986 (Fig. 1). This version has been reproduced many times, although it fails to show the true path of the ITF through the Indonesian Straits and neglects the formation of Antarctic Bottom Water A more sophisticated conveyor belt, developed by Gordon (1991) and updated by Schmitz (1996) and Lumpkin & Speer (2007) attempts to show a more three-dimensional view of global ocean circulation (Fig. 2). The structure of these later diagrams emphasize the fact the ITF is the only low latitude connection between ocean basins, and the largest interocean flow outside of the Southern Ocean. As the conceptual framework of a conveyor belt expands to include more details about ocean

circulation, it builds in complexity and becomes more difficult to show schematically. However, a more crucial concern is that it gives the impression of an ocean governed by a strong mean flow, with little, if any, variability. Such a deterministic view of ocean circulation is contradicted by observational evidence, including evidence that the ITF is highly variable in its transport not only of volume, but of heat and salt as well.

Transport of volume, heat and salt Whether ocean circulation is best described by the mean flow or by temporally variant eddies and currents, transport of mass, heat and salt must balance over long timescales. The World Ocean Circulation Experiment (WOCE) project attempted to quantify these fluxes using observational data. These data, along with satellite and other data, are then incorporated into general circulation models which can provide an estimate for both the mean state and the variability of these fluxes (see Wunsch et al. 2009 for a full discussion of these methods). Volume transport is the rate of volume flow across a unit area, calculated as the product of the integrated velocity through the area and the cross section of that area. Given that the mass of seawater is not variable, conversions between volume and mass transport are simply a matter of units and the two terms are often used interchangeably. The

Fig. 1. The ocean conveyor belt as proposed by Broecker (1987).

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Fig. 2. A more accurate schematic of global ocean circulation (Gordon 1991).

volume transport of the ITF has been estimated from to 5 to 15 Sv (Gordon 1986; Godfrey 1996; Gordon et al. 2009). An accurate estimation of volume transport is required for any calculation of heat or salt transport. However, the calculation of heat transport, more precisely referred to as internal energy transport (see Warren 1999 for discussion), also requires the use of a reference temperature. This constraint is inherent to the calculation and arises from the need to consider not only the flow through a given section but the eventual export (or previous import) of those waters as well. The heat transport is the difference in heat capacities of the water being imported and the water being exported. Thus, all heat transport calculations are somewhat dependent on the choice of reference temperature, although volume transport has a far larger effect. In the case of the ITF, heat transport estimates range from 0.24 PW (Vranes et al. 2002) (1 PW ¼ 1015 watts) to 1.15 PW (Schiller et al. 1998), and reference temperature estimates range from 0 8C to 3.4 8C, the presumed temperature of Pacific Ocean inflow or Indian Ocean outflow. Because salt comprises only a small fraction of seawater, it is computationally simpler to think of the flux of freshwater, rather than the flux of salt. The ITF transports water from the fresher Pacific Ocean to the saltier Indian Ocean. Like heat flux, freshwater flux considers the difference in salinity

between the inflow and outflow waters. In this case, the water entering the Indian Ocean via the ITF is less saline than the water exiting the Indian Ocean. The difference results in a freshwater transport of 0.23 Sv. Since the Indian Ocean is not becoming increasingly fresh, freshwater must be removed via net evaporation over the Indian Ocean (Talley 2008).

Wind-driven circulation Although much of the large-scale circulation in the ocean is ascribed to thermohaline forcing, it can also be understood in terms of global patterns of wind forcing. Wind stress, the drag exerted by air moving over the ocean, can be used to calculate the volume of water integrated from the sea floor to the sea surface being transported north or south, assuming an ocean in which local velocity is governed by density differences on a rotating Earth. This relationship, the Sverdrup balance (Sverdrup 1947), was extended to calculate flow around islands (Godfrey 1989). According to Godfrey’s island rule, the transport around an island can be calculated using the integral of the wind stress along a closed contour encompassing the island and the entire width of the ocean basin located eastward of that island. In the case of the ITF, the relevant island is Australia-New Guinea and the path can be seen in Figure 3. Godfrey’s original

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calculation led to a long-term mean ITF of 16 + 4 Sv. The island rule was later modified to include the effects of friction and bottom topography (Wajsowicz 1993), showing that the Indonesian Seas modify the ITF and reduce its magnitude. That study further suggested that the island rule could be extended to include interannual variability of ITF in addition to the long-term mean. Wajsowicz (1996) used observational wind data to calculate the interannual ITF variability. Using an ocean circulation model, Humphries & Webb (2008) compared interannual variability of the ITF to that produced by the island rule. They found that the island rule agreed well with the model when run using annually repeating monthly climatological winds, but less so when the model was run with more realistic winds. This may be due to the time required for signals generated within the Pacific Ocean to propagate to the ITF. However, given that the same wind stress data were used in both the model and the island rule calculation, it is not possible to conclusively quantify the relationship between the in situ wind stress and the in situ transport.

Regional circulation and climate Western equatorial Pacific Ocean The western equatorial Pacific Ocean contains a complex set of currents and several eddy systems (Fig. 4). The dominant currents are the westwardflowing North and South Equatorial Currents and the eastward-flowing North and South Equatorial Counter Currents and Equatorial Undercurrent. The inflow area for the ITF lies at approximately the same latitude as the North Equatorial Counter Current, in an area with two substantial eddies: the Mindanao Eddy to the north and the Halmahera Eddy to the south. North Pacific Intermediate Water (NPIW) is defined as the salinity minimum in the subtropical

North Pacific (Talley 1993). That water mass is formed by the mixing of cold, fresh subpolar waters with saltier subtropical waters and occupies the intermediate depths (300–700 m) of the northern equatorial Pacific Ocean. In the south, the thermocline contains Antarctic Intermediate Water (AIW), which is colder and saltier than NPIW and is formed by overturning in the Southern Ocean. Overlying AIW is South Pacific Equatorial Water, which is also colder and saltier than its northern counterpart, North Pacific Equatorial Water. Godfrey et al. (1993) argued that although the relatively fresh waters of the ITF suggested a northern source for the throughflow, the source must actually be to the south to account for the island rule’s dependence on wind stress to the south of the ITF. Although most of the South Equatorial Current retroflects to join the North Equatorial Counter Current, some could conceivably leak from the Halmahera eddy to form all or part of the ITF. South Pacific water is distinctly saltier than both North Pacific and ITF water, so they argued that this South Pacific source water was freshened by rainfall so that it appeared to be North Pacific water by the time it reached the ITF. The distinction here between North and South Pacific water is partly semantic, given that all North Pacific water was, at some time in the past, South Pacific water. North Pacific water forms when South Pacific water, following the South Equatorial Current, turns back to the east in the Halmahera Eddy (Fig. 4). That water flows eastward along the North Equatorial Countercurrent, then turns again to flow westward along the North Equatorial Current. Only when it reaches the western boundary of the Pacific a second time does it join the Mindanao Current and enter the Indonesian Straits. Gordon (1995) argued that the ITF source water was from the North Pacific and that its characteristic low salinity was the result of the excess precipitation throughout that region. He did, however, mention

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Fig. 4. Mean currents of the Pacific Ocean and Indian Oceans during the spring (top panel) and fall (bottom panel). The following currents are shown: North Equatorial Current (NEC), North Equatorial Counter Current (NECC), South Equatorial Current (SEC), South Equatorial Counter Current (SECC), Equatorial Counter Current (ECC), Somali Current (SC), South Java Current (SJC), Leeuwin Current (LC) and SW Monsoon Current (SWMC). Note that in the Indian Ocean, the SC and SJC reverse, while the ECC and NEC are replaced by the SWMC.

that a small amount of South Pacific water might enter the ITF below the thermocline. Nof (1996) used a theoretical approach to show that the complex geography of the Indonesian archipelago and the existence of retroflecting currents controls the source water of the ITF and leads to an ITF that is primarily from the North Pacific with only a small contribution from the South Pacific. Later studies of water mass analysis have confirmed the primacy of the North Pacific as the source of the ITF (Ilahude & Gordon 1996). The surface conditions are dominated by the trade winds, which blow from east to west and

towards the equator in both hemispheres. They meet at the intertropical convergence zone, which has an average location of 58N. The trade winds establish a sea level gradient, with higher sea level in the western side of the basin, and an associated gradient in the thermocline, with a deeper, thicker layer of warm water in the west forming the western Pacific warm pool.

El Nin˜o-Southern Oscillation El Nin˜o-Southern Oscillation (ENSO) is centred in the Pacific Ocean but can influence global climate

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through oceanic and atmospheric teleconnections. Its effect on the ITF is far more localized, as it changes the ITF source waters directly. When the trade winds slacken and the zonal sea surface height and temperature gradients weaken or reverse, creating an El Nin˜o event, the magnitude of the ITF decreases. The opposite effect is seen during a La Nin˜a event, with increased trade winds, an increased sea surface height and temperature gradient, and an anomalously large ITF. Meyers (1996) used a twelve-year record of observational hydrological data from the outflow of the ITF to construct a record of ITF variability. He found that the strongest signal within the data corresponded to the ENSO signal. The same result has been seen in modelling studies (Clark & Liu 1994; England & Huang 2005). The strength of ENSO can be recorded by various indices, among them the NINO3.4 index, which measures the sea surface temperature anomaly in the eastern-central Pacific. Although major ENSO events can be very clearly seen in all indices and measures of ENSO, smaller events can elude classification and remain a subject of scientific debate (Meyers et al. 2007).

Equatorial Indian Ocean The Indian Ocean is notable for its seasonal variability. The SE Asian monsoon dominates the equatorial Indian Ocean and leads to current reversals. During the winter (all seasons are defined relative to the North Hemisphere), the monsoon brings northeasterly winds from the high pressure centre over Asian landmass. During this time, the current system resembles that of other ocean basins, with a North Equatorial Current, a South Equatorial Current, and an Equatorial Counter Current (Fig. 4, top panel). In the west, the Somali Current brings water south from the North Equatorial Current to the Equatorial Counter Current. In the east, the South Java Current flows toward the Indonesian archipelago and curves to the south, bringing water from the Equatorial Counter Current to the South Equatorial Current. The situation reverses during the summer, when there is an atmospheric low over the Asian landmass and corresponding southwesterly winds (Fig. 4, bottom panel). The Somali Current reverses to flow to the north and South Java Current reverses to flow to the west. In place of the North Equatorial Current and Equatorial Countercurrent, the SW Monsoon Current flows from west to east. During both seasons, the Leeuwin Current flows southward along the western coast of Australia. The transitional periods between summer and winter monsoons, typically during May to June and October to November, are accompanied by Wyrtki jets, which are strong equatorial currents travelling from west to east. Over the most of the

Indian Ocean and the Indonesian Seas, the monsoonal winds are directed from the NW in winter and from the SE in summer. In the Indian Ocean, the warm pool is on the eastern side of the basin and can be seen as an extension of the western Pacific warm pool. Accordingly, the water in the eastern Indian Ocean is warmer and fresher than the water in the western Indian Ocean.

Indian Ocean Dipole Just as an El Nin˜o event is associated with a slackening or reversal of the trade winds, the Indian Ocean can experience a similar phenomenon when the summer southwesterly monsoon winds blow from east instead. This leads to anomalous warming and freshening of the western Indian Ocean and the upwelling of cooler water in the eastern Indian Ocean. This is considered the positive phase of the Indian Ocean Dipole (IOD). In its negative mode, the IOD strengthens the average conditions of the Indian Ocean, further warming the east and cooling the west (Saji et al. 1999). The IOD and ENSO can co-vary, prompting a debate over the existence of the IOD as an independent climate phenomenon or merely as an extension of ENSO dynamics (see Meyers et al. 2007, for a review). Positive IOD events generally occur during El Nin˜o years, and negative IOD events generally occur during La Nin˜a years. They can, however, occur independently (Yamagata et al. 2004). Like ENSO, the IOD is tracked by an index of sea surface temperatures, called the Dipole Mode Index (DMI). In this case the difference in sea surface temperature anomaly between the western and southeastern Indian Ocean. The IOD can also be tracked by considering just the eastern sea surface temperature anomalies (Meyers et al. 2007). As in the case of ENSO, there remains uncertainty in the classification of many years as IOD events.

The Indonesian Seas The complex geography of the Indonesian Seas can be seen in Figure 5. The area shows complex bathymetry as well, with deep basins and many sills (Gordon et al. 2003a). The main path of the ITF consists of water entering from the North Pacific between the Philippines and New Guinea into the Sulawesi (Celebes) Sea and continuing through Makassar Strait. From there, the water can exit via Lombok Strait or circulate through the Banda and Flores Seas and can enter the Indian Ocean via Ombai Strait or Timor Passage. In addition, deeper South Pacific waters enter Lifamatola Passage and provide cooler, saltier water to ventilate the Banda

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Fig. 5. Bathymetry of the Indonesian Seas. The direction of mean flow through each strait is indicated with arrows.

Sea (van Aken et al. 2009). Surface water also enters the Flores Sea from the South China Sea via Karimata Strait (Qu et al. 2005). The Indonesian Seas experience significant freshwater input from rainfall and river runoff and are warmed by surface heat fluxes south of Makassar Strait (Wijffels et al. 2008). Internal tides are trapped within the various semi-enclosed seas, resulting in strong vertical diffusivity that mixes the buoyant surface water downwards. Although the presence of these strong tides does not change the magnitude of the ITF, it changes the temperature and salinity of the outflow water (Koch-Larrouy et al. 2007).

Local forcing of the ITF Pressure driven flow. Wyrtki (1987) observed that a pressure gradient between the western Pacific Ocean and the eastern Indian Ocean must exist and govern the ITF. This is true whether that gradient exists due to the island rule, thermohaline circulation, or any other cause. Although pressure differences alone cannot give a numerical value for the ITF, they can quantify its variability relative to an unknown mean. To investigate this, Wyrtki compared sea level from Davao in the Philippines and Darwin in Australia to develop a time series

of ITF variability relative to an unknown mean. He hypothesized that the trade winds in the Pacific built up a pressure head from the Pacific Ocean to the Indian Ocean. However, the sea surface signal at Darwin is more representative of the Pacific Ocean than the Indian Ocean, which may explain why the resultant time series showed no correlation with ENSO. Since modern observations show a strong ENSO signal, this approach must be modified by using data that are more representative of the Indian Ocean. Furthermore, differences in the sea surface do not fully describe differences in pressure below the sea surface. The ITF has been shown to be primarily baroclinic (Waworuntu et al. 2001), meaning that pressure surfaces, such as the sea surface, are inclined relative to density surfaces. To address the problem of accurately representing sea surface height in both the Indian and Pacific Oceans, Potemra et al. (1997) and Potemra (2005) used sea level derived from satellite altimetry at multiple locations on the Pacific and Indian sides of the ITF to calculate the pressure difference between the two ocean basins. In addition, they expanded the concept by allowing the sea level changes to lead or lag the ITF signal to account for the time needed for signals to propagate through the basins. The resultant ITF series does not show a correlation with ENSO, which may

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again be explained by the baroclinic nature of the flow. Potemra et al. (2003) noted that the ENSO signal propagates through deeper layers of the ITF. To address the issue of subsurface forcing on the ITF, Burnett et al. (2000a, b) explored the possibility that drag from the complex bathymetry of the Indonesian Seas could cancel the pressure head from the Pacific Ocean, meaning that the pressure head would not control the ITF. The model used in that study is barotropic, neglecting the relative tilt of pressure surfaces to density surfaces, meaning that the pressure head was calculated solely as the difference in sea surface heights. The results showed that transport through individual channels would change in response to increased or decreased sea surface height differences, but that the total ITF would not. Burnett et al. (2003) and Kamenkovich et al. (2003) further explored the relationship between the interocean pressure difference and the total ITF transport in a two-part study. Their results emphasized the importance of bathymetry and indicated the pressure head, as calculated by a variety of different methods, was correlated with the seasonal variation of the ITF but could not uniquely determine its value. That is, a particular pressure head could exist for multiple values of ITF transport. They found that the transports from the Pacific currents into and away from the Indonesian Seas also influenced the total transport of the ITF. The interocean pressure gradient can be used to provide a numerical value for velocity in addition to its variability. This is done by calculating a geostrophic balance, which assumes that velocity must balance a pressure gradient, the effect of the Earth’s rotation and friction. Unlike the pressure difference alone, which can at best calculate the variability of a flow relative to an unknown mean, geostrophy provides a velocity relative to a deeper layer, ideally one presumed to have no motion. This method was used by Meyers et al. (1995) and Meyers (1996) with observational temperature data collected in the Indonesian Seas and Indian Ocean by volunteer merchant ships using expendable bathythermographs (XBTs). The XBT temperature data and mean salinity were used to calculate velocity in the top 400 m of the water column. They found a strong relationship between monsoon winds, which can change the pressure gradient from the Indian side, and the ITF. Potemra et al. (2002) also used this method to calculate velocities in the outflow channels and compare them to in-situ observational data and model output. They found the flow to be in geostrophic balance. Wijffels et al. (2008) used an expanded XBT dataset and found an average ITF of 8.9 Sv. However, they suggest that this estimate is too low

because it does not account for the deep portion of the ITF. Geostrophic balance calculations are useful for the individual outflow straits, but are difficult to apply to the inflow straits and thus to the ITF as a whole. This is because geostrophic conditions do not hold along the equator and in narrow straits, such as Makassar (Burnett et al. 2000a). This is particularly problematic because most of the ITF is transported through Makassar Strait. In addition, the Banda Sea acts as a capacitor, holding water for a varying amount of time before it is released into the Indian Ocean (Gordon & Susanto 2001). Fieux et al. (1996) noted that geostrophic currents could also be affected by friction resulting from the proximity of strong currents to coastlines. Song (2006) used both sea surface height and ocean bottom pressure to estimate the ITF, combining the techniques of geostrophy and hydraulic control of a strait. The study found the combination of both factors gave a better approximation to the total ITF than did either one alone. Tillinger & Gordon (2009) used interocean pressure differences calculated on layers of constant density to construct a timeseries of ITF variability. The resultant variability was found to match observational data on an interannual timescale and island rule calculations on a decadal timescale. Local wind driven flow. In addition to the geostrophic pressure-driven flow, a portion of the ITF is ageostrophic and driven directly by local winds. The wind drives Ekman transport, which relates local wind stress to transport, modulated by the rotation of the Earth. The ITF can be divided into several vertical layers, with the top layer forced by surface winds and the deeper layers responding to remote forcing (Potemra et al. 2002). The Ekman and geostrophic transports may interact, with the southward Ekman flow raising sea level and consequently deepening the thermocline and changing the associated pressure gradient. Although the existence of Ekman transport in the ITF is clear, its magnitude is uncertain. Murray & Arief (1988) noted that the surface flow could reverse to the north and attributed those events to cyclonic wind events that are part of the monsoonal cycle. Potemra et al. (1997) found that the ageostrophic Ekman transport was much smaller than the geostrophic transport but could account for as much as 1.8 Sv or, 25% of the total ITF, and was highly variable. Sprintall & Liu (2005) used scatterometer and SST data to more precisely calculate the Ekman contribution to the ITF, and found that Ekman transport can be the same order of magnitude as the geostrophic flow through the outflow straits, but reverses with the phase of the monsoon.

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The summer monsoon is associated with sustained winds leading to southward transport. The winter monsoon leads to strong but sporadic wind bursts over the Indonesian Seas, which lead to northward transport that balances the previous southward transport. So despite the significant contribution of Ekman transport to the total ITF at the outflow straits on sub-annual timescales, it does not significantly impact the ITF and timescales longer than one year. These reversals were confirmed by observational data from within the straits (Sprintall et al. 2009). Tillinger & Gordon (2009) found that the surface layer flow within Makassar Strait correlated with local wind variability. Between Australia and Java, further from the outflow straits, calculations from wind data suggest that Ekman transport accounts for 40% of the total transport over a twenty-year period (Wijffels et al. 2008).

Individual straits The International Nusantara Stratification and Transport (INSTANT) program deployed 11 moorings over a three-year period to obtain simultaneous observations of the major straits of the ITF: Makassar Strait, Lifamatola Passage, Lombok Strait, Ombai Strait, and Timor Passage (Fig. 5). Earlier studies of the ITF exist, but they did not observe all of the passages simultaneously. By convention, transport values are described here as positive when they are directed from the Pacific to the Indian Ocean. The most recent transport values from INSTANT (shown in Table 1) give a total transport of 15 Sv. Makassar Strait. Makassar Strait was monitored by two moorings placed at its narrowest constriction, the 45 km wide Labani Channel, analysed by Gordon et al. (2008). The resultant velocity profile (Fig. 6, top panel) shows a peak within the thermocline, at 110 to 140 m, with maximum velocities of nearly 1 m/s. The surface flow shows average velocities of c. 0.5 m/s, dropping to nearly zero below the 680 m depth of the Dewakan sill. However, velocities remain slightly positive on average down to a depth of 1500 m. A similar profile was seen during the Arus Lintas Indonen (ARLINDO) observational period of 1997, when moorings were placed in the Labani Channel (Gordon et al. 1999; Susanto & Gordon 2005). The summer monsoon is associated with greater maximum velocities, but lower velocities in the deeper flow below c. 250 m. Transport (Fig. 7, top panel) was calculated by multiplying the velocity by the depth-dependent cross section of the Labani channel. The minimum transport, 8.8 Sv, is seen from October to December, when the monsoon is in a transitional phase. The April to June transition also shows low transport

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Table 1. Transport and TWT through the straits of the ITF Strait Makassar Strait Karimata Strait Lifamatola Passage Lombok Strait Ombai Strait Timor Passage

Transport (Sv)

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11.6 1.4* 2.7 2.6 4.9 7.5

15.6 – 3.2 21.5 15.2 17.8

*There are no observational data available for Karimata Strait. This value is from the modelling results of Tozuka et al. (2007). No TWT was available.

(11.7 Sv), while transport is higher from July to September (12.6 Sv) and reaches a maximum of 13.1 Sv in January through March. The mean transport during INSTANT was 11.6 Sv, which is an increase of 27% from the ARLINDO value of 9.2 Sv. The difference can be attributed primarily to the El Nin˜o event of 1997– 1998, which would have substantially changed conditions in the region. The INSTANT period was primarily neutral or in a weak El Nin˜o, and ended with a La Nin˜a event. Both observational periods occurred during positive IOD events. Tillinger & Gordon (2009) suggest that the surface flow, normally due to Ekman transport and therefore independent of the pressure-driven flow, would have been dominated by ENSO signal during the ARLINDO period, lowering the total transport. These observational data confirmed the strong effect of ENSO on the volume of ITF transport. The transport-weighted temperature (TWT) of the flow is the average temperature in the channel, weighted by the transport at different depths. Gordon et al. (2008) found an average TWT of 15.6 8C. The seasonal pattern of TWT follows the seasonal transport pattern, with a minimum in the October to December transitional period between monsoons and a peak in the January to March winter monsoon. During ARLINDO, Vranes et al. (2002) found a TWT of 15.2 8C. Since the TWT is a function of the vertical distributions of temperature and velocity, it can be considered the heat transport per unit volume transport. A change in the total transport but not in the velocity distribution will not change the TWT, but it will change the heat transport (see section 1.2 for a discussion of heat transport). In periods of greater transport, the location of the maximum current speed deepens and therefore contains cooler water. Therefore when volume transport increases, the TWT decreases, but the heat transport increases. This disconnect between the TWT and the heat transport is a result

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of both the intensified flow within the thermocline through Makassar and the tendency of the velocity maximum to deepen when total transport increases (Tillinger & Gordon 2010). Gordon et al. (2003b) attribute the subsurface transport maximum and the corresponding low TWT as the result of fresh Java Sea water pushed in to the southern end of Makassar Strait during the winter monsoon. This thin sheet of freshwater in the top 50 to 100 m prevents the flow of warm Pacific surface water into the Indian Ocean. Karimata Strait. An alternative explanation for the profile of transport through Makassar Strait is the circulation of water from the South China Sea through the Karimata Strait, known as the South China Sea Throughflow (SCST). Qu et al. (2005) suggest that Pacific water travels between Taiwan and the Philippines through the Luzon Strait to the South China Sea, through Karimata Strait into the Java Sea, then northward through Makassar Strait to return to the Pacific Ocean. This process is a direct response to the winds in the western Pacific Ocean and therefore contains a strong ENSO

signal. With the inclusion of the SCST, the ITF through Makassar Strait is a combination of two circulations rotating in the same direction. Both are forced by the Pacific wind regime: Makassar contains a thermocline flow around Australia while the SCST contains a surface flow around the Philippines. Modelling results suggest that without this flow, the vertical profile of the ITF through Makassar Strait would be constant through the thermocline and surface layers, leading to a higher TWT and heat transport (Qu et al. 2005). By modelling the ITF with and without the SCST, Tozuka et al. (2007) find that the SCST lowers the volume transport of ITF through Makassar Strait by 1.5 Sv and decreases its TWT by 2.1 8C. Other studies (Gordon 2005) suggest that the main impact of the SCST on the ITF is the addition of freshwater to the Indonesian Seas, which changes the hydrographic characteristics of the outflow water without necessarily modulating the inflow water. Koch-Larrouy et al. (2008) demonstrated the importance of the SCST to the salinity budget of the Indonesian Seas and the Indian Ocean despite its small mass contribution.

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Fig. 7. Transport timeseries during the INSTANT period. Data for the Makassar Strait are filtered at 30 days (Gordon et al. 2008). A 15-day running mean has been applied to the Lifamatola timeseries (van Aken et al. 2009). Data for Lombok, Ombai and Timor were filtered at four days and daily values extracted (Sprintall et al. 2009).

Lifamatola passage. While Makassar Strait provides most of the volume transport of the ITF, it cannot ventilate the deep Banda Sea. The deep ITF transport travels an eastern route through the Maluku Sea to Lifamatola Passage, which has a sill depth near 2000 m (van Aken et al. 1988). This cold water sinks to depths of 5000 m, and then mixes and upwells to near 1000 m (van Aken et al. 1991). Water mass analysis from WOCE data has shown that the water in the deep Banda Sea is derived from saltier South Pacific water. Those data also suggest an intermediate ITF of 3 to 7 Sv (Talley & Sprintall 2005). Observational data from INSTANT were gathered using moorings located slightly downstream of the Lifamatola sill for 34 months and analysed by van Aken et al. (2009). Due to instrument error and unexpectedly strong tides, there were

difficulties with data collection, particularly during the first half of the observational period. From January 2004 through July 2005, uninterrupted current data are available from 1000 to 1500 m and from July 2005 through April 2006, they are available from 1000 to 2000 m. Limited velocity data are available from 1000 to 300 m during the second half of the observational period. Most of the variability in the currents comes from tidal forcing, showing fortnightly variability. The average flow (Fig. 6, second panel) is near zero or towards the Pacific Ocean at c. 0.1 m/s in the upper part of the water column, from 250 m to 1250 m (data shallower than 250 m are not available). Below 1250 m, the flow is directed towards the Indian Ocean at velocities increasing to 0.7 m/s near 1950 m. Velocities decrease below that depth to c. 0.5 m/s at 2000 m, probably due

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to frictional interaction with the sea floor. As with Makassar Strait, transport was calculated by multiplying the velocity by the depth-dependent cross section of the channel, which decreases from 36 km at 1000 m depth width to only 10 km wide at 1750 m. The resulting transport ranges from 0.8 Sv towards the Pacific Ocean to 5.2 Sv towards the Indian Ocean, with an average of 2.7 Sv towards the Indian Ocean (Fig. 7, second panel). This is comprised of a 2.5 Sv flow towards the Indian Ocean below 1250 m, and a return flow of c. 0.9– 1.3 Sv towards the Pacific Ocean above 1250 m. The TWT of the flow through the Lifamatola Passage below 1250 m is 3.2 8C. It is expected to be much lower than the TWT of the flow through Makassar Strait due to the much greater depth, and therefore the colder temperature, of the flow. Combining the TWTs of Makassar Strait and Lifamatola Passage during non-contemporaneous measurements suggests an overall TWT of 12.2 8C. Lombok Strait. Lombok Strait is the westernmost path through which ITF water can enter the Indian Ocean. It is 35 km wide and 300 m deep, in contrast with the numerous nearby straits which are shallower than 50 m and thus do not appreciably contribute to the ITF. Previous observational data suggested a total transport of 1.7 Sv (Murray & Arief 1988), with brief periods of northward transport. This was confirmed in a 1996–1997 study using shallow pressure gauges, which found an upper layer (0 to 100 m) that was primarily directed towards the Indian Ocean but periodically reversed and flowed towards the Pacific Ocean. A deeper layer (100 m to 200 m) consistently transported water to the Indian Ocean, with no flow beneath that. The average transport from that study was 2.6 Sv (Hautala et al. 2001). During INSTANT, two moorings were placed in the strait. The eastern mooring was at a depth of 1144 m and remained in place from January 2004 through December 2006, and the western mooring was at a depth of 921 m and remained in place from January 2004 through June of 2005. The results are analysed by Sprintall et al. (2009). The western side of Lombok Strait showed slightly higher velocities than did the eastern side, but the two were highly positively correlated. The flow through Lombok Strait (Fig. 6, third panel) is primarily surface intensified with a maximum at 50 to 80 m with velocities of c. 0.6 m/s. Velocity decreases quickly to the western sidewall and more slowly to the eastern sidewall and weakens considerably below 150 m. The flow was towards the Indian Ocean, except for weak surface reversals during the winter monsoon. The maximum surface velocity occurs in June or July, with a subsurface velocity maximum in August. The average transport

(Fig. 7, third panel) is 2.6 Sv, with a TWT of 21.5 8C. This very warm outflow temperature is due to the shallow sill in Lombok Strait and the surface intensified flow. The TWT is at a maximum during the summer monsoon and at a minimum during the winter monsoon, when the surface flow is reversed. Ombai Strait. Ombai Strait, 35 km wide and 3250 m deep, is located to the east of Lombok Strait. An average transport of 5 Sv was calculated from a one-year current mooring during JADE (Molcard et al. 2001). That study found an upper layer flow in the top 200 m with mean velocities of 0.6 m/s and some as high as 1.5 m/s. There is a deeper layer flow from 200 to 400 m of 0.2 m/s, and the flow decreases to nearly zero by 1200 m. The transport maximum was found during the summer monsoon and minimum during the winter monsoon, which included a reversal in the surface flow. Two moorings were placed in Ombai Strait for the INSTANT observational period and were analysed by Sprintall et al. (2009). The northern mooring was located at a depth of 1329 m and was in place from January 2004 through December 2006, while the southern mooring was at a depth of 3224 m from August 2003 through December 2006 (the same location as the Molcard et al. (2001) mooring described in the previous paragraph). Although these sites are less than 15 km apart, the southern mooring recorded stronger flow. Despite its sill depth of 1450 m, the outflow of Ombai Strait is limited by the c. 900 m Sumba Strait and the c. 1150 Savu/Dao Strait. Ombai Strait (Fig. 6, fourth panel) shows two distinct velocity maxima of c. 0.45 m/s, one near the surface and another near 150 m, in the thermocline, both located slightly towards the western side of the channel. In the subsurface flow is directed away from the Indian Ocean with flows of 0.1 to nearly 0.2 m/s which extend down to 80 m. The flow in the top 300 m of Ombai Strait displays variability similar to that of Lombok Strait, with weak surface reversals during the winter monsoon and velocity maxima during the summer monsoon. Below 300 m, the flow reverses towards the Indian Ocean semi-annually, during the transitions between monsoons. Transport through Ombai Strait (Fig. 7, fourth panel) has a total average transport of 4.9 Sv, with maxima in August and February and minima in May and November. The average TWT is 15.2 8C, but this value increases during the semiannual flow reversal and decreases when all of the transport is directed towards the Indian Ocean. Timor passage. The widest of the outflow channels, Timor Passage is located the furthest south and transports the largest amount of ITF outflow

INDONESIAN THROUGHFLOW OCEANOGRAPHY

water. It is 160 km wide, with depths of 1250 m in the east to 1890 m in the west. Two current meters were deployed there during JADE and a mean transport of 4.3 Sv was found, more than half of it in the upper 200 m (Molcard et al. 1996). Variability was found on annual, semi-annual, and monthly timescales, with a maximum in spring and summer and minimum in winter. Timor Passage was monitored by four moorings during INSTANT, again analysed by Sprintall et al. (2009). From NW to SE, the moorings were located at: 741 m depth (from January 2004 to December 2006), 1890 m depth (from December 2003 to December 2006), 1386 m depth (from December 2003 to December 2006), and 902 m depth (from December 2003 to December 2006). Flow through Timor passage (Fig. 6, bottom panel) is nearly always towards the Indian Ocean, with reversals at the northernmost mooring suggesting the presence of rotating eddies. The maximum velocity is located at the surface with flows of 0.5 m/s. Although most of the flow is found in the upper 200 m, there is a weak secondary core of increased velocity near 1200 m, with speeds up to 0.1 m/s. Transport within Timor Passage (Fig. 7, bottom panel) does not co-vary with Lombok and Ombai Straits. Reversals are seen during the monsoon transitional periods in the flow between c. 800 to c. 1800 m, with the strongest reversals below 1400 m. The mean transport is 7.5 Sv, or half of the total outflow, with variability dominated by the annual cycle. The large difference in transport between JADE and INSTANT is due to a difference in the width of the strait used to calculate transport. The earlier study assumed that the ITF encompassed 85 km of channel width, but the INSTANT measurements, using four moorings instead of two, showed that the ITF has significant velocity through the full 160 km width of the channel. The associated TWT is 17.8 8C, which is warmer than Ombai Strait but colder than Lombok Strait. The total outflow is dominated by Timor Passage transport. During INSTANT, the total outflow value was 15 Sv, which is larger than previous estimates. Most of that transport was confined to the upper 300 m.

Conclusions and future work The ITF has now been considered on global, local, and regional scales. It sits at the intersection of ENSO, IOD, and the SE Asian monsoon and provides a conduit from the Pacific Ocean to the Indian Ocean. Broadly speaking, the ITF is controlled by a pressure gradient in the thermocline and wind forcing in the surface layer, with a deep transport presumably driven by pressure as well. But the existence of the pressure gradient between the Pacific and Indian Oceans can be seen to be

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the result of a geographical configuration of landmasses in which the Indian Ocean lacks a northern basin to export heat. Luyten et al. (1983) presented a theoretical model of circulation proposing that the ocean can be divided into multiple layers of constant density. The uppermost layer, located in the tropics is driven directly by Ekman pumping, while the deeper layers interact with the atmosphere only where they outcrop at higher latitudes. The Indian Ocean is therefore ventilated by the northern Pacific Ocean since it has no northern basin of its own. The Indian Ocean remains at a steady temperature because its high evaporation rate leads to cooling through latent heat loss. Questions remain about the ITF on all of these scales. The INSTANT program and the growing availability of high quality observational data will allow models with realistic values for the mass and heat transport of the ITF to answer many of these questions. They will also allow for a comprehensive understanding of the flow of heat, freshwater, and mass between ocean basins. Within the observational data, however, there remains significant uncertainty about the shallow flow through Lifamatola Strait and the contribution and importance of the SCST to the ITF. Although it is understood how the monsoon and ENSO affect the ITF, it is still not known if there are feedbacks by which the ITF can influence those phenomena. A growing appreciation of the IOD as an independent climate phenomenon offers another avenue of research, as the relationship between the IOD and the ITF is not well documented. A deeper understanding of the ITF can be expected to provide the basis for better regional forecasts and climate prediction. It will also contribute to our understanding of the ITF in previous geological and climatic conditions. The INSTANT data analysis is funded by National Science Foundation Grant OCE 07-25935. I thank Arnold L. Gordon, Xiaojun Yuan, Robert Hall, Janet Sprintall and Derek Vance for helpful reviews.

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Indonesian Throughflow variability during the last 140 ka: the Timor Sea outflow ANN HOLBOURN1*, WOLFGANG KUHNT1 & JIAN XU1,2 1

Institute of Geosciences, Christian-Albrechts-University, D-24118 Kiel, Germany 2

State Key Laboratory of Continental Dynamics and Department of Geology, Northwest University, 229 North Taibai Rd., Xi’an 710069, China *Corresponding author (e-mail: [email protected])

Abstract: The transfer of surface and intermediate water from the Pacific to Indian Ocean through the Indonesian passages (Indonesian Throughflow: ITF) strongly influences the heat and freshwater budgets of tropical water masses, in turn affecting global climate. Here, we use combined d18O and Mg/Ca analyses of surface and thermocline planktonic foraminifers to estimate variations in sea surface temperature, salinity and mixed layer thickness over the last 140 ka. Comparison of water mass properties reveals a steeper thermocline temperature gradient in the Timor Strait than in the eastern Indian Ocean during glacials, implying a decrease in ITF cool thermocline outflow. A major freshening and cooling of thermocline waters occurred at c. 9.5 ka, when sea level rose above a critical threshold, allowing establishment of a shallow marine connection from the South China Sea to the Java Sea. Comparison of benthic d13C profiles (c. 1800 to 3000 m water depth) suggests vigorous mixing of Indian Ocean and ITF outflow intermediate waters during interglacials. In contrast, deep and intermediate water masses became more stratified during glacials. Lower d13C values at c. 3000 m water depth reflect a decrease in deepwater ventilation, probably related to slowdown of the global thermohaline circulation during glacials.

Regional oceanography The Indonesian Throughflow (ITF) transports c. 12 –16 Sv of surface and intermediate waters from the Pacific to the Indian Ocean through the Indonesian seas (Gordon 2005). Recent international oceanographic monitoring projects (ARLINDO and INSTANT, Sprintall et al. 2004; Gordon 2005; van Aken et al. 2009) revealed that a dominant component of the ITF inflow is relatively cool and fresh thermocline water originating from the North Pacific. Today, most of the ITF inflow water enters the Indonesian Seas by crossing the sill south of Mindanao, then travels through the Makassar Strait. The ITF inflow is additionally cooled en route by seasonal freshwater lenses within its pathway, which effectively block the warm surface component during the boreal winter monsoon (Gordon et al. 2003). As a result, the main transport of the ITF occurs not at the surface but within the thermocline. Thus, the ITF outflow into the Indian Ocean through the three main exit passages (Lombok, Ombai and Timor Straits) does, overall, cool rather than warm the tropical Indian Ocean. Approximately half of the ITF transport entering the Indian Ocean takes place through the Timor Strait (Sprintall et al. 2009). The Timor Sea and eastern Indian Ocean are additionally influenced by warm tropical surface

water masses from the northern part of the NW Australian shelf, partly feeding the South Equatorial Current and the Leeuwin Current, which continue westward at c. 128S and southward along the West Australian coast. These warm surface currents are mainly active during austral autumn and winter (April –September), when easterly trade winds prevail (Godfrey & Weaver 1991; Tomczak & Godfrey 1994; Peter et al. 2005). Tropical surface water masses clearly dominate the Timor Sea, resulting in warm (c. 29.5 8C and c. 27.5 8C), low-salinity (c. 34.1 psu and c. 34.4 psu; psu, practical salinity unit) surface water and a deep thermocline (¼18 8C isotherm, 170 m) [May 2001 and September 2005 CTD (conductivity, temperature, density) data given by Bassinot et al. 2002; Kuhnt et al. 2005]. While sea surface temperature (SST) decreases and sea surface salinity (SSS) increases from NE to SW in the Timor Sea, thermocline water masses exhibit a reversed latitudinal temperature and salinity gradient (Fig. 1). Coolest and freshest thermocline waters (at c. 150 m water depth) occur along the southern margin of the Indonesian islands. Toward the East, these waters are directly derived from the cool and fresh ITF, whereas toward the West (off Sumatra), there is additional influence of wind-driven upwelling of deeper thermocline water. These cool, fresh upper thermocline waters become warmer and saltier

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 283–303. DOI: 10.1144/SP355.14 0305-8719/11/$15.00 # The Geological Society of London 2011.

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(a)

10°N 28 WPWP

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Karimata Str.

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24 MD01-2378

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West Australian Current

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Fig. 1. (a) Annual average SST and (b) thermocline (c. 200 m water depth) temperatures in the Timor Sea (WOA 2005 data, Locarnini et al. 2006). Note reverse NE–SW gradient in SST and thermocline temperatures.

towards the south, when mixing with Indian Ocean intermediate water occurs. While the Makassar Strait provides the main route for the shallow throughflow of waters from the Pacific to the Indian Ocean, the shallow sill depth of c. 680 m at its southern end prevents a deep throughflow from entering the Indonesian Seas. The only entrance allowing flushing of the deep basins within the Indonesian archipelago is the Lifamatola Passage east of Sulawesi with a sill depth of c. 2000 m (van Riel 1956; Broecker et al. 1986). Initial tracer measurements showed that ventilation by a deep overflow from the Pacific Ocean occurred through this passage (van Aken et al. 1988; Gordon et al. 2003), although measurements of flow variability and mass transport were scarce and rather uncertain. However, recent

measurements from a current meter mooring deployed across the Lifamatola Passage revealed that a substantial inflow (2.5 Sv) of cool, deepwater (3.2 8C) occurs from the Maluku Sea into the Seram Sea at depths below c. 1250 m (van Aken et al. 2009). This deep inflow eventually travels though the Banda Sea, joining the main shallow inflow from the Makassar Strait. Thus, the ITF outflow into the Indian Ocean consists of a shallow- and deep-water component originally derived from Pacific thermocline and intermediate water masses.

Approach and objectives Previous studies of Late Pleistocene palaeoceanography and palaeoclimate within the ITF outflow

ITF VARIABILITY

area in the Timor Sea focused on four major topics. (1) Palaeoproductivity reconstructions based on opal, organic matter and chlorin fluxes and benthic foraminiferal assemblages, revealed marked precessional variability as well as a strong glacial –interglacial contrast, which were linked to changes in monsoonal wind and oceanic circulation patterns (Mu¨ller & Opdyke 2000; Holbourn et al. 2005; Murgese & De Deckker 2005, 2007; Kawamura et al. 2006; Du¨rkop et al. 2008). (2) Land derived pollen records allowed reconstruction of regional climate variability, which appeared to be mainly controlled by variations in sea level and monsoonal intensity (van der Kaars et al. 2000; van der Kaars & De Deckker 2002, 2003). (3) Studies of clay minerals, marine v. terrestrial organic matter and microfossil variations provided estimates of past sea level offshore Australia and sedimentation regimes in the Timor Sea (Gingele et al. 2002; Moreno et al. 2008; De Deckker & Yokohama 2009). (4) Surface salinity temperature reconstructions were initially based on planktonic foraminiferal counts and d18O (Martinez et al. 1997, 1998, 1999; Spooner et al. 2005). The new planktonic foraminiferal Mg/Ca derived temperature proxy has so far been applied to reconstruct temperature and seawater d18O variability over the late Pleistocene and Holocene in the central part of the West Pacific Warm Pool (Lea et al. 2000; Visser et al. 2003; Stott et al. 2007; Oppo et al. 2009) and in distal ITF outflow locations in the Eastern Indian Ocean (Xu et al. 2006, 2008; Zuraida et al. 2009; Mohtadi et al. 2010). Although these studies provided new insights into local hydrological variability, they could not fully resolve regional shifts in water mass boundaries and inherent fluctuations in ITF intensity and thermal structure. The frontal area between ITF derived cool, fresh thermocline water and relatively warm and salty Indian Ocean thermocline water in the Timor Sea offers a unique opportunity to reconstruct ITF outflow variability over the last 140 ka. Our approach is to reconstruct the gradient between surface and upper thermocline temperatures and salinities in one proximal area with strong thermocline flow (Timor Strait) and a contrasting distal area, where ITF waters are mixing with eastern Indian Ocean waters at the southern edge of the main ITF outflow. During periods of decreased ITF thermocline flow, the eastern Indian Ocean will be more influenced by warm and salty tropical Indian Ocean waters and the upper ocean thermal gradient will be decreased. We additionally use benthic d18O and d13C records at three sensitive locations within the Timor Sea to investigate changes in intermediate and deepwater masses possibly related to variability of the deep ITF outflow. The main objectives of this work are thus: (1) to

285

provide a synthesis of surface and upper thermocline stable isotope and Mg/Ca derived temperature data as well as benthic stable isotope records from sensitive locations in the Timor Sea to monitor variations in the shallow and deep outflow components of the ITF over the last two glacial cycles (2) to investigate major controls on ITF variability during the late Pleistocene, in particular relations to sea level, trade-monsoon wind systems and insolation forcing.

Material and methods Sonne-185 (VITAL) and IMAGES MD122 (WEPAMA) cruises The Sonne-185 ‘VITAL’ Cruise from Darwin (Australia) to Jakarta (Indonesia) from 15th September to 6th October 2005 retrieved sediment and water samples at 58 stations along two main transects across the Timor Strait between the Arafura Sea and East Timor and across the outflow of the ITF into the Indian Ocean between the Sunda Islands and NW Australia (Fig. 2). The onboard CTD with water sampler was deployed at six stations, a multicorer equipped with a selfregistering CTD to obtain additional CTD records of the water profile and bottom water was deployed at 54 stations. For this study we selected four cores from key locations and integrated new isotope and Mg/Ca data with published data from the Timor Sea (Table 1). SONNE-185 Piston-Core 18460 (8847.3860 S, 128838.4850 E; 1875 m water depth) was recovered at the tip of an elevated promontory at the northern flank of the Timor Trough, east of eastern Timor underneath the main flow of the ITF in the Timor Strait (Kuhnt et al. 2005). The sediment consists of undisturbed, homogenous, carbonate-rich nannoplankton ooze with high clay content, abundant planktonic foraminifers and relatively rare benthic foraminifers. Sediment samples (c. 30 –40 cc from 1 cm thick sediment slices) were analysed at 10 cm intervals between 0 and 1470 cm (0–142 ka) corresponding approximately to millennial resolution. Additional centennial resolution sample sets were analysed over the last glacial termination. SONNE-185 Kastenlot-Core SO18462 (9805.3170 S, 129814.2240 E, 1421 m water depth) was retrieved in the lower part of the Arafura slope transect at the southern margin of the Timor Strait. This core is situated south of the main path of the ITF in the Timor Strait and therefore represents the northernmost part of the warm water body that covers the eastern Arafura shelf. The sediment composition is similar to that in Core SO18460. Samples were analysed in 1 cm

286 A. HOLBOURN ET AL. Fig. 2. ITF Outflow area with RV Sonne 185 VITAL stations and Core MD01-2378 location. Cores 18460, 18462 and CTD stations 18458 and 18465 are located within the ITF outflow in the Timor Strait, whereas Cores MD01-2378, 18506 and Core and CTD stations 18479 are within the southern frontal area of the ITF outflow and mixing zone with Indian Ocean thermocline waters. CTD stations and cores used in this study are circled, black dots indicate stations used for core-top proxy calibration.

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287

Table 1. New and published stable isotope and Mg/Ca data from the Timor Sea integrated in this work Data

Core/location

Time interval (ka)

Reference

Benthic stable isotopes

MD01-2378

0 – 150 ka

Benthic stable isotopes Benthic stable isotopes Benthic stable isotopes Planktonic stable isotopes

S018460 SO18479 Timor Sea MD01-2378

0 – 128 ka 0 – 140 ka core tops 0 – 150 ka

Planktonic stable isotopes Planktonic stable isotopes Mg/Ca

S018460 S018462 MD01-2378

0 – 128 ka 0 – 20 ka 0 – 150 ka

Mg/Ca Mg/Ca

SO18460 S018462

0 – 128 ka 0 – 20 ka

Holbourn et al. (2005) Du¨rkop et al. (2008) This study This study This study Holbourn et al. (2005) Du¨rkop et al. (2008) Xu et al. (2006, 2008) This study This study Xu et al. (2006, 2008) Zuraida et al. (2009) This study This study

resolution for intervals spanning the last glacial maximum, Termination I and the Holocene. SONNE-185 Piston-Core SO18479 (12827.1590 S, 121822.3950 E, 2974 m water depth) is located in the eastern Indian Ocean, and represents the northernmost and deepest site along the SW Ashmore Reef transect. This core has the highest sedimentation rates of all cores studied (c. 15 cm/ka) due to increased clay accumulation. Sediment samples were analysed in 10 cm (representing c. 0.7 ka) resolution throughout the core. IMAGES Calypso-Core MD01-2378 (1384.950 S, 121847.270 E; 1783 m water depth) was recovered at the northwestern margin of the Scott Plateau at the southern margin of the main outflow of the ITF in the Timor Sea during the IMAGES MD122 (WEPAMA) Cruise in May 2001. The sediment consists of undisturbed, homogenous, carbonaterich nannoplankton ooze with abundant planktonic foraminifers and relatively rare benthic foraminifers. Sediment samples (c. 30 –40 cc from 1 cm thick sediment slices) were analysed at 10 cm intervals between 0 and 1470 cm (0–142 ka) in approximately millennial resolution (Holbourn et al. 2005). Additional centennial resolution data sets were analysed between 0 and 450 cm (Holocene and Termination I) at 2 cm resolution (Xu et al. 2008), between 450 and 895 cm (MIS 3) at 1 cm resolution (Du¨rkop et al. 2008; Zuraida et al. 2009) and between 1310 and 1471 cm (MIS 5e and Termination II) at 1 cm resolution (Xu et al. 2006, 2008).

Stable isotopes Samples were dried in an oven below 40 8C, disaggregated by soaking in water, then wet sieved over a 63 mm screen. Residues were dried on a sheet of filter paper below 40 8C, then sieved into 63– 150 mm, 150 –250 mm, 250– 630 mm fractions.

For planktonic isotope analysis, we selected 20 tests of Globigerinoides ruber (white) and Pulleniatina obliquiloculata from the size fraction of 250– 315 mm. Globigerinoides ruber is a well-known surface dweller, whereas P. obliquiloculata has a preferred depth habitat within the upper thermocline (Xu et al. 2006; Cle´roux et al. 2007; Farmer et al. 2007; Mohtadi et al. 2009). A study of 33 coretop samples from the Timor Sea indicates that P. obliquiloculata Mg/Ca-based temperatures are generally comparable with WOA05 annual mean temperatures at water depth of 100–125 m (Locarnini et al. 2006), supporting a thermocline depth habitat for this species (Zuraida et al. 2009). For benthic isotope analysis, we selected 3 to 6 tests (.250 mm) of the epifaunal benthic foraminifer Planulina wuellerstorfi. In a few samples, where benthic foraminiferal density was low, a smaller number (1–2) of specimens was analysed. All tests were checked for cement encrustations and infillings before being broken into large fragments, then cleaned in alcohol in an ultrasonic bath and dried at 40 8C. Stable carbon and oxygen isotope measurements were made with the Finnigan MAT 251 mass spectrometer at the Leibniz Laboratory, Kiel University. The instrument is coupled on-line to a Carbo-Kiel Device (Type I) for automated CO2 preparation from carbonate samples for isotopic analysis. Samples were reacted by individual acid addition. The mean external error and reproducibility (1s) of carbonate standards is better than +0.07‰ and +0.05‰ for d18O and d13C, respectively. Results were calibrated using the NIST (National Institute of Standards and Technology, Gaithersburg, Maryland) carbonate isotope standard and NBS (National Bureau of Standard) 20 and in addition NBS 19 and 18, and are reported on the PeeDee Belemnite (PDB) scale. Replicate samples of G. ruber and

288

A. HOLBOURN ET AL.

P. obliquiloculata indicate that the mean reproducibility (1s) is +0.11‰ for d18O and +0.13‰ for d13C. The mean reproducibility of paired samples of P. wuellerstorfi is better than +0.08‰ for d18O and d13C (see also replicate measurements of sample sets published in Holbourn et al. 2005; Xu et al. 2006; Du¨rkop et al. 2008).

Mg/Ca palaeothermometry Mg/Ca ratios were measured on c. 30 tests of G. ruber and P. obliquiloculata, from the same size fraction used for stable isotope analysis. To assess overall reproducibility, we duplicated measurements of 24 randomly selected samples. Foraminiferal tests, weighing about 0.2–0.8 mg per sample, were gently crushed under the microscope, and cleaned of contaminant phases using the standard cleaning procedure with reductive step (Martin & Lea 2002). Samples were analysed on an ICP-OES (inductively coupled plasma – optical emission spectrometer) (Spectro Ciros SOP) with cooled cyclonic spraychamber and microconcentric nebulization (200 ml min21) at the Institute of Geosciences, Kiel University. Intensity ratio calibration followed the method of de Villiers et al. (2002). Internal analytical precision from replicate measurements is better than 0.1– 0.2% (relative standard deviation), which corresponds to +0.02 8C. Replicate analyses showed a standard deviation of 0.14 mmol/mol, equivalent to +0.7 8C. Consistency of results was checked by analysing sets of standards obtained from M. Greaves, University of Cambridge. The validity of Mg/Ca ratio was checked by evaluating the consistency of Ca concentration before and after cleaning. Samples with a reduction in Ca concentration of more than 20% were rejected. Fe/Ca, Al/Ca and Mn/Ca ratios were additionally used to monitor cleaning efficacy, and samples with a significant correlation between Fe/Ca, Al/Ca, Mn/Ca and Mg/Ca values were excluded, following the method used by Schmidt et al. (2004). SST was calculated from the Mg/Ca ratios of G. ruber, using the equation developed by Anand et al. (2003): Mg/Ca ¼ 0.38 (+0.02) exp 0.090 (+0.003) T, where a change in Mg/Ca of 9.0 + 0.3% is equivalent to a change of 1 8C in temperature, based on calibration of Mg/Ca in 12 species from sediment traps in the Sargasso Sea to d18O derived temperatures. Following this approach, the accuracy in estimating calcification temperature is +1.2 8C. Thermocline water temperature derived from P. obliquiloculata Mg/Ca was based on a species-specific calibrated equation assuming an exponential constant of 0.09 (Anand et al. 2003). The equation is expressed as Mg/Ca ¼ 0.328 (+0.007) exp 0.090 (+0.003) T. No correction

was applied for the reductive step in our cleaning protocol, which generally results in a small decrease in Mg/Ca, leading to temperature underestimation of up to c. 0.6 8C (Rosenthal et al. 2004).

Salinity reconstructions We calculated surface and thermocline water d18O composition (d18Osw) from paired Mg/Ca and d18O measurements of G. ruber and P. obliquiloculata, respectively. We used the palaeotemperature equation of Bemis et al. (1998) and Thunell et al. (1999) expressed as: d18 Osw (V-SMOW) ¼ 0:27 þ (T(8C)  16:5 þ 4:8  d18 Ocalcite (V-PDB))=4:8: Since d18Osw is corrected for the effect of temperature on the oxygen isotope fractionation between the foraminiferal test and water, d18Osw is only influenced by the global change in d18O related to continental ice volume and local d18O variations related to the salinity of surface and thermocline water masses. As the d18Osw curves (bottom, thermocline and surface waters) carry the same ice volume effect, deviations mainly reflect changes in salinity. The present day salinity/seawater d18O relationship in the Timor outflow area is close to the average low latitude value of c. 0.3‰ d18O/psu (Ru¨hlemann et al. 1999) and thermocline d18Osw and salinity variability in R/V Sonne Cruise 185 CTD-measurements from the outflow area is in the order of 0.2‰ d18O and 0.6 psu, respectively (Fig. 3).

Results Chronology Age models for the Holocene and Termination I in Cores MD01-2378, SO18460, SO18462 are based on AMS (accelerator mass spectrometry)14C dates of the surface dwelling G. ruber. AMS14C ages of Core MD01-2378 are from Holbourn et al. (2005) and Xu et al. (2008); AMS14C ages of Cores SO18460 and SO18462 are from Xu et al. (2010). Conventional AMS14C ages were converted into calendar ages following Fairbanks et al. (2005) and a reservoir correction of 300 years (applicable for the Timor Sea following Butzin et al. 2005) was applied. Age models for the interval 18.15 ka to 62 ka (MIS 2 and MIS 3) are based on the correlation of benthic oxygen isotope records in Cores MD012378, SO18460 and SO18479 to the d18O record of the EDML ice core, recovered from Dronning Maud Land within the European Project for Ice Coring in Antarctica (EPICA Community members

ITF VARIABILITY

289

Fig. 3. CTD temperature-salinity profiles at Stations 18458 and 18465 within the main ITF outflow into the Timor Sea, and at Station 18479 in the eastern Indian Ocean (see locations in Fig. 2). Measurements were taken during the inter-monsoonal season of September/October 2005, when ITF surface outflow was at a maximum (unpublished data from Sonne 185 VITAL cruise). Approximate depth habitat of P. obliquiloculata is shaded blue. Rectangles indicate temperature ranges for P. obliquiloculata and G. ruber in six core tops from the Scott Plateau (Xu et al. 2006). Note temperature difference of 3 –6 8C and salinity difference of 0.2– 0.3 psu between ITF and Indian Ocean water masses at habitat depth of P. obliquiloculata. Decrease in thermocline temperature is steep between 90 and 110 m water depth within the ITF outflow and more gradual between 120 and 1.50 m water depth in the eastern Indian Ocean.

2004). This ice core has been synchronized to the layer-counted NGRIP ice core following the new Greenland Ice Core Chronology (GICC05) timescale (Andersen et al. 2006; EPICA Community members 2006). Eight tie points were used for correlation to the GICC05 chronology over the interval 18.15 ka to 62 ka (Table 2). The interval older than 62 ka was correlated to the POP-Chronology, based on the benthic d18O record from Core MD95-2042 (Martrat et al. 2007). Four additional correlation points with age assignments following Shackleton et al. (2002, 2003) were used to constrain the chronology over the MIS 5.5 interval (see details in Xu et al. 2006).

G. ruber d18O and Mg/Ca to reconstruct surface water characteristics SST estimates from Core MD01-2378 (eastern Indian Ocean) and Core SO118460 (Timor Strait) spanning the last two glacial cycles show a high degree of similarity with glacial minimum values of c. 24 –25 8C and interglacial maxima of c. 28– 29 8C (Fig. 4). Slightly warmer surface

temperatures prevailed in the eastern Indian Ocean, except during MIS (Marine Isotope Stage) 5.3 and the late Holocene, when surface waters were somewhat warmer in the Timor Strait. Today these two sites are characterized by almost identical annual average SST of c. 28.2 –28.4 8C (Fig. 1). Calcification temperature corrected surface water d18Osw values are also close in the Timor Strait (SO18460) and eastern Indian Ocean (MD01-2378). Late Holocene d18Osw varies between 20.2 and 20.3‰ at both sites, which is c. 0.2‰ less than surface water d18Osw measured in late September –early October (unpublished Sonne-185 CTD data). This slight offset may be explained by the seasonal variability of SSS, which is higher at the end of the dry season, when water samples were taken during the Sonne-185 cruise. The offset may also be due to the fact that G. ruberbased d18Osw mainly records austral summer hydrographic conditions in this area. The glacial– interglacial d18Osw gradient is c. 1.3–1.5‰ from the Last Glacial Maximum (LGM) to the late Holocene, which cannot be fully accounted for by the overall change in global ice volume. The residual 0.1 –0.3‰ indicates elevated regional glacial SSS of c. 0.3–1 psu at both sites.

290

Table 2. Age tie points in Cores SO18460, SO18479, MD01-2378 Age (ka)

SO18479 Depth (cm)

MD01-2378 Depth (cm)

Description

References

268 414 – 535 – 573 – 678 733 810 940 960 1010 1110 1150 1240 1390 1460 1530 1570 –

260 360 400 500 620 670 760 810 – 1090 1260 – – 1400 1420 1440 1560 1600 1660 1690 1720

342 506 554 – 642 664 742 – 817 901 971 991 1021 1081 – 1161 1281 1326 1383 1410 1431

d18O maximum prior to deglaciation d18O maximum after AIM 4 event d18O maximum before AIM 4 event d18O maximum after AIM 8 (¼A1) event Last d18O minimum of AIM 8 (¼A1) event d18O maximum before AIM 8 (¼A1) event d18O minimum in centre of AIM 12 (¼A2) event d18O maximum before AIM 12 (¼A2) event d18O maximum before DO 14 event Last MIS 4 d18O maximum First MIS 4 d18O maximum Third MIS 5.1 d18O minimum Second MIS 5.1 d18O minimum First MIS 5.1 d18O minimum Last MIS 5.3 d18O maximum Last MIS 5.3 d18O minimum First MIS 5.4 d18O maximum end of MIS 5.5 plateau Beginning of MIS 5.5 plateau Midpoint Termination 2 Last MIS 6d18O maximum

EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) EPICA (GICCO5 timescale) NGRIP (GICCO5 timescale) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD95-2042 (Martrat et al. 2007) MD01-2378 (Xu et al. 2006) MD01-2378 (Xu et al. 2006) MD01-2378 (Xu et al. 2006) MD01-2378 (Xu et al. 2006)

A. HOLBOURN ET AL.

18.15 27.45 30.65 36.75 38.15 39.85 47.25 48.85 54.49 62.00 70.00 74.00 77.00 86.00 88.00 90.50 111.00 116.00 128.00 132.00 135.00

SO18460 Depth (cm)

ITF VARIABILITY

291

Fig. 4. Comparison of SST and sea surface water d18O variations over the last 150 ka in the Timor Sea (Cores SO18460, MD01-2378) and in the centre of the West Pacific Warm Pool (ODP Site 806; data from Lea et al. 2000).

A comparison of Mg/Ca-derived SST and calcification temperature corrected surface water d18Osw (1) in the centre of the ITF outflow within the Timor Strait (SO18460), (2) at the southern margin of the Timor outflow into the eastern Indian Ocean (MD01-2378) and (3) in the centre of the West Pacific Warm Pool (ODP 806; Lea et al. 2000) indicates highest SST and SSS (d18Osw) within the West Pacific Warm Pool, and lowest SST and SSS within the ITF outflow area (Fig. 4). SST in the Timor Strait and eastern Indian Ocean approach values in the centre of West Pacific Warm Pool during warm phases with elevated sea level (MIS 1, MIS 5.1, MIS 5.3 and MIS 5.5). However, differences increase markedly during glacials (up to 3 8C) due to the pronounced glacial –interglacial contrast in Timor Sea SST. In the Timor Strait and eastern Indian Ocean, SSS are consistently lower than in the West Pacific Warm

Pool, particularly during interglacials, when values are c. 0.4 –0.7‰ lower.

Multispecies planktonic d18O and Mg/Ca to reconstruct upper water column variability Longterm variability during the last two glacial cycles. Changes in the ITF vertical profile over the last 140 ka are monitored by comparing the gradient between SST and upper thermocline temperatures in Core SO18460 in the Timor Strait and Core MD01-2378 within the mixing zone between ITF and eastern Indian Ocean water masses (Fig. 5). Thermocline temperatures derived from P. obliquiloculata Mg/Ca are substantially lower (by c. 2– 3 8C) in the Timor Strait during the LGM, MIS 3, MIS 5.2 and MIS 5.4, whereas thermocline water masses remain almost identical at both locations during Termination 1, MIS 5.1, MIS 5.3 and MIS

292

A. HOLBOURN ET AL. 26

(a)

MD01-2378

Thermocline Temperature (°C)

24

22

20

°C

SO18460 18

ΔT surface - thermocline

thermocline shoaling

SO18460

16 8 7

(b)

5

thermocline deepening

MD01-2378

6

4

°C

3 2

Δ SST

18460 - 2378

(c) 0

Δ Thermocline Temperature (°C) 18460 - 2378

(d)

°C

1 0 –1

weak thermocline ITF

–2

°C

–3 –4

MD01-2378 (e)

2.5

δ18Owuellerstorfi

3

SO18460

3.5

(‰ vs. PDB)

strong thermocline ITF

–1

4

0

20

40

60

80

100

120

140

Age (ka) Fig. 5. (a)–(b) Changes in the upper water column temperature gradient in Core MD01-2378, situated in the southwestern Timor Sea within the mixing zone with Indian Ocean thermocline waters, and Core SO18460, located within the ITF outflow in the Timor Strait, over the last 140 ka. (c) – (d) Differences in surface and thermocline temperatures between Cores MD01-2378 and SO18460. Major divergence in thermocline temperatures during MIS 2-4, MIS 5.2 and MIS 5.4 (light brown shading) suggests weakened Australian monsoon and intensified SE trades. During these intervals, thermocline outflow does not influence the distal eastern Indian Ocean Site MD01-2378. Present day thermocline temperature is c. 1 8C cooler at Site SO18460. (e) Benthic d18O for stratigraphic reference in Cores MD01-2378 and SO18460.

ITF VARIABILITY

5.5. The temperature difference between surface and thermocline waters (DT ) in Cores SO18460 and MD01-2378 provides a proxy for thermocline depth during the last two glacial cycles (Fig. 5). Increases in DT during the LGM, MIS 3, MIS 5.2 and MIS 5.4 imply that the thermocline became shallower in the Timor Strait during these intervals. These divergences indicate significant changes in the upper water column structure and thermocline depth gradient in the Timor Strait and eastern Indian Ocean over the last two glacial cycles. ITF response to rising sea level during termination I. Figure 6 shows that thermocline d18Osw decreased in parallel to sea-surface d18Osw from the LGM (approx. 21 –18 ka) to earliest Holocene c. 10 ka) in the Timor Strait. This decrease by c. 0.6‰ can be accounted for by the global sea level rise of c. 60– 70 m. However, a remarkable decoupling between surface and thermocline d18Osw occurs at c. 9.5 ka, as thermocline d18Osw rapidly drops by 0.5–0.7‰ and surface d18Osw remains nearly constant or decreases slightly by 0.1–0.2‰. This rapid change in thermocline d18Osw reflects a substantial freshening of upper thermocline water masses, which cannot be related to changes in local precipitation. The concurrent abrupt increase in the thermal gradient between surface and upper thermocline by 3 –4 8C is mainly driven by thermocline cooling and/or shoaling. These changes occurred at a time, when global sea level rose above a critical threshold of c. 40 m (Lambeck & Chappell 2001), which corresponds to the sill depth of the Karimata Strait, providing the main connection between the South China Sea and Java Sea (Fig. 6).

Benthic d18O and d13C to track deepwater variability Modern d18O and d13C in the Timor Sea. We investigated present day benthic d18O and d13C variability in multicore tops along two depth transects in the Timor Strait and eastern Indian Ocean (Figs 2 & 7). Along both transects, d18O values vary between c. 0.7 and 2.8‰, exhibiting a consistent increasing trend with water depth. d13C values overall vary from c. 0.8 to 0.3‰ along both transects, showing a coherent decreasing trend with depth but only for stations between c. 500 and 1500 m water depth. At depth .1500 m, d13C values show a wide scatter, reaching c. 0.5 –0.6‰ for stations between c. 1500 and 3000 m in the western Timor Sea and c. 0.3–0.4‰ for stations at c. 2000 m in the Timor Strait. This geographical partition suggests that deepwater masses in the Timor Strait and western Timor Sea become increasingly differentiated at depth .1500 m. Southern-sourced

293

Indian Ocean intermediate (IOCW) waters are characterized by higher d13C values between c. 0.6 and 1.1‰, as shown by the World Ocean Circulation Experiment (WOCE) transect I10 in the eastern Indian Ocean. In contrast, the ITF deep outflow from the Banda Sea exhibits lower d13C values, as it is relatively depleted in 13C due to enhanced productivity in the Banda Sea (Gordon 2005). Thus, the differences in d13C values below 1500 m appear to reflect the relative influence of these different water masses with higher values along the eastern Indian Ocean transect, where IOCW is more influential, and lower values along the Timor Strait transect, where the ITF deep outflow is dominant. Stations above 1500 m water depth are predominantly influenced by the ITF deep outflow, irrespective of their location within the Timor Sea.

d18O and d13C variability over the last 140 ka. We monitored benthic d18O and d13C variability in Core SO18460 (1883 m water depth) situated within the main ITF path in the Timor Strait and Cores MD01-2378 (1783 m water depth) and SO18479 (2983 m water depth) located in the eastern Indian Ocean at more distal positions from the ITF outflow (Fig. 8). The d18O profiles in Cores SO18460 and MD01-2378 are almost identical, whereas values are generally 0.3– 04‰ lower in Core SO18479, reflecting differences in water mass density at this deeper site. In contrast, d13C records exhibit distinct patterns of variability over the last two glacial cycles: (1) d13C values converge in all three cores during warmer intervals (Holocene, MIS 5.5), implying a consistent deepwater mass throughout the Timor Sea during these periods; (2) d13C values are quite similar in Cores SO18460 and SO18479 during relatively cold intervals (later part of MIS 3, MIS 5.2 and MIS 5.4), but significantly higher (by 0.3–05‰) in Core MD01-2378, suggesting a different water mass prevailed in the southwestern Timor Sea during these intervals; (3) d13C values diverge in all three sites during glacial periods (MIS 2, later part of MIS 3, MIS 4 and MIS 6), implying that deep and intermediate water masses substantially differ in different areas of the Timor Sea.

Discussion Variability of ITF surface-thermocline outflow Seasonal controls on ITF outflow: a model for longterm change? The modern ITF outflow is dominated by a thermocline flow of cooler and fresher water originating from the North Pacific with some advection of South China Sea surface and South Pacific intermediate waters in the Java and Banda

1.0

0.5

0

–0.5

–100

SO18462

SO18462

5

0m

0

SO18462

5

freshening

thermocline

10

–50

Sill Depth of Karimata Strait

1.5

294

shoaling

MWP 1a

MWP 1a

15

15 Sealevel (Lambeck & Chappell 2001)

20

20

SO18460

SO18460

SO18460

25 1.5

1.0

0.5

Surface δ18Oseawater (G. ruber)

0 2.0

1.5

1.0

0.5

0

–0.5

Thermocline δ18Oseawater (P. obliquiloculata)

2

4

6

8

25

ΔTSST-Thermocline Temperature (°C)

Fig. 6. ITF response to rising sea level during Termination 1. d18Osw reconstruction for surface waters is based on d18O and Mg/Ca derived temperatures in G. ruber; d18Osw reconstruction for upper thermocline waters is based on d18O and Mg/Ca derived temperatures in P. obliquiloculata. Cores SO18460 and SO18462 are both located within the Timor Strait (see Fig. 2).

A. HOLBOURN ET AL.

Age (ka)

10

Age (ka)

0

ITF VARIABILITY

295

0 thermocline ITF

thermocline ITF

Water depth (m)

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W

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Seas (Gordon 2005). This cool, fresh and nutrient-rich ‘Australasian Mediterranean Water’ (AAMW), characterized by a comparatively shallow thermocline and low salinity gradient, dominates at Site SO18460 within the Timor Strait. In contrast, the upper water column differs significantly at the more distal Site MD01-2378, located within the mixing zone between AAMW and Indian Ocean waters. Site MD01-2378 is influenced by the ITF outflow (relatively cool and fresh upper thermocline) at times of intensified thermocline ITF, whereas upper ocean water masses at this site retain characteristics of the tropical Indian Ocean (deep thermocline, relatively warm, saline and nutrient depleted water masses at habitat depth of P. obliquiloculata), when the thermocline ITF is relatively weak and the ITF outflow becomes dominated by surface flow (Fig. 9). Today, the vertical structure of the Timor Sea outflow is strongly influenced by seasonal fluctuations in the trade wind-monsoon systems. Recent INSTANT mooring data from the Timor Strait indicate a clear minimum in surface flow during austral summer when the Australian monsoon is strong, whereas thermocline flow remains relatively consistent at c. 150 m water depth (Atmadipoera et al. 2009; Sprintall et al. 2009). The surface outflow from the Timor Strait is intensified during austral winter (Tomczak & Godfrey 1994; Sprintall et al. 2009), when trade winds push warm surface waters from the NW Australian shelf westwards

resulting in deepening of the thermocline in the southwestern Timor Sea (Fig. 9). The capping effect of this light surface water prevents upwelling in the southern part of the Timor Sea along the Australian Shelf, resulting in lowered primary productivity (Takahashi & Okada 2000). In contrast, trade winds promote upwelling off the Sumatra– Java –Timor coast, and as a result the NE –SW gradient in SST increases during austral winter in the Timor Sea (Sprintall et al. 2003). Influence of ITCZ position and trade-monsoon wind systems. On millennial to orbital timescales, the relative intensity of SE trades and NW monsoonal winds over the Timor Sea is closely related to the position of the Intertropical Convergence Zone (ITCZ) during austral summer (Griffiths et al. 2009). The most important factors influencing the seasonal displacement of the ITCZ are the interhemispheric temperature contrast (Broccoli et al. 2006) and radiative forcing over Australia, the Australasian archipelago and Asia (Wyrwoll et al. 2007). When the displacement of the ITCZ extends to NW Australia, as it does today, strong monsoonal winds prevail over the Timor passage during austral summer (Fig. 10). In contrast, when the ITCZ remains further north and does not reach NW Australia, SE trade winds continue to be influential over NW Australia and the Timor Sea over most of the year. Thus, on longer timescales, the ITCZ position and monsoon intensity over the

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Fig. 8. Comparison of benthic d18O and d13C in Core MD01-2378 (1783 m water depth, eastern Indian Ocean, black), Core SO18460 (1883 m water depth, Timor Strait, red), and Core SO18479 (2983 m water depth, eastern Indian Ocean, blue). Note similarity in d13C values in the Timor Strait and eastern Indian Ocean during warmer stages (intensified ITF outflow). In contrast, d13C values diverge during glacial periods, implying increased differentiation of deep and intermediate water masses within the Timor Sea. Abrupt increases in d13C at the onset of Terminations 1 and 2 in Core SO18479 (purple bands) suggest rapid improvement in deepwater ventilation in the Timor Sea during early deglaciation.

region are strongly controlled by precessional interhemispheric insolation forcing (Holbourn et al. 2005; Wyrwoll et al. 2007). Our proxy records suggest that extended intervals of thermocline deepening in the southwestern Timor Sea (Core MD01-2378) and thermocline shallowing in the Timor Strait (Core SO18460) occurred between 112 –86 ka and 56– 20 ka (Fig. 5), implying an overall decrease in ITF thermocline outflow. These decreases may partly reflect general slowdowns in the thermohaline circulation but also likely relate to periodic weakening of the Australian summer monsoon and intensification of trades during long-term sea level falls. Spooner et al. (2005)

also reported strongly reduced monsoonal rainfall during MIS 3 and MIS 2 in the Banda Sea, east of Timor. Widespread exposure of the Sunda and Sahul shelves during prolonged glacial episodes may have triggered changes in radiative forcing that helped to lock the ITCZ closer to the equator, resulting in substantial weakening of the Australian summer monsoon (Turney et al. 2004; Spooner et al. 2005). However during the MIS 4 sea level lowstand, no change in ITF thermocline flow is detected in our proxy records, implying that additional factors are at play in controlling monsoon evolution and ITF variability in this area during sea level lowstands.

ITF VARIABILITY

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Fig. 10. Mean January–February precipitation rates (in mm/day) and 925 hPa wind vectors (arrows) in the Australia– Asian monsoon region derived from CMAP (Xie & Arkin 1997) and NCAR reanalysis (1951– 2000) (Wang et al. 2003). (a) Austral summer ITCZ position is indicated by dashed blue line. Today, maximum summer monsoon rainfall occurs over Timor. (b) Northward shift of austral summer ITCZ position, indicated by dashed red line would result in stronger (almost annual) influence of trade winds over the Timor Sea.

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High resolution thermocline temperature reconstructions in Core MD01-2378 also provide evidence for substantial ITF variability on millennial timescales during MIS 3 in the southwestern Timor Sea (Zuraida et al. 2009). Rapid cooling of the northern hemisphere and warming of the southern hemisphere during Heinrich Events are reflected by local changes in the upper ocean thermal gradient. Zuraida et al. (2009) reported successive episodes of thermocline warming during Heinrich Events 3, 4 and 5, and interpreted these events as overall decreases in ITF intensity due to a weakened thermohaline circulation (Bond et al. 1993; Paillard & Labeyrie 1994; Ganopolski & Rahmstorf 2001; Clark et al. 2007). Each of these warming pulses is preceded by a short-lived cooling of the thermocline, suggesting transient increases in ITF thermocline flow in the early phases of Heinrich Events 3, 4 and 5 (Fig. 11). These sudden increases in cool thermocline flow are

consistent with brief episodes of intensification of the Australian summer monsoon following southward displacement of the ITCZ due to southern hemisphere warming and northern hemisphere cooling. Influence of sea level on ITF variability during Termination 1. A major difference between the glacial and interglacial palaeogeography of the ITF pathways is the presence of a marine connection between the South China Sea and Java Sea with a sill depth of 40 m during warmer climate phases. Today, this open gateway allows a major transfer of freshwater from the South China Sea into the main ITF path, which leads to development of a freshwater plume at the southern end of the Makassar Strait. This in turn results in the blockage of the warm surface water flow through the Makassar Strait and promotes a cool ITF thermocline flow into the Banda Sea and Timor Sea

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ITF VARIABILITY

(Gordon et al. 2003). Our records from the main ITF outflow through the Timor Strait (Cores SO18460 and SO18462) indicate that major freshening and cooling of upper thermocline waters occurred at c. 9.5 ka (Fig. 6), which closely corresponds to the time, when eustatic sea level rose to 40 m below present day level (Lambeck & Chappell 2001) and allowed the establishment of a shallow marine connection from the South China Sea to the Java Sea through the Karimata Strait. Another major change in the upper ocean structure of the Timor Strait at c. 14.5 ka appears also to be controlled by changes in sea level and regional topography. The trend in thermocline shoaling, which prevailed in the Timor Strait since the LGM and was probably driven by strong trades, became reversed from c. 14.5 to 9.5 ka, as shown by proxy data from Cores SO18460 and SO18462 (Fig. 6). We suggest that this deepening of the thermocline within the main ITF outflow was related to the relatively swift flooding of tropical shallow land in the Australasian region including the Sunda shelf during Meltwater Pulse 1a (14.6–14.3 ka, Hanebuth et al. 2000). The change in radiative forcing associated with rapid flooding of vast tropical shelf areas possibly led to a southward shift of the ITCZ position during austral summer, which was coupled with intensification of the Australian summer monsoon.

Variability of ITF deepwater outflow into the Timor Sea Benthic d13C data in Cores MD01-2378, SO18460 and SO18479 allows close tracking of the evolution of deepwater masses in the Timor Sea during the last two glacial cycles (Fig. 8). The difference in the benthic d13C signals from the Timor Strait and eastern Indian Ocean is mainly related to the degree of mixing between relatively older, poorly ventilated AAMW, which carries a large component of old Pacific intermediate water, and younger, better ventilated southern sourced IOCW. Intensification of the deep ITF outflow results in increased transport of AAMW from the Indonesian archipelago (Banda Sea) into the Timor Sea. As this low d13C water spreads out and mixes with IOCW, the d13C gradient between the more proximal Core SO18460 and distal Core MD01-2378 decreases. Conversely, when the deep ITF outflow is reduced or the influx of better ventilated IOCW into the Timor Sea becomes enhanced, the d13C gradient between the cores increases. As the three cores are located in relatively similar productivity regimes, changes in local productivity, which may for instance occur on glacial –interglacial timescales

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(Holbourn et al. 2005), do not primarily affect the d13C gradient between cores. Core SO18460 in the Timor Strait and Core SHI-9014 (Ahmad et al. 1995) in the Banda Sea exhibit virtually identical d13C profiles, implying that Core SO18460 remained bathed in deepwater sourced from the Banda Sea over the last two glacial cycles. During warmer intervals (Holocene, MIS 5.5), d13C values converge in all three cores from the Timor Sea, suggesting vigorous mixing of Indian Ocean and Timor outflow deep and intermediate waters. However, differences in density would have inhibited mixing at depth at the location of Core SO18479 (2983 m water depth), which is significantly below the sill depth of the Timor Strait (c. 1890 m) and probably remained strongly influenced by IOCW. During relatively cold intervals (earlier part of MIS 3, MIS 5.2 to MIS 5.4), d13C values are quite similar in Cores SO18460 and SO18479 but are significantly higher (by 0.3– 05‰) in Core MD01-2378, suggesting that the ITF outflow into the eastern Timor Sea became restricted, mixing decreased and better ventilated IOCW increasingly dominated in the western Timor Sea. During glacial periods (MIS 2, later part of MIS 3, MIS 4 and MIS 6), d13C values diverge in all three cores, indicating that the ITF deep outflow declined further and that deep and intermediate water masses differed significantly on a regional scale within the basin. Lower d13C values at the deeper Site SO18479 most likely reflect a decrease in deepwater ventilation probably related to major slowdown of the global thermohaline circulation during glacial periods. At the beginning of Terminations 1 and 2 (c. 18 and 135 ka), the prominent increases in d13C in Core SO18479 (Fig. 8) suggest rapid improvement in deepwater ventilation in the Timor Sea. Waelbroeck et al. (2006) interpreted a major d13C increase at c. 2000 m water depth in the tropical eastern Indian Ocean at the end of the last glacial (18.2 ka) as reflecting the arrival of better ventilated intermediate waters from the Antarctic circumpolar current, which rapidly replaced the ‘old’ glacial Indian Ocean deepwater. The d13C record from Core SO18479 supports that such water mass renewals were pervasive in the eastern Indian Ocean at the end of glacial intervals, even extending to the deeper part of the western Timor Sea. Further rapid switches in deep/intermediate water sourcing are also evident within the later part of MIS 3 in the Timor Sea, resulting in short periods of lowered d13C gradients between Cores MD01-2378 and SO18460 (Fig. 8). These intervals probably coincide with episodes of re-organization in deep and intermediate circulation associated with high latitude climate change during Heinrich events and Antarctic warming events.

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Conclusions Comparison of millennial records over the last 140 ka reveals no significant difference in SST between Cores SO18460, SO18462 and MD012378, indicating that a uniform surface water mass extended from the Timor Strait to the southeastern Timor Sea. In contrast, thermocline water temperatures were c. 2– 3 8C lower in the Timor Strait than in the eastern Indian Ocean during glacials and stadials, implying a decrease in the transport of thermocline waters. A stepwise increase of cool, fresh thermocline ITF flow occurred during Termination I with a rapid intensification in the late part of Termination I (c. 9.5 ka), when sea level reached a critical threshold of 240 m, corresponding to the sill depth of the Karimata Strait. Opening of the connection between the South China Sea and Java Sea allowed advection of low salinity waters into the main ITF pathway, prohibiting the warm surface flow through the Makassar Strait. Thus, sea level variations affecting land– sea distribution and gateway configuration and periodic re-organizations in the trade-monsoonal wind systems related to ITCZ position and regional radiative forcing had a major impact on the heat and freshwater transport within the ITF over the last 140 ka. Comparison of benthic d13C profiles in the Timor Strait and eastern Indian Ocean (c. 1800 to 3000 m water depth) indicates that the deep ITF outflow into the Timor Sea exhibits a strong glacial –interglacial contrast. During glacials (MIS 2, later part of MIS 3, MIS 4 and MIS 6), d13C values diverge in all three cores, indicating that mixing between IOCW and AAMW was reduced and deep and intermediate water masses became more stratified within the Timor Sea. Lower d13C values at c. 3000 m water depth most likely reflect a decrease in deepwater ventilation, probably related to major slowdown of the global thermohaline circulation during glacial periods. Abrupt increases in d13C at the beginning of Terminations 1 and 2 (c. 18 and 135 ka) signal major improvements in deep water ventilation through arrival of younger, nutrient depleted southern sourced waters into the Timor Sea. Therefore, variations in the ITF deep outflow over the last 140 ka appear mainly related to changes in the intensity of the global thermohaline circulation and in deep and intermediate water sourcing. This research used samples and data collected during the IMAGES WEPAMA cruise and Sonne-185 VITAL cruise. We thank the crews of RV ‘Marion Dufresne’ and RV ‘Sonne’, the Institut Franc¸ais de Recherche et Technologie Polaires and the German Ministry of Education, Science and Technology for their support. We also thank

Robert Hall, Bob Morley and an anonymous reviewer for constructive reviews. We gratefully acknowledge funding from the Deutsche Forschungsgemeinschaft (DFG-grant KU649/28-1) and the German Ministry of Education, Science and Technology (BMBF-grant 03G0185A).

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The impact of ocean gateways on ENSO variability in the Miocene ANNA S. VON DER HEYDT* & HENK A. DIJKSTRA Institute for Marine and Atmospheric Research Utrecht, Utrecht University, Utrecht, The Netherlands *Corresponding author (e-mail: [email protected]) Abstract: The existence of El Nin˜o/Southern Oscillation (ENSO) variability in past climates is still debated. Based on evidence from geological records indicating a different long-term mean climate in the tropical Pacific, a permanent El Nin˜o state has been hyothesized to exist prior to the Plio-Pleistocene transition. However, model studies of past climate and geological records suggest that ENSO variability has existed on Earth as far back as in the Eocene and the Miocene. In the early-to-middle Miocene, climate was not only warmer than today, but oceanic gateways such as the Indonesian Passage and the Central American Seaway established deep connections between the main ocean basins. Here, we analyse the effect of increased levels of atmospheric greenhouse gases and open tropical gateways on the amplitude, period and pattern of ENSO variability using results of fully coupled climate model simulations. While our model shows only small changes in ENSO variability under increased greenhouse gas levels, it suggests a significantly stronger and less frequent ENSO due to altered oceanic gateways. In particular, a deeper and more open Indonesian Passage does not prevent a Western Pacific warm pool from developing, but it allows the warm pool to shift into the Indian Ocean.

The El Nin˜o/Southern Oscillation phenomenon (ENSO) is one of the strongest and most studied modes of climate variability on inter-annual timescales. Although the centre of variability is located in the tropical Pacific Ocean, where strong sea surface temperature variations are observed, El Nin˜o affects climate, societies and ecosystems in many regions of the world (Philander 1990). The future development of this type of inter-annual variability remains controversial, mainly because the interaction between long-term mean climate changes and inter-annual variability is poorly known. State-of-the-art coupled climate models therefore predict a variety of changes for ENSO variability under increasing greenhouse gases (Collins 2005; van Oldenborgh et al. 2005). A promising method to improve our understanding of the behaviour of ENSO variability under different long-term mean climate states is to study past warm climates. Changes in ENSO variability have been suggested for the early- to mid-Pliocene, where global mean temperatures were about 3 8C higher than today (Ravelo et al. 2004). These studies indicate that prior to 2.7 Ma a permanent El Nin˜o state has persisted with a significantly reduced zonal sea surface temperature (SST) gradient in the tropical Pacific (Wara et al. 2005; Fedorov et al. 2006) and extra-tropical climate changes resembling those observed during an El Nin˜o event today (Molnar & Cane 2002). Most climate model studies for past warm climates as well as geological records with a high temporal resolution, that

allows to resolve actual ENSO variability, however, show persistent ENSO variability in the Eocene (Huber & Sloan 2001), Miocene (Galeotti et al. 2010) and the Pliocene (Haywood et al. 2007). Apart from being warmer, these past climate states were characterized by different locations of the continents, and therefore changed interconnections between the main ocean basins. In this paper, we will focus on the warm early-to-middle Miocene period. The Atlantic and Pacific Ocean basins were not only connected at high southern latitudes via the Antarctic Circumpolar Current (ACC) as they are today, but there was also a low-latitude connection through the Central American Seaway in the early Miocene. In addition, Australia was located further south and the Indonesian Archipelago had not yet uplifted, such that there was a wide and deep ocean connection between tropical Pacific and Indian Ocean. Reconstructions suggest that around the middle Miocene climatic optimum (about 10 Ma) global mean surface temperatures were 2–3 8C higher than today (Kuerschner et al. 2008). Estimates of atmospheric CO2 vary between low values of 250 ppmv (Pearson & Palmer 2000; Demicco et al. 2003; Royer 2006) and higher values of 400 ppmv to as high as 700 ppmv (Pagani et al. 2005; Royer 2006; Kuerschner et al. 2008). The different proxies that were used for these estimates are, however, consistent in the general trends they exhibit during the early-to-middle Miocene: In the earliest Miocene, CO2-concentrations dropped, probably leading to Antarctic glaciation, while

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 305–318. DOI: 10.1144/SP355.15 0305-8719/11/$15.00 # The Geological Society of London 2011.

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afterwards CO2-concentrations increased again culminating in the middle Miocene climatic optimum (16 Ma) with CO2 concentrations around 400 – 500 ppmv or even higher. In the early Miocene, several cold glaciation events were recorded (with the most prominent event named Mi1 across the Oligocene –Miocene boundary) with warmer phases in between. In this paper, we focus on the warm phases of the early-to-middle Miocene with relatively high atmospheric CO2-concentrations. By analysing the results of fully-coupled climate model simulations with early Miocene continental geometry and elevated atmospheric CO2 concentrations as well as present-day simulations our aim here is to study the effect of low-latitude ocean gateways on ENSO variability and the mean tropical Pacific sea surface temperature structure. More specifically, we will compare the mean climatology in the tropical Pacific as well as the amplitude, frequency and pattern of ENSO in simulations with open and closed tropical gateways and different atmospheric CO2 concentrations. The paper is organized as follows. In the next section the climate model experiments and boundary conditions are described. We present our results with respect to mean climatology and ENSO variability in section 3. In section 4 we discuss and summarize our findings.

Model description The model simulations were performed using the Community Climate System Model (CCSM 1.4) developed at the National Center for Atmospheric Research (NCAR), (Boville & Gent 1998). This model, in which the atmosphere, ocean, sea-ice and land surface components are fully coupled with detailed parametrizations of physical processes in each component, simulates the evolution of climate under external forcing conditions without the use of so-called flux corrections (Blackmon et al. 2001), that is, the surface heat and fresh water fluxes between ocean and atmosphere freely evolve in the model. The atmospheric model is the Community Climate Model (CCM), a general circulation model, which was run with a 3.758  3.758 horizontal grid resolution with 18 levels in the vertical, with the highest level at about 35 km (Kiehl et al. 1998). It is coupled to the Land Surface Model (LSM), a one-dimensional model of energy, momentum, water and CO2 exchange between the atmosphere and the land surface. LSM distinguishes between specified vegetation types and contains a comprehensive treatment of surface processes (Bonan 1998). The ocean model is the NCAR CSM Ocean Model (NCOM), a general circulation model on a stretched grid with 0.98 meridional

grid spacing at the equator, 1.88 at high latitudes, 3.68 zonal grid spacing, and 25 vertical levels (Large et al. 1997). The sea-ice model (CSIM) includes thermodynamic ice processes and ice dynamics and has the same resolution as the ocean model (Weatherly et al. 1998). The CCSM has been shown to reasonably reproduce ENSO variability and its teleconnections in the present-day climate (Blackmon et al. 2001; OttoBliesner & Brady 2001; Kang et al. 2002). For example, in the present-day climate, the zonal temperature gradient along the equator has a realistic amplitude. However, the cold tongue is located more to the west and the mean SST pattern is more symmetric than in observations. Amplitude, phase and propagation features of SST and thermocline depth anomalies are reasonably simulated. The dominant period of the simulated ENSO variability is in the 2–4 year range, which is slightly shorter than the observed period ranging between 3– 8 years. In this paper, we use the output of three simulations, which are extensions of those described by von der Heydt & Dijkstra (2005, 2006). These are an early Miocene simulation (20 Ma) with idealized Miocene boundary conditions and a (constant) atmospheric CO2 concentration of 710 ppmv, a present-day control simulation (0 Ma) with a 280 ppmv atmospheric CO2, and a present-day simulation with the same high atmospheric CO2 concentration (710 ppmv) as the Miocene run (0 Ma2.5  CO2). The chosen Miocene CO2 level is still high compared to most estimates of earlyto-middle Miocene concentrations, because the CCSM is known to have a relatively low climate sensitivity. Our aim here is not to accurately simulate the early Miocene climate, but rather explore the effect of both atmospheric CO2 level and oceanic gateways on ENSO variability, which justifies the rather high CO2 levels. For the Miocene simulation, we used idealized continental boundary conditions following plate tectonic reconstructions from the Ocean Drilling Stratigraphic Network (ODSN) and idealized vegetation and soil distributions of C. R. Scotese (see http://www.scotese. com). In order to purely study the effect of the opening or closing of ocean gateways we use the ocean model with a flat bottom of 5000 m depth. The land is also kept flat with a constant elevation of 350 m. The main difference with the present continental geometry is that the Central American Seaway is open and the Indonesian Throughflow is a wide and deep passage instead of the present narrow and almost closed seaway (see Fig. 1c). The solar constant is set to the present-day value. For the present-day simulations we have also used idealized boundary conditions, that is, a flat bottomed ocean of 5000 m depth, flat land (350 m

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Fig. 1. Annual mean sea surface temperature (SST) for (a) the present-day control (0 Ma) simulation, (b) the present-day high CO2 (0 Ma2.5  CO2) simulation and (c) the Miocene (20 Ma) simulation. The thick black contours in (c) indicate the Miocene continental geometry used in the simulation. Contour intervals are 2 8C. Data are averaged over the last 50 years of simulation.

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high) and zonally constant idealized land surface types. By using flat land in all simulations, we explicitly exclude effects of the Tibetan Plateau/Himalaya uplift on the West Pacific warm pool via the possible development of the Australasian Monsoon system. We have chosen for this option to purely study the effect of ocean gateways. The ocean and atmosphere models are initialized with idealized zonally constant SST and atmospheric temperature profiles. The atmosphere is initialized with a zonal velocity field determined by the thermal wind balance. In the ocean, the salinity is taken initially constant, and the velocities are zero. The model is started with no sea-ice. From these initial conditions, the CCSM was integrated to equilibrium for all three cases using an iterative procedure as described in von der Heydt & Dijkstra (2006). After the spin-up, the coupled model was integrated for 395 years for the 20 Ma simulation and 200 years for the 0 Ma and 0 Ma2.5  CO2 simulations, respectively. Here, we analyse monthly time series of the last 150 years of the coupled simulation for all three simulations. In the following analysis, we average over the last 50 years of each simulation to show the mean climatology, while we use the full 150-year-long time series for timedependent properties.

Results The overall global ocean properties of the three simulations have been described by von der Heydt & Dijkstra (2006). Therefore, we summarize here only the main differences for global sea surface temperatures (SST) and then focus on the tropical areas. The global annual mean SST distributions for all three simulations are shown in Figure 1. The overall pattern of the Miocene SST distribution (Fig. 1c) looks very similar to the present-day control simulation 0 Ma (Fig. 1a) and the present-day simulation with high CO2 level 0 Ma2.5  CO2 (Fig. 1b), but there exist some notable differences. First of all, the global average SST in the 20 Ma simulation is 21.6 8C, which is about 3.5 8C higher than in the 0 Ma simulation and 1.5 8C higher than in the 0 Ma2.5  CO2 simulation. This indicates that the warmer climate of the Miocene is not only due to higher CO2 levels, but also due to the altered land–sea distribution and probably different land surface types. When comparing global average surface air temperatures instead, the relative importance of these two contributions only slightly changes: The global average surface air temperature in the Miocene is 17.2 8C, again 3.5 8C higher than the 0 Ma simulation, but only 18C higher than the 0 Ma2.5  CO2 simulation.

The effect of increasing atmospheric CO2 levels alone, results in an almost uniform warming of the sea surface temperature (compare Fig. 1a, b), while the different continental geometry of the Miocene introduces more regional changes (compare Fig. 1b, c). At high northern latitudes, in particular in the North Atlantic, the simulated Miocene SSTs are colder than the 0 Ma and 0 Ma2.5  CO2 simulations. This is due to the weaker Gulf stream and a very weak meridional overturning in the Atlantic compared to the presentday simulations as discussed in von der Heydt & Dijkstra (2006). In the tropical oceans, the strongest warming in the Miocene occurs in the western Pacific and the Indian Ocean, which emphasizes the importance of the deep and wide Indonesian Passage in the Miocene for tropical current systems and variability.

Changes in the mean climate of the tropical Indo-Pacific Ocean The almost uniform warming due to a 2.5 times higher atmospheric CO2 concentration is again visible in a difference plot of the 0 Ma2.5CO2 and the 0 Ma simulations for the tropical Pacific and Indian Ocean (Fig. 2a). The SST at the equator in the Pacific increases slightly more than in the subtropics. This effect has been observed in many other coupled GCMs under global warming scenarios (Liu et al. 2005; Philip & van Oldenborgh 2006). Along the equator the SST also increases almost uniformly, slightly more in the Pacific than in the Indian Ocean (Fig. 2c). The positions of the minimum SST in the equatorial cold tongue and the maximum SST in the western warm pool are not affected by the increased CO2 levels in this model. When comparing the 20 Ma run with the 0 Ma2.5  CO2 simulation in Figure 2b much more pronounced differences are observed, which are mostly due to shifts of large scale SST patterns. Remember, that in these two simulations the atmospheric CO2 level is the same. In the eastern Pacific, the equatorial cold tongue changes its source from the upwelling regions south of the equator along the South American coast in the 0 Ma2.5  CO2 simulation to another upwelling region along the northern coast of South America in the 20 Ma simulation. This is also visible in Figure 1b, c. The zonal position of the cold tongue (minimum SST along the equator) is slightly shifted to the east (Fig. 2c), however, when the cold tongue position is measured in terms of the distance to the South American continent, this may be interpreted as an actual westward shift of the cold tongue in the Miocene. In the western Pacific and Indian Ocean the SST changes are

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Fig. 2. Simulated difference in annual mean SST in the tropical Pacific and Indian Oceans between (a) present high CO2 and present-day control (0 Ma2.5  CO2 minus 0 Ma) and (b) Miocene and present high CO2 (20 Ma minus 0 Ma2.5  CO2). Contour intervals are 0.5 8C. (c) Simulated SST along the equator (averaged from 28S – 28N) in the Indian and Pacific Oceans for all three simulations. Data are averaged over the last 50 years of simulation.

even more dramatic. The West Pacific warm pool shifts to the west in the Miocene simulation and covers the eastern and central tropical Indian Ocean due to the open connection between Indian and Pacific Oceans. The maximum SST along the equator (averaged from 28S – 28N) in the warm pool hardly changes between the 0 Ma2.5  CO2 and the 0 Ma simulations, but in the Miocene simulation its position is strongly shifted westwards (Fig. 2c). A main ingredient to theories of ENSO variability is the mean tilt of the thermocline in the tropical Pacific. In the present-day climate, the thermocline is relatively deep in the West Pacific warm pool

and shallow in the east where the cold tongue is located. For the present-day climate, it has been shown that the thermocline simulated by the CCSM is slightly more diffusive than in observations, but still realistically simulated (Otto-Bliesner & Brady 2001). We use here the 20 8C-isotherm as an approximation for the thermocline in the 0 Ma simulation as is generally done in model studies (Otto-Bliesner & Brady 2001; Philip & van Oldenborgh 2006). For the simulations with elevated CO2, we use the 22 8C-isotherm in case of the 0 Ma2.5  CO2 run and the 23 8C-isotherm in case of the 20 Ma run, respectively. These isotherms approximate for each simulation the depth of the

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maximum vertical gradient in temperature (not shown). The depth of the thermocline as approximated by these isotherms is shown in Figure 3a–c for all three simulations. The general structure of the thermocline depth in the two simulations with present-day continental geometry is very similar, and only slightly affected by higher atmospheric CO2 levels. In the warmer climate of the 0 Ma2.5  CO2 simulation, the thermocline in the West Pacific is not as deep as in the 0 Ma simulation, indicating that with increased

CO2 levels the mean thermocline tilt along the equator is slightly reduced (see Fig. 3a, b, d). Opening the Central American Seaway and the Indonesian Passage has a much stronger impact on the thermocline structure in the equatorial ocean as becomes evident when comparing the Miocene simulation 20 Ma with the 0 Ma2.5  CO2 simulation. As was seen in the SST distribution, the upwelling region feeding the cold tongue in the eastern Pacific extends to the NE into the Central American Seaway in the Miocene simulation

Fig. 3. Annual mean thermocline depth (TCD) for (a) the present-day control 0 Ma simulation approximated by the depth of the 20 8C isotherm, (b) the present-day high CO2 0 Ma2.5  CO2 simulation approximated by the depth of the 22 8C isotherm, and (c) the Miocene 20 Ma simulation approximated by the depth of 23 8C isotherm. Contour intervals are 10 m. Panel (d) shows the difference between present high CO2 and present-day control (0 Ma2.5  CO2 minus 0 Ma) with contour intervals of 5 m and (e) shows the difference between Miocene and present high CO2 (20 Ma minus 0 Ma2.5  CO2) with contour intervals of 10 m. Data are averaged over the last 50 years of simulation.

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(Fig. 3b, c). In the central western Pacific the equatorial thermocline becomes shallower in the Miocene simulation, but at the same time it is much deeper in the eastern Indian Ocean (west of 1308E), indicating a westward shift of the western Pacific warm pool (Fig. 3e). Overall, the mean climatology in the tropical Pacific is only weakly affected by higher CO2 levels in our model simulations (compare 0 Ma and 0 Ma2.5  CO2), which is consistent with earlier studies using the CCSM (Zelle et al. 2005). The SST and thermocline depth patterns are, however, significantly changed when using the Miocene continental geometry with an open

(a)

(b)

(c)

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Central American Seaway and a mostly unblocked Indonesian Passage. This can be explained by strong surface current reorganizations not only in the opened gateways but in the whole equatorial ocean as becomes visible in Figures 4 and 5. Figure 4 shows the depth-integrated horizontal velocities for all three simulations. In the presentday simulations (Fig. 4a, b) there is counterclockwise (clockwise) flow between the equator and 158N (158S). Just north of the equator in the central Pacific where these two flow cells interfere, the North and South Equatorial Currents with the North Equatorial Counter Current between them become visible. In the Miocene simulation

Ma

0 Ma2.5 × CO2

Ma

Fig. 4. Annual mean vertically integrated horizontal flow velocity (units m2/s) in the Indian and Pacific oceans for (a) the present-day control 0 Ma simulation, (b) the present-day high CO2 0 Ma2.5  CO2 simulation and (c) the Miocene 20 Ma simulation. Data are averaged over the last 50 years of simulation.

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(a)

Ma

(d)

Ma

(b)

0 Ma2.5 × CO2

(e)

0 Ma2.5 × CO2

(c)

Ma

(f)

Ma

Fig. 5. Annual mean zonal (west to east) velocity in cm/s as a function of depth and latitude at two zonal locations, 908E in the eastern Indian Ocean (panels a, b, c) and 1108W in the eastern Pacific Ocean (panels d, e, f). (a) 0 Ma simulation at 908E; (b) 0 Ma2.5  CO2 simulation at 908E; (c) 20 Ma simulation at 908E; (d) 0 Ma simulation at 1108W; (e) 0 Ma2.5  CO2 simulation at 1108W; (f) 20 Ma simulation at 1108W. Solid (dashed) contour lines indicate eastward (westward) flow. Contour intervals are 10 cm/s in panels a, b, c and 15 cm/s in panels d, e, f, respectively. Data are averaged over the last 50 years of simulation.

(Fig. 4c) both transport cells (north and south of the equator) become much stronger. The northern (counter-clockwise) flow cell extends into the Atlantic Ocean and results in an eastward extension and intensification of the (eastward flowing) North Equatorial Counter Current (NECC). The right

column of Figure 5 shows the zonal flow velocity as a function of depth and latitude in the eastern tropical Pacific (at 1108W) and clearly reveals the intense eastward flow at 5– 108N in the Miocene simulation (Fig. 5f) while in the present-day simulations at this longitude the NECC is already very

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weak (Fig. 5d, e). In the western part of the basin, the positive flow cell just south of the equator extends into the Indian Ocean and forms a large Indo-Pacific gyre (Fig. 4c), while there is only weak flow through the Indonesian Passage in the presentday simulations (Fig. 4a, b). Associated with these gyres, in the Miocene simulation an analogue of the NECC and the Equatorial Undercurrent join forming a strong eastward flowing (subsurface) current in the eastern Indian Ocean (see Fig. 5c for a cross section at 908E). The flow reorganizations just south of the equator result in the strong westward shift of the Pacific warm pool into the Indian Ocean in the Miocene simulation.

Changes in ENSO variability Based on monthly time series of the last 150 years of each simulation, in this section we analyse the interannual variability associated with ENSO. All three simulations show El Nin˜o variability on timescales of 2–8 years. The amplitude of ENSO is defined as the standard deviation of SST in a 58S – 58N region in the Pacific Ocean as given in Table 1 for the F (1508E–1608W), NINO34 (1708W–1208W) and NINO3 (1408W–908W) regions. While there is a slight (most likely not significant) increase in ENSO amplitude in the present-day simulation with high CO2 level 0 Ma2.5  CO2, in the Miocene simulation the ENSO amplitude is much larger in all regions. The strongest increase in amplitude due to the continental geometry (compare 20 Ma and 0 Ma2.5  CO2) is in the most western index region (NINO) and is about 80% increase. In the central (NINO34) and eastern (NINO3) index regions, the amplitude increases by 40% and 30%, respectively. The dominant period of ENSO variability in the present-day control simulation is between 2 and 2.5 years as can be seen from the spectrum of the NINO34 index shown as blue line in Figure 6a. This is somewhat shorter than the observed period and can be explained by a too narrow wind-response to SST anomalies in the CCSM, which inhibits the

Table 1. Amplitude of ENSO variability in the three simulations as measured by the standard deviations in 8C of monthly time series of three different ENSO indices (NINO, NINO34 and NINO3)

0 Ma 0 Ma2.5  CO2 20 Ma

NINO

NINO34

NINO3

0.83 0.86 1.58

1.05 1.15 1.61

1.02 1.08 1.40

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exitation of Rossby waves and thereby influences the period of variability (Kirtman 1997; Zelle et al. 2005). This is also the reason for the relative insensitivity of ENSO in the CCSM to changes in the background temperature. In previous simulations with increasing CO2 levels only weak changes in ENSO amplitude and period have been found (Zelle et al. 2005). This is generally confirmed by our present-day simulation with high CO2 level 0 Ma2.5  CO2. The most dominant ENSO period becomes slightly longer and there appears another peak in the spectrum around 2.7 years (see Fig. 6a, red line). When comparing the spectra of NINO34 for the 0 Ma2.5  CO2 and the Miocene 20 Ma simulation (Fig. 6b), it turns out that the dominant Miocene ENSO period is longer, around 2.7 years, most likely due to the larger zonal extent of the Indo-Pacific basin. In addition, there appear peaks in the spectrum at 4.5 and 6–8 years, which are however not significant at the 99% level. The strong increase in ENSO amplitude in the Miocene simulation is partly due to a change in the pattern of variability and partly due to a real increase in amplitude. The pattern of ENSO is defined as the first Empirical Orthogonal Function (EOF) of SST in the region between 608E–608W and 158S– 158N, spanning the tropical Indian and Pacific oceans. For the present-day climate, the EOFs are usually calculated only in the Pacific region between 1308E and 808W (Zelle et al. 2005). For our 0 Ma and 0 Ma2.5  CO2 simulations we calculated EOFs for both regions and found almost the same patterns of variability. Before calculating the EOFs the monthly time series of SST were linearly de-trended and the annual cycle was removed. No additional filtering was applied. The two most dominant EOFs and the part of total variance they explain are shown in Figure 7 for the three simulations. In all cases, the first EOF explains about 20% of total variance and the patterns look very similar. In the Miocene simulation (Fig. 7c) the pattern of variability extends more westward into the open Indonesian Passage and the Indian Ocean, which may explain the longer ENSO period as compared to the presentday simulations. The period of ENSO is among others determined by the transit time of Rossby waves across the Pacific basin and therefore, should become longer in a wider basin. The principal components corresponding to the first EOF vary at almost the same frequencies as the NINO34 index for all three simulations (not shown). The second EOF, which explains about 7– 8% total variance, indicates strong variability in the Indonesian and Indian Ocean region in all simulations. However, for the two present-day simulations (Fig. 7a, b) this EOF corresponds to longer periods of variability (5– 8 years). For the Miocene

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(a) Ma

0 Ma2.5 × CO2 NINO34 index

(b) 0 Ma2.5 × CO2 NINO34 index

20 Ma NINO34 index

Fig. 6. Spectra of the 150-year-long monthly NINO34 index time series for the three simulations. (a) 0 Ma NINO34 index (solid blue line) together with the significance levels based on an AR(1) process (dotted blue lines) compared with the 0 Ma2.5CO2 NINO34 index (solid red line) together with the significance levels (dotted red lines). (b) Miocene NINO34 index (solid black line) together with the 99, 95 and 90% significance levels (dotted black lines) compared with the 0 Ma2.5CO2 NINO34 index (solid red line) together with the significance levels (dotted red lines).

simulation instead the second EOF varies on the same timescales as the first EOF (2.5–3 years) and therefore, the combined patterns of the first two EOFs can be viewed as the real Miocene ENSO pattern. This again emphasizes the relevance of the zonal width of the basin: In the Miocene ENSO variability extends in to the Indian Ocean and partly into the Central American Seaway and, therefore, has a lower frequency. At the same time, the larger area leads to a stronger coupling

between atmosphere and ocean, which increases the amplitude of ENSO variability.

Discussion and conclusions Using the output of three simulations within version 1.4 of the fully coupled climate model CCSM, we have analysed the effect of (i) high values of atmospheric CO2 levels and (ii) open

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Fig. 7. Simulated pattern of ENSO variability as expressed in terms of the first two EOFs of SST in the area 158S–158N and 608E –608W. For each simulation the upper panel shows the first EOF and the lower panel the second EOF. (a) Present-day control (0 Ma), (b) present-day high CO2 (0 Ma2.5  CO2) and (c) Miocene (20 Ma). The annual cycle has been removed from the 150-year-long monthly time series prior to EOF calculation. The top right corner of each panel shows the percentvariance that is explained by the EOF. Shown are de-normalized patterns, that is, with units of 8C. Contour intervals are 0.2 8C.

tropical gateways on the behaviour of ENSO variability. For this goal we have compared three simulations: a present-day control simulation, one

simulation with present-day conditions but 2.5 times pre-industrial CO2 concentration, and one simulation with Miocene boundary conditions and

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the same high (2.5 pre-industrial) CO2 concentration. The effect of high CO2 levels on ENSO variability is relatively small as has been found before in a transient study with increasing atmospheric CO2 levels (Zelle et al. 2005). The analysis of Zelle et al. (2005) has revealed that this is due to a relatively weak coupling strength between the tropical ocean and atmosphere in the CCSM 1.4 model, which is mainly caused by deficiencies in the simulated wind response to SST anomalies. In our study we have used higher values of CO2 and an equilibrium simulation instead of a transient increase in greenhouse gases, which results in slightly stronger changes in ENSO than in the study by Zelle et al. (2005). In particular, we find a slight trend towards longer periods and larger amplitudes of ENSO under higher CO2 levels. The open gateways in the Miocene simulation, that is, the Central American Seaway and the Indonesian Passage, on the other hand, have a stronger effect on both the long-term mean climate in the tropical Pacific and ENSO variability. The West Pacific warm pool shifts westward and covers the central and eastern Indian Ocean in the Miocene simulation. In the East, the position of the cold tongue, measured in terms of distance from the South American coast shifts slightly westward. The open gateways also lead to a general intensification and extension of the equatorial Pacific current systems, in particular, the North Equatorial Counter Current (NECC) and the Equatorial Undercurrent (EUC). ENSO variability in the Miocene simulation shows a significantly larger amplitude and a longer period than the two present-day simulations. The longer period can be explained by a changed pattern of variability that extends into the

Indian Ocean and the Caribbean Sea, while the larger amplitude may be a result of stronger atmosphere –ocean coupling in the larger IndoPacific area. Studies based on planktic foraminiferal isotope ratios have indicated a waxing and waning of the West Pacific warm pool through the middle Miocene (Nathan & Leckie 2009). They concluded that in times with relatively low sea level stand (or major steps in the restriction of the Indonesian Passage) a warm pool could develop, together with a more intense EUC bringing nutrients to the eastern tropical Pacific. On the other hand, in times with less restricted Indonesian Passage a so-called El Nin˜o-like state developed without a West Pacific warm pool. Our model results indicate, that most likely a warm pool did exist even with a completely unrestricted Indonesian Passage, however, it was located farther west in the Indian Ocean and is, therefore, not recorded in the records analysed by Nathan & Leckie (2009). Together with the warm pool, the region with a strong EUC extends westward into the Indian Ocean in our simulations. This suggests, that during the gradual closure of the Indonesian Passage, small changes in the depth or width of the Passage could result in a weakening or displacement of this extended EUC, but most likely not to a further strengthening as was suggested by Nathan & Leckie (2009). The seemingly reduced SST difference between east and west equatorial Pacific in Pliocene records (Wara et al. 2005), could be (partly) explained by a shift of the East Pacific cold tongue under increased greenhouse gases and an open Central American seaway. Although our Miocene simulation suggests almost no shift in the absolute position of the cold

Table 2. Simulated seasonality of the SST in the eastern and western tropical Pacific. Shown are SST in 8C, averaged over two regions: NINO3-region 2108 – 2708E, 58S– 58N, and NINO-region 1508 –2008E, 58S –58N. 25 years of simulation are used to calculate the seasonal cycle. DSST is the difference between warmest and coldest month temperatures Experiment

Seasonal cycle

NINO-region

NINO3-region

0 Ma

Annual mean Warmest month Coldest month DSST Annual mean Warmest month Coldest month DSST Annual mean Warmest month Coldest month DSST

28.78 29.95 27.61 2.34 30.79 31.99 29.60 2.39 30.93 32.78 28.14 4.64

25.68 27.86 24.13 3.73 27.52 29.69 26.00 3.69 28.19 30.27 25.92 4.35

0 Ma2.5  CO2

20 Ma

˜ O VARIABILITY MIOCENE EL NIN

tongue, the distance of the minimum SST along the equator to the South American coast becomes larger, which may be interpreted as an effective westward shift. Another factor that may mask the zonal SST gradient in the tropical Pacific in reconstructions is the size of seasonal fluctuations. Significant temperature differences in the growth seasons of planktonic foraminifera or algae could strongly influence temperature reconstructions. In our simulations, the seasonal variation of SSTs in the tropical Pacific is about the same in the two present-day simulations. In the Miocene simulation, however, the seasonality becomes significantly stronger and the warm season covers a larger part of the year. In Table 2 the simulated seasonal cycle of SST in an eastern and a western tropical Pacific region is summarized. In conclusion, our model results indicate that gateway changes in the tropical Pacific can have a strong impact on both the climatology as well as the amplitude, period and pattern of ENSO variability. In particular, it seems unlikely that ENSO variability has not existed in Miocene times, because the warm pool/cold tongue SST structure in the tropical Pacific remains established even under significantly altered continental boundary conditions. On the contrary, our simulations suggest a larger amplitude and period of Miocene ENSO variability due to the zonally extended basin. The changes in the warm pool and cold tongue as well as in the pattern of ENSO variability, however, emphasize the need for a good spatial coverage of palaeoceanographic records in order to be able to detect them. A. von der Heydt acknowledges personal support from the Netherlands Organisation for Scientific Research (NWO) through a VENI grant. Computer resources were funded under project SC-192 by the National Computing Facilities Foundation (NCF) with financial support from the Netherlands Organisation for Scientific Research (NWO).

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Fedorov, A., Dekens, P. et al. 2006. The Pliocene paradox (Mechanisms for a permanent El Nin˜o). Science, 312, 5779. Galeotti, S., von der Heydt, A. et al. 2010. Evidence for active ENSO in the late Miocene greenhouse climate. Geology, 38, 419– 422. Haywood, A. M., Valdes, P. J. & Peck, V. L. 2007. A permanent El Nin˜o-like state during the Pliocene? Paleoceanography, 22, PA1213. Huber, M. & Sloan, L. C. 2001. Heat transport, deep waters, and thermal gradients: coupled simulations of an Eocene greenhouse climate. Geophysical Research Letters, 28, 3481–3484. Kang, I., Jin, K. et al. 2002. Intercomparison of atmospheric GCM simulated anomalies associated with the 1997/98 El Nin˜o. Journal of Climate, 15, 2791– 2805. Kiehl, J. T., Hack, J. J., Bonan, G. B. & Boville, B. A. 1998. The national center for atmospheric research community climate model CCM3. Journal of Climate, 11, 1131– 1149. Kirtman, B. P. 1997. Oceanic Rossby wave dynamics and the ENSO period in a coupled model. Journal of Climate, 10, 1690– 1704. Kuerschner, W. M., Kvacek, Z. & Dilcher, D. L. 2008. The impact of Miocene atmospheric carbon dioxide fluctuations on climate and the evolution of terrestial ecosystems. Proceedings of the National Academy of Sciences, 105, 449–453. Large, W. G., Danabasoglu, G. & Doney, S. C. 1997. Sensitivity to surface forcing and boundary layer mixing in a global ocean model: annual-mean climatology. Journal of Physical Oceanography, 27, 2418–2447. Liu, Z., Vavrus, S., He, F., Wen, N. & Zhong, Y. 2005. Rethinking tropical ocean response to global warming: the enhanced equatorial warming. Journal of Climate, 18, 4684–4700. Molnar, P. & Cane, M. A. 2002. El Nin˜o’s tropical climate and teleconnections as a blue-print for pre-ice age climates. Paleoceanography, 17, 1021. Nathan, S. A. & Leckie, R. M. 2009. Early history of the Western Pacific Warm Pool during the middle to late Miocene (similar to 13.2–5.8 Ma): role of sea-level change and implications for equatorial circulation. Palaeogeography, Palaeoclimatology, Palaeoecology, 274, 140–159. Otto-Bliesner, B. L. & Brady, E. C. 2001. Tropical Pacific variability in the NCAR climate system model. Journal of Climate, 14, 3587– 3607. Pagani, M., Zachos, J. C., Freeman, K. H., Tipple, B. & Bohaty, S. 2005. Marked decline in atmospheric carbon dioxide concentrations during the Paleogene. Science, 309, 600–603. Pearson, P. N. & Palmer, M. 2000. Atmospheric carbon dioxide concentrations over the last 60 million years. Nature, 406, 695–699. Philander, S. G. 1990. El Nin˜o and the Southern Oscillation. Academic Press, New York. Philip, S. & van Oldenborgh, G. J. 2006. Shifts in ENSO coupling processes under global warming. Geophysical Research Letters, 33, L11704. Ravelo, A. C., Andreasen, D. H., Lyle, M., Olivarez, A. & Wara, M. W. 2004. Regional climate shifts caused by gradual global cooling in the Pliocene epoch. Nature, 429, 263– 267.

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Neogene climate history of the Makassar Straits, Indonesia ROBERT J. MORLEY1,2* & HARSANTI P. MORLEY1 1

Palynova Limited, 1 Mow Fen Road, Littleport, Cambs, CB6 1PY, UK 2

Dept Geology, Royal Holloway University, Egham, Surrey, UK *Corresponding author (e-mail: [email protected])

Abstract: The Neogene climate history of the Makassar Straits has been assessed by combining palynological studies of two Late Quaternary cores from the ocean floor with analyses of petroleum exploration wells from the Makassar Straits, Indonesia, penetrating the Early Pleistocene to Middle Miocene. The two Late Quaternary cores span 30 ka, located offshore the Mahakam Delta, east Kalimantan, and 95 ka, from offshore south Sulawesi. The first provides a record of the vegetation and climate history of the Mahakam catchment, and indicates rain forests through the last 30 ka, but with a cooler last glacial maximum, whereas the second provides a record of vegetation of the Java Sea region and south Sulawesi, and indicates extensive grasslands, suggesting a distinctly seasonal climate, during the last glacial maximum. Based on a climate model constructed from the cores which link sea level change with changes of temperature and seasonality, the history of vegetation and climate for the Makassar Straits is then extrapolated back to the Middle Miocene using the record obtained from the two exploration wells. Results show that the equatorial climate has been everwet since the Middle Miocene, but at subequatorial latitudes seasonal climates became established from the Late Pliocene onward.

The Makassar Straits (Fig. 1) lie at the centre of the SE Asian rain forest block, the second most extensive, and probably the most diverse, of the three main tropical rain forest blocks, and is currently maintained by tropical everwet equatorial climates (Whitmore 1975; Morley 2000). Understanding the climate history of this area, in terms of both the Quaternary timescale and deep time, is an essential prerequisite to understanding the development of the region’s flora and fauna. This study reviews evidence for palaeoclimates through this area back to the latest Middle Miocene. Firstly palynological evidence is reviewed for climate history from terrestrial localities, spanning the last glacial maximum (LGM) and Holocene. The terrestrial records are then compared with analyses from deep sea cores which cover the same stratigraphic interval, one in an equatorial position from the toe of the Makassar deep sea fan, and the other in a subequatorial setting from the southern end of the Makassar Straits offshore Sulawesi. The history of vegetation and climate for the Makassar Straits is then extrapolated back to the latest Middle Miocene for the Kutei Basin and Late Miocene for Tarakan Basin (Fig. 1) using the palynological record obtained from shelf and deepwater exploration wells from the region. Interpretation is based on a simple climate model constructed from these two cores which links sea level change with change of temperature and seasonality.

Late Quaternary climate change The Terrestrial record Palynological studies from three localities on Borneo (Fig. 1) provide a record back to the LGM: Anshari et al. (2001, 2004) examined sediment cores from Lake Sentarum in Western Kalimantan; studies of the Sebangau peats in southern Kalimantan were undertaken by Morley (1981), Page et al. (2004) and Kershaw et al. (2001); and the Mahakam Delta succession was studied by Caratini & Tissot (1985). The Lake Sentarum core by Anshari et al. (2001, 2004) provides a record from a peat core back to about 30 ka and indicates that everwet rain forest encompassed the area throughout this period. However, during the last glacial maximum, the climate was cooler, based on the increased representation of pollen of montane taxa at this time. Unfortunately the record may not be complete; there are breaks in the succession from 4–13 ka and from 18 –26 ka, but there is convincing evidence for everwet climates over part of the LGM from 13– 17 ka. The record for seasonal climate indicators, such as grasses, is negligible, until the Late Holocene, during which period grasses may have formed part of the swamp vegetation and are likely to be of anthropogenic origin (Beccari 1904; van Steenis 1957), rather than indicating seasonality of climate.

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 319–332. DOI: 10.1144/SP355.16 0305-8719/11/$15.00 # The Geological Society of London 2011.

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300 ka for the 101 m horizon. Unfortunately the assumption that the .30 ka date reflects the LGM has been used to model the Mahakam succession by Sydow (1996) and Roberts & Sydow (1996) and the likelihood is that modern sedimentation rates for the delta have been greatly overestimated. The significance of the grass pollen acme at this horizon is also disputable, as the event has not been duplicated in detailed analyses of a well dated marine core from the Mahakam Fan (see below).

The Marine record

Fig. 1. Positions of Papalang-10 and Sangkarang-16 gravity cores, of Attaka Well B in Kutei Basin and Well A in Tarakan Basin, and other localities mentioned in the text.

For the Sebangau region, Morley (1981) showed that thick peat swamps in this area developed over seasonal grasslands, which were dated to the LGM by Page et al. (2004) and later by Kershaw et al. (2001). The climatic ecotone for the Bornean LGM, with seasonal climates in the south, and everwet climates at equatorial latitudes, was used to restrain the atmospheric circulation model for the region at the time of the last glacial by Cannon et al. (2009). Caratini & Tissot (1985) undertook palynological analyses on a 600 m (very discontinuous) core from the Mahakam Delta, which spanned the entire Quaternary period. They radiocarbon dated the upper part, which suggested that Holocene sediments extended to 71 m (although dates were internally inconsistent), and a radiocarbon date just below this depth at 121 m yielded an age of .30 ka. A horizon just above this depth, at 101 m, yielded common grass pollen, and this has been widely used to indicate seasonal climates at the LGM in this area (Whitmore 1987; Flenley 1998; Bird et al. 2005). However, there are major problems with both the age and climate inferences from this study. Firstly, it is unlikely that the grass pollen is from the LGM; the radiocarbon date is at best a ‘minimum’ date, and if we apply regional subsidence rates to this core, based on well-restrained nannofossil dates for the Plio-Pleistocene from nearby petroleum exploration wells of 3 m per 10 ka (see below), it seems likely that the 30 ka date for the 121 m horizon is far too young. Simple application of subsidence rates would suggest an age of

A generalized vegetational history for the whole Mahakam River catchment has been obtained from a 7 m gravity core from the toe of the Mahakam Fan, taken at 2341 m water depth (Figs 1 & 2) at latitude 08, during the 2003 Unocal Multibeam project for the Papalang Block (Decker et al. 2004). In addition to palynology, the core was studied for foraminifera and oxygen isotopes, which provided good time control and indicates that the core spans approximately 30 ka with continuous sedimentation (Morley et al. 2004) with none of the time gaps seen in terrestrial cores. The data indicates that the Mahakam catchment has experienced an everwet climate throughout the last 30 ka, but that the period of the LGM was much cooler, indicated by the much more common occurrence of pollen from montane forests prior to the Holocene. Grass pollen was present throughout in low, but consistent frequencies of 2–5% with minima during the LGM. Grass pollen is generally associated with sedge pollen in this section, and it is likely that both are derived from swamp vegetation rather than indicating seasonality of climate. The high grass pollen values seen intermittently in the Misedor core of Caratini & Tissot (1985) are therefore likely to reflect localized grass swamps rather than indicating regional climate change as was previously supposed. A second core (Sankarang-16) was obtained from 783 m water depth from offshore southern Sulawesi, at 68S (Figs 1 & 3), penetrating just 5 m of sediment but with a record dated by O18 to about 95 ka at the base. Here, a strong seasonal climate signal is shown from the distribution of grass pollen, with low Gramineae pollen values during isotope stage 5a, during which time the climate was everwet, intermediate values during stages 4 and 3, and very high values during stage 2. Within the stage 2 interval grass pollen is accompanied by common charcoal debris, suggesting the widespread burning of open woodland or savannah vegetation, and thus confirming strong seasonality of climate in this area during the LGM. The postglacial sea level rise is reflected in both the Papalang-10 and Sangkarang-16 cores from the

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Fig. 2. Papalang-10 core, from toe of Mahakam Fan, 2341 m water depth (courtesy of Anadarko and Migas), palynology and foraminifera. Main components of palynomorph assemblage groupings as follows: Backmangrove’, Sonneratia and Brownlowia pollen; ‘Coastal’, Terminalia, Thespesia, Barringtonia pollen; ‘Rain forest’ pollen of Dipterocarpaceae, Burseraceae, Leguminosae and many others; ‘Kerapah’, Casuarina and Dacrydium pollen; ‘Peat swamp’, Austrobuxus, Blumeodendron, Cephalomappa, Calophyllum, Durio and Sapotaceae pollen; ‘Seasonal’, Gramineae and Compositae pollen; ‘Lower Montane’, Lithocarpus type pollen; ‘Upper Montane’, Podocarpus, Phyllocladus and Alnus pollen. The groups reflect the main vegetation types from which the fossil pollen is thought to be derived – the taxa are not wholly restricted to the vegetation types indicated.

mangrove pollen record, which exhibits a strong abundance peak coinciding with the period of sea level rise and subsequent delta coastline progradation. There is a strong correlation in marine sediments across the broad region between acmes of mangrove pollen and periods of rapid sea level rise, with major mangrove acmes coinciding with strong negative deflections of the O18 curve (Morley 1996; Grindrod et al. 1999). For the SE Asian region, mangrove acmes from marine successions are particularly useful since they specifically relate to periods of rapid eustatic rise rather than to gradual tectonic subsidence. For the Mahakam Delta, as noted above, regional subsidence is of the order of 3 m per 10 ka, whereas periods of sea level rise following glacial terminations characterized by mangrove pollen acmes (Figs 2 & 3) are at least an order of magnitude higher, typically 110 m per 10 ka. The palynological response of such periods of rapid sea level rise may thus be preserved irrespective of tectonic subsidence rates. Combining results from both Papalang-10 and Sangkarang-16 cores, the climate scenario for the

(Bornean side of the) Makassar Straits based on the combination of the pollen record from terrestrial sources, and the Holocene mangrove pollen acme is shown in Figure 4a, with rain forests and everwet climates being restricted to the equatorial belt during the LGM, and with significantly cooler temperatures across the region. This model has been extrapolated further into two ‘climatograms’ for the equatorial and subequatorial regions in Figure 4b; this figure provides the key to palaeoclimatic interpretations for the Neogene presented below.

Neogene climate change For the Neogene, the palynological record can be used to identify climate change (a) by comparing palynological signals from rain forest and seasonal climate vegetation, and (b) from the mangrove pollen record. Periods of rapid eustatic change, identified from mangrove pollen acmes, can be used to infer a change from cooler to warmer

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Fig. 3. Sangkarang-16 core, offshore Sulawesi, from 783 m water depth (modified from Morley et al. 2004). Refer to Figure 2 for content of palynomorph groups.

climates, since eustatic sea level change corresponds directly to sequestration of polar ice, which is itself controlled by global climate change (Chappell & Shackleton 1986). It is difficult to use

palynology to identify temperature change in the Neogene from hinterland elements in the same manner as in the Quaternary; the main palynomorph used to characterize cooler Quaternary climates is

Fig. 4. (a) Makassar Straits climate model for last 100 ka, and (b), climatograms used to characterize Neogene climates (see below).

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Lithocarpus type pollen (Oak family), which many palynologists find difficult to identify in Tertiary sediments. For the Neogene, analyses have been obtained from cuttings samples rather than cores; the limitations of using cuttings samples, and the limitations of resolution imposed from the analysis of samples at typically 10 or 20 m intervals, needs to be taken into account (e.g. mixing of sediments from adjacent stratigraphic levels may blur microfossil signals). However, the clarity of assemblage changes suggests that such mixing is minimal. The succession of events also needs to be carefully considered in the light of sedimentation rates, which can be obtained from nannofossil analyses by reference to the timing of successive nannofossil extinction horizons (e.g. see International Stratigraphic Commission website www.stratigraphy.org, 2009).

Kutei Basin, record from Attaka Well B The Attaka Well B is located to the NE of the Mahakam Delta in the Attaka Field (Fig. 1). It was studied for foraminifera, nannofossils and palynology with foraminiferal and nannofossil analyses at 10 m intervals and palynology at 20 m intervals. Dating the succession in this area is problematic due to its marginal marine, deltaic setting, but, by applying a ‘focused nannofossil’ approach where any horizon which yielded a foraminiferal pulse was examined for nannofossils using multiple slides, a reliable breakdown eventually emerged from the presence of age-restricted nannofossils (Fig. 5). Some intervals were poorly dated, but the widely correlatable Rt marker based on a resistivity log shift (e.g. Waryan et al. 2001), dated to 7.1 Ma in slope wells (Morley et al. 2006) provides a critical tie. The succession was studied for palynology and foraminifera in order to interpret a ‘sequence’ succession of transgressive/regressive packages, with the aim of tying the succession into the sequence biostratigraphic framework of Morley et al. (2006), for the Kutei deepwater acreage. Nannofossil analysis indicated that the succession ranges in age from Pleistocene at the top to basal Late Miocene/topmost Middle Miocene at the very base, with sedimentation falling into three broad packages: (a) an ‘early’ Late Miocene interval, spanning nannofossil zones NN8–11B (top zone NN11B dates the Rt resistivity marker on the slope based on top Discoaster bergeni following Morley et al. 2006), with very high sedimentation rates, over twice that of the modern Mahakam Delta; (b) the ‘late’ Late Miocene to Early Pliocene (nannofossil zones NN11B to NN15), with much reduced sedimentation rates, half that of modern Mahakam, and (c) the Late Pliocene and Pleistocene

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(nannofossil zones NN16 –19), with intermediate sedimentation rates, comparable to that seen in the current Mahakam Delta. The Kutei sequence succession of Morley et al. (2006) could be readily interpreted through most of the well, except through the DL sequence (8.5 to 7 Ma) where a larger number of transgressive/regressive packages (thirteen in all) could be seen compared to five typically seen in most deepwater successions (Fig. 5). The reliability of data from the basal Late Miocene, within NN8–NN10, from 12 0000 to 78600 , is low, and therefore discussion is restricted to the succession from base NN11 (about 8.5 Ma) to the Pleistocene, about 1 Ma. Transgressive/regressive packages could be modelled throughout the well using the following criteria: (a) acmes of riparian pollen and (b) mangrove pollen acmes characterize transgressive intervals, followed by (c) foraminiferal acmes at the maximum flood, and (d) rain forest elements predominating during regressive episodes. The succession is discussed below under these three headings. Due to the nature of the sampling with cuttings samples, not all the indicators are likely to be present in each package, and clearly some are better defined than others, but the main ‘pulse beat’ of the succession is clear. The value of the ‘riparian’ group in characterizing transgressive packages in very proximal successions is shown by the palynological succession of the modern Kutei lakes, where during the early Holocene with rising sea levels, palynomorph assemblages are dominated by riparian pollen, and this is then replaced by rain forest elements within the late Holocene as sea levels stabilized (Hope et al. 2005). ‘Early’ Late Miocene, 7860 0 to 4200 0 , 8.5 to 7.1 Ma. This succession is quite accurately dated using the lowest occurrence of Discoaster quinqueramus to date the base, and the Rt marker dated at 7.1 Ma at the top. There are at least 13 transgressive/regressive cycles in this interval, giving a period of 107 ka per cycle, close to the Earth’s eccentricity cycle. The cycles show a strong rhythmical pattern, within mangrove and hinterland pollen groups which are mainly thought to reflect changes in lowland and coastal vegetation associated with oscillating sea levels and patterns of coastal progradation. The very high subsidence rate over this interval has been matched by an equal rate of sedimentation, ensuring that although water depths remained more or less constant; the character of each transgressive/regressive cycle could be recorded using 600 sample spacing. ‘Late’ Late Miocene to Early Pliocene, 4200 0 to 2850 0 , 7.1 to 3.4 Ma. There are about six

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Fig. 5. Palynological and foraminiferal record for Attaka Well B, located to the NW of Mahakam Delta, showing age, nannofossil zones, sedimentation rates, (black line shows actual rates uncorrected for compaction, (a) 7.5 m/10 ka;

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Fig. 5. (Continued) (b) 1.1 m/10 ka; (c) 2.7 m/10 ka), mangrove pollen acmes, hinterland trends, foraminiferal abundance and palaeoenvironments (courtesy of Chevron IndoAsia and Migas). See Figure 2 for details of palynological groups.

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palynological cycles in this interval, with each cycle representing about 600 ka. Each cycle shows abundant riparian pollen at the base, with rain forest elements increasing farther up the section, and some acmes of mangrove pollen. It is thought that bypass is an important process through this interval, and that only a part of the sedimentary succession is preserved, otherwise, more palynomorph cycles, or more homogeneous assemblages, would be expected. Late Pliocene to Pleistocene, 2850 0 to 1230 0 , 3.4 to 1.5 Ma. The intermediate sedimentation rates through this interval ensure that high resolution transgressive/regressive cycles are preserved, with strong mangrove pollen acmes reflecting the high amplitude sea level oscillations of the Quaternary. Examination of the broader scale changes within hinterland pollen shows some additional trends; cycles show increasing percentages of riparian pollen and decreasing ‘rain forest’ pollen through Sequence DPP up to cycle 5, this pattern is then repeated through cycles 6 to 13. The reverse trend is shown from the latest Miocene through to the Pleistocene. These trends are thought to reflect relative differences in transportation patterns of lowland and coastal plain pollen onto the shelf with changes in position of the well in relation to the former coastline through successive sea level cycles. The trends roughly parallel the abundance of foraminifera recorded through the well. Evidence for climate change. Patterns of climate change have been interpreted using the model shown in Figure 4a, b. There is minimal evidence for change in seasonality of climate through individual cycles irrespective of stratigraphic position; rain forest elements dominate all cycles from base to top, from the base of the Late Miocene, up to the Quaternary. The only interval with regular grass pollen is the Early Pliocene, where very low numbers of Gramineae pollen were recorded in Sequences AT and SA. The main conclusion, therefore, is that everwet climates have occupied the equatorial Makassar Straits, and within the Mahakam catchment area in particular, from the at least basal Late Miocene onward, irrespective of the global climate scenario, during intervals characterized by both high and low sea levels, and thus warmer and cooler climate. Evidence of periodic eustatic sea level rise, providing a proxy for increasing temperatures, based on mangrove pollen acmes, can be seen throughout the succession. Mangrove pollen acmes are particularly prominent in the Quaternary, where a succession of strong acmes mirrors the rapid sea level rises of glacial terminations. Within the Late Miocene and Pliocene, they are more subdued, but

still prominent, with clear acmes during transgressive episodes of the DL Sequences 2, 5, 9 and 11, and the MDB sequence. It is highly possible that the 20 m cuttings analyses are insufficient to pick up all of the mangrove pollen acmes which are likely to occur through the sedimentary succession, but sufficient events have been recorded to approximately illustrate the pulse beat of the succession. Although there is no significant record of seasonality of climate from the Attaka succession, this may be due to periods of more seasonal climate being restricted to times when sea levels dropped below the shelf edge, and hence are not preserved in the lithological succession on the shelf. The nature of the record during periods of lowest sea level would be expected to be recorded in well sections from the slope and basin floor. Several well sections of comparable age have been studied from the more distal settings on the slope, and most are rich in pollen and spores, as is characteristic of slope deposits formed predominantly from turbidites (Morley et al. 2004). Although chronostratigraphic control on the slope is greater compared to that on the shelf, the resolution of the sequence packages is less for the same sample spacing since the sedimentation rate is less. For this reason and also due to the increased transportation and mixing of pollen, the picture of vegetation change on the slope is less clear than on the shelf. However, the same main trend is apparent. As for the succession on the shelf, the record is essentially of rain forest pollen, reflecting widespread rain forests and of continuous everwet climates with minimal evidence for pollen suggesting seasonal climate vegetation. The slope record thus provides confirmation that the Mahakam catchment area has experienced an essentially everwet climate without interruption from the 8.5 Ma to the present. From the broad Kutei region, palynological data from wells studied in a similar manner to the Attaka Well B suggests that everwet climates characterized the region to at least the topmost Middle Miocene.

Tarakan Basin The second dataset used is from an unnamed well (Well A) drilled on the mid slope in the Tarakan Basin (Figs 1 & 6). As for the Kutei wells, samples were studied for nannofossils at 10 m intervals to give chronostratigraphic control, then for foraminifera and palynology at 20 m intervals to provide a basis for sequence biostratigraphic interpretation. Again the succession is very rich in pollen and spores, and sequences were modelled using acmes of mangrove pollen in shales reflecting condensed

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Fig. 6. Palynological and foraminiferal record for Tarakan Well A, showing age, nannofossil zones, mangrove pollen acmes, hinterland pollen trends and foraminiferal abundance (courtesy of Migas). This well is from the mid/lower slope. See Figure 2 for details of palynomorph groups.

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Fig. 6. Continued.

sections. Within the upper part of the well, foraminiferal acmes are also very useful for sequence modelling, but below the Pliocene, the prevalence of carbonate dissolution makes it very difficult to use foraminifera to model sequences. The Tarakan well succession is characterized by strong mangrove pollen acmes during the Quaternary, in the manner of the Attaka Well B. In the Pliocene and Late Miocene, mangrove acmes are regularly present but not with such strong acmes as for the Quaternary. The Pliocene and Late Miocene show reasonable cyclicity within rain forest elements. The most notable features of Tarakan Well A is the gradual increase in abundance of seasonal climate elements within the Late Pliocene and Pleistocene, and within the lower part of the

well, acmes of Eugeissona pollen coinciding with lowstands. Evidence for climate change. For most of the Late Miocene and Early Pliocene, the pollen record suggests the continuous representation of rain forests and of everwet climates. Cyclical change occurs throughout the succession, mostly reflecting change of coastal and lowland vegetation in response to sea level change and coastal progradation, but with minimal evidence for seasonality of climate. The only instance of greater seasonality is the presence of low percentages of grass pollen in the CA2A interval (Fig. 6). Within the Late Pliocene and Quaternary, however, there is a gradual, but persistent increase of seasonal climate elements, mainly of grass and

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Compositae pollen. The abundance of seasonal climate elements varies through each sequence, suggesting alternating more seasonal and more everwet climates, but resolution is insufficient to clearly show detailed changes in every package. The presence of seasonal climate pollen is a persistent feature of each of the Neogene successions so far studied in the Tarakan Basin and suggests that during the Quaternary and Late Pliocene, seasonal climates, probably supporting open canopy woodland, were a persistent feature of low sea level periods in the area of NE Borneo. Evidence for temperature change is forthcoming from the repeated occurrence of mangrove pollen acmes, most persistent within the Quaternary, when periods of rapid sea level rise characterized each glacial termination, but also present through the Late Miocene and Pliocene, especially within the DL2, MDBA, CA1, CA2, DS and DB intervals.

The regional picture The Bornean side of the Makassar Straits has thus experienced essentially everwet climates since the Middle Miocene, with virtually no excursions of seasonal climates supporting more open vegetation until the Late Pliocene, and then only in the NE Bornean area. Most palynological studies undertaken so far from the Makassar Neogene have been from Borneo-sourced sediments and may not apply to the eastern side of the Makassar Straits. Palynological studies (R. J. Morley unpublished) of the most easterly, basinal Kutei wells, in which sediments may have been sourced partly from Sulawesi, indicate increased grass pollen signals for the Late Pliocene, suggesting seasonality of climate in the same manner as in the Tarakan wells. Within the Early Pliocene grass pollen values are reduced, suggesting everwet climates, and then increased values occur in the Late Miocene in a manner not seen in slope wells from the Borneo side. This suggests that the climate history for Sulawesi may have followed a different course to Borneo during the Late Miocene. There is currently minimal data on the Early Miocene and Oligocene climate history for the Makassar Straits, and to infer palaeoclimates for this period it is necessary to look at data from outside the region to the West Natuna Basin (Morley et al. 2003) and the Malay Basin (Morley & Shamsuddin 2006; Shamsuddin & Morley 2006) where more comprehensive palaeoclimatic records are available (Fig. 7). From a regional perspective, the most notable features are the presence of seasonal climates in the Oligocene, but with the wettest climates characterizing the Early Miocene, and therefore it is likely that Early Miocene climates

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in the Makassar region were also everwet. Seasonal-everwet climate oscillations characterize much of the Middle Miocene to Pliocene in Natuna and the Malay Basin; in contrast, Makassar Late Miocene and Pliocene climates are much wetter and more uniform.

Conclusions The Island of Borneo currently has an essentially everwet climate, but at the time of the LGM, it was divided into an everwet belt in the equatorial zone, and with a distinctly seasonal climate supporting grasslands in the south, extending also to the area now covered by the Java Sea. The suggestion of seasonality of climate for the LGM for the Mahakam Delta area by Caratini & Tissot (1985), based on grass pollen acmes, is shown to be in error. Intervals yielding grass pollen from the Mirador core are likely to be much older than the LGM, and also such acmes have not been duplicated in a deep sea core from the Mahakam Fan. This suggests that grass pollen acmes are more likely to reflect localized grass swamps than indicating regional climate change. At equatorial latitudes in Borneo, everwet climates have been the rule for the entire Pliocene and the whole of the Late Miocene, with minimal evidence for seasonality of climate, even at times when sea levels were below the shelf edge. Mangrove pollen acmes provide a proxy for sea level and temperature change, with strong, repeated abundance fluctuations reflecting the pronounced sea level/temperature changes of the Quaternary, with more subdued, but still prominent fluctuations during the Late Miocene and Pliocene. For the period from 7.0 to 8.5 Ma, sea level/climate cycles seem to show a distinct periodicity of about 107 ka, approximating the Earth’s eccentricity cycle. In the region of NE Borneo, everwet climates were the rule for the Late Miocene and Early Pliocene, but seasonal climates were prominent at times of low sea level during the Late Pliocene and Quaternary, probably supporting open canopy woodland. The Neogene climate history of the Sulawesi side of the Makassar Straits may have parallelled that of NE Borneo during the Quaternary and Late Pliocene. However, during the Late Miocene, the climate history may have taken a different course, since (R. J. Morley unpublished) well sections closer to Sulawesi yield increased grass pollen values not seen in shelf or slope wells from the Borneo side of the Straits. There is not a palynological record for the Early Miocene of the Makassar Straits region, but for the

330 R. J. MORLEY & H. P. MORLEY Fig. 7. Climate trends in Malay (Morley & Shamsuddin 2006) and Natuna Basins (Morley et al. 2003; Morley 2006) compared to schematic climate history for Borneo margin of Makassar Straits. Alpha-numeric terminology for West Natuna basins indicate biostratigraphically defined sequences of Morley et al. (2003).

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Malay and Natuna Basins to the north, the wettest climates occurred during the Early Miocene, and therefore it is likely that Early Miocene climates in the Makassar region were also everwet. The current evidence suggests that equatorial Borneo has experienced everwet climates without interruption from the topmost Middle Miocene, and most likely from the base Miocene. Such climates would have been implemental in maintaining rich and diverse rain forests which would have facilitated the diversification and maintenance of all elements of the Bornean rain forest fauna and flora, possibly helping to explain why some areas of Borneo, such as the Lambir forest Reserve in Sarawak, contain the world’s most floristically diverse rain forests (Wright 2001; Lee et al. 2002). The authors are grateful to Chevron IndoAsia and Migas for permission to present palynological data from the Attaka Well B, and to Migas for permission to present palynological data from Tarakan Well A, and Anadarko Indonesia and Migas for the permission to discuss the Papalang-10 bottom core. Foraminiferal and nannofossil analyses were performed by Sukarno and Wantoro of Corelab Jakarta, whereas the palynological analyses on the three exploration wells were undertaken by Yuli Pane of Corelab Jakarta.

References Anshari, G., Kershaw, A. P. & van der Kaars, W. A. 2001. A late Pleistocene and Holocene pollen and charcoal record from peat swamp forest, Lake Sentarum Wildlife Reserve, West Kalimantan, Indonesia. Palaeogeography, Palaeoclimatology, Palaeoecology, 171, 213– 228. Anshari, G., Kershaw, A. P. & van der Kaars, W. A. 2004. Environmental change and peatland forest dynamics in the Lake Sentarum area, West Kalimantan, Indonesia. Quaternary Science, 19, 637– 655. Beccari, O. 1904. Wanderings in the Great Forests of Borneo. (reprinted 1986 with introduction by Earl of Cranbrook), Oxford University Press, Oxford. Bird, M. I., Taylor, D. & Hunt, C. 2005. Environments of insular Southeast Asia during the Last Glacial Period: a savanna corridor in Sundaland? Quaternary Science Reviews, 24, 2228–2242. Cannon, C. H., Morley, R. J. & Bush, A. B. G. 2009. The current refugial rainforests of Southeast Asia: major biogeographic implications. Proceedings of the North American Society, 106, 1188– 1193. Caratini, C. & Tissot, C. 1985. Le Sondage Misedor, Etudes palynologique. Etude Ge´ographie Tropicale, 3, CNRS, Centre d’Etudes de Geographie Tropicale, University of Bordeaux, France. Chappell, J. & Shackleton, N. J. 1986. Oxygen isotopes and sea level. Nature, 324, 137– 140. Decker, J., Teas, P. A., Schneider, R. D., Saller, A. H. & Orange, D. L. 2004. Modern deep sea sedimentation in the Makassar Strait: Insights from highresolution multibeam bathymetry and backscatter,

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sub-bottom profiles, and USBL-navigated cores. In: Noble, R. A., Argenton, A. & Caughey, C. A. (eds) Deepwater and Frontier Exploration in Asia & Australasia. Proceedings of the International Geoscience Conference, Indonesian Petroleum Association, Jakarta, DFE-04-PO-042, 377– 388. Flenley, J. R. 1998. Tropical forests under the climates of the last 30 000 years. Climatic Change, 39, 177– 197. Grindrod, J., Moss, P. & van der Kaars, S. 1999. Late quaternary cycles of mangrove development and decline on the north Australian continental shelf. Journal of Quaternary Science, 14, 465– 470. Hope, G., Unna, C. & Syaiful, A. 2005. The stratigraphy and fire history of the Kutai Peatlands, Kalimantan, Indonesia. Quaternary Research, 64, 407– 417. INTERNATIONAL STRATIGRAPHIC COMMISSION 2009. ISC website. http://www.stratigraphy.org. Kershaw, A. P., Penny, D., van der Kaars, S., Anshari, G. & Thamotherampillai, A. 2001. Vegetation and climate in lowland SE Asia at the Last Glacial Maximum. In: Metcalfe, I., Smith, J. M. B., Morwood, M. & Davidson, I. (eds) Faunal and Floral Migrations and Evolution in SE Asia– Australasia, A.A. Balkema, Lisse, 227– 235. Lee, H. S., Davies, S. J. et al. 2002. Floristic and structural diversity of mixed dipterocarp forest in Lambir Hills National Park, Sarawak, Malaysia. Journal of Tropical Forest Science, 14, 379– 400. Morley, R. J. 1981. Development and vegetation dynamics of a lowland ombrogenous peat swamp in Kalimatan Tengah, Indonesia. Journal of Biogeography, 8, 383– 404. Morley, R. J. 1996. Biostratigraphic characterisation of systems tracts in Tertiary sedimentary basins. In: Proceedings of the International Symposium on Sequence Stratigraphy in SE Asia. IPA, Jakarta, 49– 71. Morley, R. J. 2000. Origin and Evolution of Tropical Rain Forests. Wiley & Sons, London. Morley, R. J. 2006. Cretaceous and Tertiary climate change and the past distribution of megathermal rain forests. In: Bush, M. & Flenley, J. R. (eds) Tropical Rainforest Responses to Climatic Change. PraxisSpringer, Chichester, 1 –54. Morley, R. J. & Shamsuddin, J. 2006. The Sequence Biostratigraphy and Chronostratigraphy of the Malay Basin. PGCE Malaysia Proceedings, Kuala Lumpur, 77–78. Morley, R. J., Morley, H. P. & Restrepo-Pace, P. 2003. Unravelling the tectonically controlled stratigraphy of the West Natuna Basin by means of palaeoderived Mid Tertiary climate changes. In: 29th IPA Proceedings. 1. Morley, R. J., Morley, H. P., Wonders, A. A. H. W., Sukarno & Van Der Kaars, S. 2004. Biostratigraphy of Modern (Holocene and Late Pleistocene) Sediment Cores from Makassar Straits. In: Proceedings Deepwater and Frontier Exploration in Asia & Australasia, Indonesian Petroleum Association, Jakarta, December 2004, DFE04-PO-26. Morley, R. J., Decker, J., Morley, H. P. & Smith, S. 2006. Development of high resolution biostratigraphic

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framework for Kutei Basin. In: Proceedings Jakarta 2006 International Geoscience Conference, August 14–16, 2006, IPA Jakarta, 06-PG-27. Page, S. E., Wu, R. A. J., Weiss, St. D., Rieley, J. O., Shotyk, W. & Limin, S. H. 2004. A record of Late Pleistocene and Holocene carbon accumulation and climate change from an equatorial peat bog (Kalimantan, Indonesia): implications for past, present and future carbon dynamics. Journal of Quaternary Science, 19, 625–635. Roberts, H. L. & Sydow, J. 1996. The offshore Mahakam Delta: stratigraphic response of Late Pleistocene-to-modern Sea Level Cycle. In: Proceedings Indonesian Petroleum Association, TwentyFifth Silver Anniversary Convention, October 1996, 147– 161, IPA Jakarta, IPA-96-1.1.031. Shamsuddin, J. & Morley, R. J. 2006. Integration of biostratigraphy with seismic for sequence stratigraphic interpretation in the Malay Basin. In: PGCE Malaysia Proceedings, Kuala Lumpur, 101– 102. Sydow, J. 1996. Holocene to Late Pleistocene stratigraphy of the Mahakam delta, Kalimantan, Indonesia,

PhD Dissertation. Department of Oceanography and Coastal Sciences, Louisiana State University, Baton Rouge, LA. van Steenis, C. G. G. J. 1957. Outline of vegetation types in Indonesia and some adjacent regions. Proceedings of the eighth Pacific Science Congress (Manila), 4, 61–97. Waryan, S., Vo, D. T., Stites, J. & Swanson, M. 2001. Integrated 3D geological data into fluid flow model improves reservoir management plan: serang field case. In: Study SPE Asia Pacific Oil and Gas Conference and Exhibition, 17–19 April 2001, Jakarta, Indonesia, SPE, 68486-ms. Whitmore, T. C. 1975. Tropical Rain Forests of the Far East. Clarendon Press, Oxford. Whitmore, T. C. 1987. Introduction. In: Whitmore, T. C. (ed.) Biogeographical Evolution of the Malay Archipelago. Oxford Monographs on Biogeography 4, Oxford Scientific Publications, Oxford, 1 –4. Wright, S. J. 2001. Plant diversity in tropical forests: a review of mechanisms of species coexistence. Oecologia, 130, 1– 14.

Oligocene palynological succession from the East Java Sea EKO BUDI LELONO1 & ROBERT J. MORLEY2,3* 1

LEMIGAS, Jl. Ciledug Raya, Cipulir Kav.109, Kebayoran Lama, Jakarta Selatan 12230, Indonesia 2

Palynova Limited, 1 Mow Fen Road, Littleport, Cambs, CB6 1PY, UK

3

Department of Earth Sciences, Royal Holloway University, Egham, Surrey, UK *Corresponding author (e-mail: [email protected]) Abstract: Rich palynomorph assemblages have been obtained through a marine Oligocene succession from the East Java Sea (Indonesia) and provide the first instance of an independently dated Oligocene succession from SE Asia that has yielded a good quality palynological record. The succession has been independently dated by nannofossils and foraminifera. The palynomorph succession suggests climatic control on Oligocene vegetation, on which basis a regionally applicable zonation is proposed. The Early Oligocene is characterized by common rain forest elements, suggesting an everwet, rain forest climate. The early part of the Late Oligocene contains much reduced rain forest elements with grass pollen, indicating a more seasonal climate, whereas for the latest Late Oligocene, rain forest elements return in abundance, suggesting a superwet rain forest climate. This palynological succession is similar to that from the Sunda Basin in the West Java Sea, allowing the Sunda Basin succession to be dated by correlation. The study also extends the stratigraphic range of Dacrydium and Casuarina, two plant taxa previously thought to have dispersed from the Australian Plate into SE Asia at the time of the collision with Sunda, to well before the time of collision. A different means of dispersal for these taxa is proposed.

The area of study is located on the off-shore of North Madura which is a part of the East Java basin (Fig. 1). This is a back-arc basin situated on the southern margin of the Sundaland. This basin covers an area over 54 000 km2 with an east –west alignment and accommodates sediment with a thickness of more than 2000 m (Pusoko et al. 2005). It is well known as an important hydrocarbon province in Indonesia. East Java has been an attractive area for oil exploration and new discovery continues to occur within this basin. A regional stratigraphy of the East Java basin is shown in Figure 1. The stratigraphic succession follows an internal Lemigas report and also Yulianto et al. (2000). Sedimentation commenced during the Middle Eocene, unconformably over Cretaceous metasedimentary and igneous basement rocks and Paleocene/Late Cretaceous ‘PreNgimbang’ sediments, with deposition of the Ngimbang Formation. The Ngimbang Formation consists of sands, coals and lacustrine shales in the lower part, and subsequently marine shales and limestones deposited during the Late Eocene and Early Oligocene, and forming the synrift phase of deposition within the basin. Subsequently, following a period of nondeposition/erosion during the mid Oligocene, during a period of post-rift tectonic quiescence, thick

and extensive carbonates of the Kujung Formation formed over the major part of the area during Late Oligocene –Early Miocene (Johansen 2003). This was followed by clastics of the Tuban Formation during the latest Early Miocene. The Kujung Formation is divided into three units, generally termed Kujung Units III, II and I. Kujung unit III is the oldest (Late Oligocene) and consists of alternations of shale, sand and limestone. Kujung unit II is characterized by clastics and carbonates with a basal sandstone and formed during Late Oligocene–Early Miocene, whilst Kujung unit I is represented by widespread Early Miocene carbonate build-ups. The Kujung Formation is conformably overlain by the Tuban Formation which is dominated by claystones with the intercalation of foraminiferalrich marls forming during the Early Miocene in a deep shelf setting (Firdaus 2004), and includes the widespread Rancak reef limestone member. The younger Neogene succession consists of the Ngrayong, Wonocolo, Ledok and Mundu Formations. The Oligocene of the East Java Sea area is usually subdivided biostratigraphically using a combination of nannofossils and foraminifera. It is indeed one of the mainstay areas for larger foraminifera, with many classic papers from the onshore

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 333–345. DOI: 10.1144/SP355.17 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Location and lithostratigraphic succession. Arrow indicates interval studied. Stratigraphy from internal LEMIGAS report, Onlap curve from Vail et al. (1977) and Haq et al. (1988).

region (Van der Vlerk & Umgrove 1927; Leupold & Van der Vlerk 1931), and also for planktonic foraminifera (Bolli 1966). In the area NW of Madura, however, the Kujung Formation often contains interbedded clastics which introduce additional correlation problems, and to try to clarify correlations in this area well sections have been studied for palynology as well as the usual microfossil groups. Samples from two wells, termed Well X and Well Y, have been studied at 600 intervals for each discipline, with some analyses of sidewall cores. The traditional palynological zonation for SE Asia (Morley 1978, 1991) does not work well in this area, and so the succession has been divided into broad palynological assemblage zones, which appear to be controlled mainly by climate. Some pollen types previously found to be stratigraphically useful in the Sunda Basin have also been recorded, and can be used in this area to help define the zones. The palynological zones appear to provide the best criteria for correlation of the observed stratigraphic succession between the two wells. The age of the assemblage zones, however, is best determined on nannofossils, with support from larger foraminifera. This provides one of the first instances where an Oligocene palynological succession from the SE Asian region can be independently dated using marine fossils. For nannofossils, reference is made to the scheme of Martini (1971) as presented by Perch-Nielsen (1985). For planktonic

foraminifera, the scheme of Blow (1969, 1979) is followed using taxon ranges as presented by Bolli & Saunders (1985), whereas the larger foraminiferal zonation follows Adams (1970). The timescale used is that of Gradstein et al. (2004). Based on the close similarities of the palynological succession with that from the Sunda Basin, it is possible to apply at least part of the zonation scheme in the non-marine Sunda Basin Talang Akar and Cibulakan Formations, thus providing the first direct indication of the age of these formations. Apart form providing a new, independently dated, palynological zonation this paper will also discuss the regional implications of the climatic succession The sequence is also characterized by the common occurrence of palynomorph taxa of Australasian origin, and the biogeographical implications of these records are considered.

Overview of datasets The datasets are reviewed by discipline, in relation to the lithostratigraphic succession.

Foraminifera The succession yields variable recovery of foraminifera, consisting mainly of calcareous benthonics (Fig. 2), with larger foraminifera featuring

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Fig. 2. Foraminifera from Well X and Well Y, identifications by LEMIGAS.

prominently. Planktonics, however, are poorly represented. With respect to planktonics, the presence of Globigerina tripartita and Globorotalia opima nana in Well X indicates reference to the Late Oligocene Planktonic zone P22, but correlation to Well Y cannot be achieved as planktonics in that well are very rare indeed. Larger foraminifera provide a better basis for correlation (Fig. 2), since Lower Te foraminifera (Spiroclypeus spp without Miogypsina spp) are well represented in both wells. The Early Oligocene Tc Letter stage is represented in Well Y from the regular presence of Nummulites fitchelii, but not in Well X. However, the occurrence of Tb marker Nummulites variolensis near the base of the well, although present in both sections, may be due to reworking (see below).

Nannofossils Nannofossil recovery from both sections was good in the upper part, and zone NP25 can be interpreted from the interval between the highest occurrence of Sphenolithus ciperoensis and the highest S. distentus (Fig. 3). The base of NP24 is marked by the deepest occurrence of Sphenolithus ciperoensis, but this species appears much lower

stratigraphically in Well Y, where it is likely to be present closer to its true base. Its base in Well X is likely to be controlled by environmental constraints. There are no nannofossils to date the Early Oligocene, although the topmost Eocene nannofossil Discoaster saipanensis (top in NP20) occurs at the base of both wells, for which reworking is suspected. Eocene foraminifera and nannofossils are associated with a basal lag over basement in both Well X and Well Y. Foraminifera include Globorataloides carsoseleensis (tops in P16) and Morozovella lehneri (tops in P14) and Cribrohantkenina inflata (range P16/17) together with the Tb larger foram Nummulites variolensis. The Eocene assemblage is thus from a variety of different stratigraphic intervals, with taxa with ranges that do not overlap. The likelihood is that the basal lag contains a mixture of reworked Eocene fossils and may be of Early Oligocene age. The rarity of palynomorph markers for top Eocene would also support the idea that Eocene sediments are missing from these sections.

Palynology Rich pollen and spore assemblages were found more or less throughout the succession, and dinoflagellate cysts and other marine palynomorphs were also

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Fig. 3. Nannofossils from Well X and Well Y, identifications by LEMIGAS.

well represented. Pollen/spore assemblages can be divided into three groups, mangroves, hinterland pollen and spores. The hinterland pollen group shows the most interesting succession, with elements on the one hand suggesting everwet climates (Table 1), and seasonal elements on the other (Fig. 4). Mangrove pollen however suggests strong environmental control since mangrove pollen shows different abundance variations in the two wells. Pollen is compared to that of modern taxa where appropriate, with the suffix ‘type’ being added in instances when a pollen type compares to that of more than one extant taxon. A list of the palynomorph taxa mentioned in the text and their respective morphotaxon names is given in Table 2. The palynological slides from which assemblages were recorded are stored at LEMIGAS Stratigraphy Unit, Jakarta. Dacrydium and Casuarina:superwet elements. The hinterland pollen group shows pollen of Dacrydium and Casuarina type (Fig. 5) to be particularly well represented in the upper part of the succession. These taxa are of Gondwanan origin, and following dispersal into the SE Asian region found niches in peat swamps and Kerangas vegetation or ‘heath’ forests (Gymnostoma spp), montane forests (Dacrydium spp and Casuarina junghuhniana) and also along beaches (Casuarina equisitifolia). The occurrence of abundant Casuarina type and Dacrydium pollen in the Well X and Well Y wells suggests a peat swamp derivation

(e.g. these pollen types are common in coals from the laterally equivalent ‘Coaly’ Talang Akar Formation in the Sunda Basin, Morley 2000). Casuarina/Gymnostoma and Dacrydium do not occur in Table 1. Tropical moist climates Constant high temperature is characteristic of tropical climates, and climates in which the mean temperature of the coldest month is 188 or over is often used as a definition (Whitmore 1998), but the absence of frosts was shown to be a critical factor by Morley (2000). Also, the mean diurnal temperature range exceeds the annual range. Everwet, or perhumid tropical climates occur in those areas with typically more than 2000 mm of rain per year and where every month is wet (with 100 mm of rain or more) but there may be frequent short dry periods or a very short dry season (Richards 1996) and such climates permit the development of true evergreen rain forests. The term ‘superwet’, for climates in which dry periods are rare or absent and in which annual rainfall exceeds 3000 mm was proposed by Richards (1996) and from a geological standpoint is useful since such climates may permit growth of true ombrotrophic peats. Tropical wet seasonal climates may have a significant dry season of up to 4 months with less that 100 mm of rain per month (Richards 1996). Areas characterized by soils exhibiting a slight moisture deficit during the dry season may support semi-evergreen rain forests, whereas areas with a strong soil water deficit would support monsoon forest (Whitmore 1998).

OLIGOCENE PALYNOLOGY EAST JAVA SEA Fig. 4. Mangrove, hinterland pollen and Dacrydium/Casuarina from Well X and Well Y. Main components of palynomorph assemblage groupings as follows: ‘Backmangrove’, Sonneratioid and Brownlowia pollen; Nypa (in blue) ‘Coastal’, Terminalia, Thespesia, Barringtonia pollen; ‘Rain forest’ pollen of Dipterocarpaceae, Burseraceae, Leguminosae and many others; ‘Kerapah’, Casuarina and Dacrydium pollen; ‘Peat swamp’, Blumeodendron, Cephalomappa, Calophyllum, Durio and Sapotaceae pollen; ‘Seasonal’, Poaceae pollen; ‘Montane’, mainly Podocarpus pollen. The groups reflect the main vegetation types from which the fossil pollen is thought to be derived – the taxa are not wholly restricted to the vegetation types indicated.

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Table 2. Pollen types and pollen morphotaxa from Java Sea wells Pollen type Arenga type Austrobuxus swainii/Dissiliaria baloghioides Canthium type Casuarina type (Casuarina/ Gymnostoma) Crudia Dacrydium Eugeissona insignis type Eugeissona minor Ilex Myrtaceae Pandanus Pinus Poaceae Pometia Schoutenia

Morphotaxon

Family

Arengapollenites sp Malvacipollis diversus Harris (1965) Canthiumidites sp Casuariniidites cainozoicus Cookson & Pike (1954) Striatricolpites catatumbus Gonzalez Guzman (1967) Lygistepollenites florinii Stover & Evans (1973) Quilonipollenites sp Quilonipollenites sp Ilexpollenites spp Myrtaceidites spp Pandaniidites spp Pinuspollenites spp Monoporites annulatus van der Hammen (1954)

Arecaceae Euphorbiaceae

Echitriporites schoutenioides

the communities of the domed topotrophic/ ombrotrophic peat swamps which are widespread along the coasts of Kalimantan and Sumatra (e.g. Anderson 1963), but characterize the less well known ombrotrophic (blanket bog type) ‘Kerapah’ or watershed peats which occur locally in Borneo on poorly drained podsolic soils (Brunig 1990; Morley 2000) in areas of ‘superwet’ climate (Richards 1996). The occurrence of ‘Kerapah’ swamps is thus convincing evidence for an everwet climate (Fig. 4). Riparian elements. Pollen of riversides and alluvial swamps, especially of Pandanus and Ilex (Figs 5 & 6), show distinct trends, being common in the topmost part, above the interval with common Dacrydium pollen, and also common in the interval with regular seasonal climate elements (see below). Seasonal climate elements. The regular occurrence of Poaceae (Fig. 5) and Schoutenia pollen (Schoutenia is an element of seasonal evergreen forest) in the mid part of the succession suggest a more seasonal climate aspect to the mid part of the succession. This interval is also characterized by the regular occurrence of Malvacipollis diversus, which compares closely with modern pollen of Australian Austrobuxus swainii and Dissiliaria baloghioides (Martin 1974), and differs from SE Asian Austrobuxus spp in its shorter spines and smaller size. These are both trees of modern drier rain forests/warm temperate sclerophyll communities

Rubiaceae Casuarinaceae Leguminoseae Podocarpaceae Arecaceae Arecaceae Aquifoliaceae Myrtaceae Pandanaceae Pinaceae Poaceae Sapindaceae Tiliaceae

in SE Queensland/NW New South Wales (Pickett et al. 2004). Algal and other marine palynomorphs. Dinoflagellates are common through both wells and consist mainly of Operculodinium spp, Spiniferites spp and Homotryblium spp, but without useful ageindicative taxa. However, the dinoflagellate cysts do show two distinct acmes in the mid part of the succession (Fig. 7). There is one very strong acme within both wells, with abundant Operculodinium spp and Spiniferites spp, and below this a second less pronounced acme of the same taxa, but also with common Homotryblium spp. These acmes are coincident with the events based on hinterland pollen discussed above.

Palynological zonation scheme The main palynological events are brought together in Figure 7, and five zones are proposed in the basis of assemblage changes as follows: Zone OL-1. This zone is difficult to define, and is based essentially on the rarity of seasonal climate and riparian elements. Zone OL-2. This zone is characterized by the regular occurrence of seasonal climate elements, especially of Malvacipollis diversus (Fig. 5), with the top continuous occurrence of this species marking the top of the zone. Other seasonal climate elements are

OLIGOCENE PALYNOLOGY EAST JAVA SEA

Fig. 5. Pollen from Wells X, and Y, and from Talang Akar Formation. (1) Pometia pollen; (2, 3) Ilex pollen; (4) Casuarina type pollen; (5) Dacrydium pollen; (6) Pandanus pollen; (7) Myrtaceae pollen; (8) Poaceae pollen; (9) Schoutenia pollen; (10, 11) Malvacipollis diversus. Black line ¼ 10 microns.

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Fig. 6. Riparian and seasonal climate pollen from Well X and Well Y. See caption to Figure 4 for explanation of pollen groupings.

OLIGOCENE PALYNOLOGY EAST JAVA SEA

Fig. 7. Palynomorph taxa used to zone the succession in Well X and Well Y, including total marine dinocysts.

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Gramineae pollen and Schoutenia pollen. Riparian elements are also well represented in this interval. Zone OL-3. Zone OL-3 is characterized by common to abundant Dacrydium and Casuarina type pollen, with a strong acme of dinoflagellate cysts dominated by Operculodinium spp and Spiniferites spp marking the top of the zone but the low representation of Pandanus, Ilex and other riparian elements. Zone OL-4. Zone OL-4 is characterized by abundant Dacrydium and regular Casuarina type pollen, but the low representation of Pandanus, Ilex and other riparian elements and dinoflagellate cysts. Zone OL-5. Zone OL-5 is characterized by reduced Dacrydium and Casuarina type pollen, and increased riparian elements such as Ilex and Pandanus.

Climate history Zone OL-1 probably reflects a period with moderately wet climates since rain forest and peat swamp pollen dominate assemblages. Zone OL-2,

on the other hand, probably reflects a period of marginally seasonal climates. The assemblages through OL-2 are in fact mainly dominated by everwet rain forest climate elements, but the consistent occurrence of pollen of Poaceae, Schoutenia and of ‘Australian’ Austrobuxus/Dissiliaria (Malvacipollis diversus) and reduced representation of peat swamp elements suggests that vegetation characteristic of more seasonal climates must have had a distinct presence in the sediment source area. Within zones OL-3 and OL-4, the dominance of pollen of ‘Kerapah’ type peat swamp elements, such as Dacrydium and Casuarina type, and the disappearance of riparian elements (ecologically replaced by the Kerapah group) suggests a period of everwet to superwet climates. With the reduction of the Kerapah group in zone OL-5, and return of riparian elements, a change to a slightly less wet (but still everwet) climatic regime is indicated for the uppermost interval. It is because the palynological zones are characterized by elements suggesting regional climate change that the zones are thought to be of potentially wide correlatable significance.

Fig. 8. Climate sketches for zones OL-2, 3 and 4. superimposed on palaeogeographic maps of Hall 1998. Stipple, land area; grey shading, shelf; brick pattern, limestone platforms; triangles, volcanoes.

OLIGOCENE PALYNOLOGY EAST JAVA SEA

Discussion East Java wet climate province For most of the Oligocene in SE Asia seasonal climate assemblages are the rule (e.g. Morley 2000; Morley et al. 2003; Morley & Shamsuddin 2006) although essentially everwet climates have recently been interpreted for the Udang Formation in Block B, West Natuna (southernmost part of Natuna) by Morley et al. (2007). The wettest climates recorded for the Oligocene, however, are from the Java region and probably reflect a wet climate fringe to the eastern margin of Sundaland prior to the collision of the Australian and Asian plates at the Oligo-Miocene boundary. The period of wetter climate approximately coincides with the Late Oligocene thermal maximum as recorded from isotopes by Zachos et al. (2001). Possible

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palaeoclimate sketches for zones OL-2, and OL-3/4 are shown in Figure 8.

Correlation to Sunda Basin Zone OL-2, with regular Malvacipollis diversus, and regular seasonal climate elements, can be correlated directly to the ‘Freshwater’ Talang Akar and Cibulakan Formations in the Sunda Basin, with zones OL-3 and OL-4 corresponding to the ‘Coaly’ Talang Akar (Fig. 9). Zone OL-2 can be accurately dated using nannofossils to zone NP24 (27 –29 Ma) and this provides the first direct age attribution to the Freshwater Talang Akar/Cibulakan Formations in the Sunda/Arjuna region. The OL-2 seasonal climate event probably has a wide stratigraphic application across the Sunda region, and may correlate to palynological zones PIV– PV

Fig. 9. Correlation to Sunda Basin. See caption to Figure 4 for explanation of pollen groupings.

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Vietnam (Morley unpublished) characterized by wetter climates, but not other areas where climates were more seasonal (Morley 2010). A model to explain the dispersal of Casuarinaceae into the Sunda region remains unresolved, since dispersal via India is unlikely as there is no pollen record from the Indian subcontinent prior to the Miocene. Long distance dispersal may be a possibility. The authors are grateful to Anadarko Indonesia and MIGAS for permission to publish this paper. The foraminifera and nannofossils discussed in this paper were identified by biostratigraphers at LEMIGAS. Carlos Jaramillo and Margaret Collinson provided suggestions which have improved the manuscript.

References Fig. 10. Timing of dispersal of Dacrydium into SE and east Asia with arrows showing dispersal direction (from Morley 2010). Location of DSDP 214 and 254 shown as black circles, plate tectonic reconstruction 30 Ma from Lawver et al. 2000.

(Morley et al. 2003) in Natuna Basin and zones PR1–4 in the Malay Basin (Azmi et al. 1996).

Occurrence of pollen of Dacrydium and Casuarina type in the Oligocene of East Java This study extends the record of Dacrydium and Casuarina type pollen from the Oligo-Miocene boundary (Morley 2000) to the base Oligocene, previously brought to attention by Lelono (2007). Previously, it was thought that these two taxa dispersed into the Sunda region following the collision of the Australian and Asian plates, at the Oligo-Miocene boundary (Morley 1998, 2000), but this study suggests that this is unlikely, since at the time of the basal Oligocene, when these pollen types first appear, the Australian land mass would have been some 1000 km south of the East Java area. Recently, additional records of Dacrydium pollen have been reported from the Eocene of the Indian subcontinent (Morley 2011) and it can now be proposed that Dacrydium dispersed into SE Asia prior to the Early Oligocene via the NinetyEast Ridge (where it has been reported from the Paleocene by Kemp & Harris (1975) and the Indian Plate (Fig. 10). Its distribution across the Sunda region and Indochina is thought to be limited by palaeoclimate, explaining why it is present in some areas of the Sunda region and in the SE Asian mainland, in the Oligocene of Southern China (Sun et al. 1981) and Late Oligocene of

Adams, C. G. 1970. A reconsideration of the East Indian Letter classification of the Tertiary. Bulletin of the British Museum Natural History (Geology), 19, 87–137. Anderson, J. A. R. 1963. The flora of the peat swamp forests of Sarawak and Brunei, including a catalogue of all recorded species of flowering plants, ferns and fern allies. Gardens Bulletin, Singapore, 20, 131– 238. Azmi, M. Y., Awalludin, H., Bahari, M. N. & Morley, R. J. 1996. Integrated biostratigraphic zonation for the Malay Basin. Geological Society of Malaysia, 39, 157– 184. Blow, W. H. 1969. Late Middle Eocene to Recent planktonic foraminiferal biostratigraphy. In: Proceedings First International Conference on Planktonic Microfossils, Geneva 1967, 1, 199 –422. Blow, W. H. 1979. The Cainozoic Globigerinida: a study of the morphology, taxonomy, evolutionary relationships and stratigraphical distribution of some Globigerinida (mainly Globigerinacea). E.J.Brill, Leiden. Bolli, H. M. 1966. The planktonic foraminifera in the well Bojonegoro-1 of Java. Eclogicae Geologicae Helveticae, 59, 449–265. Bolli, H. M. & Saunders, J. B. 1985. Oligocene to Holocene low latitude planktic foraminifera. In: Bolli, H. M., Saunders, J. B. & Perch-Nielsen, K. (eds) Plankton Stratigraphy. Cambridge University Press, Cambridge, 155–262. Brunig, E. F. 1990. Oligotrophic forested wetlands in Borneo. Chapter 13. In: Lugo, A. E., Brinson, M. & Brown, S. (eds) Ecosystems of the World. Elsevier, Amsterdam, Forested wetlands, 15, 299– 344. Cookson, I. C. & Pike, K. M. 1954. Some dicotelydonous pollen types from Cainozoic depoits in the Australian region. Australian Journal of Botany, 2, 197– 219. Firdaus, I., Soeka, S., Irwansyah, & Prayitno, I. 2004. The application of quantitative biostratigraphy for identifying stratigraphic traps in East Java BlockMadura Strait. Report of Lemigas, Jakarta, In-house Research. Gonzalez Guzman, A. E. 1967. A palynological study of the Los Cuervos and Mirador Formations, Lower and Middle Eocene, Tibu area, Colombia. E.J. Brill, Leiden.

OLIGOCENE PALYNOLOGY EAST JAVA SEA Gradstein, F. M., Ogg, J. G. et al. 2004. A Geologic Time Scale 2004. Cambridge University Press, Cambridge. Hall, R. 1998. The plate tectonics of Cenozoic SE Asia and the distribution of land and sea. In: Hall, R. & Holloway, J. D. (eds) Biogeography and Geological Evolution of SE Asia. Backhuys Publishers, Leiden, The Netherlands, 98–131. Hammen, T. & van der, 1954. El desarollo de la flora Colombiana en los periodos geological I. Maestrichtiano hasta Terciario mas Inferior. Boletin Geologico, 2, 49–106. Haq, B. U., Hardenbol, J., Vail, P. R. & Baum, G. R. 1988. Mesozoic and Cenozoic chronostratigraphy and eustatic cycles. In: Wilgus, C. K., Hastings, B. S., Posamentier, H. & Van Wagoner, J. (eds) Sea Level Change: An Integrated Approach. Society of Economic Palaeontologists and Mineralogists, Tulsa, OK, Special Publication, 42, 71– 108. Harris, W. K. 1965. Basal Tertiary floras from the Princetown area, Victoria, Australia. Palaeontographica B, 115, 75–106. Johansen, K. B. 2003. Depositional geometries and hydrocarbon potential within Kujung carbonates along the Madura Platform, as revealed by 3D and 2D seismic data. In: Proceedings Indonesian Petroleum Association, 29th Annual Convention, Jakarta, IPA-3-G-174. Kemp, E. M. & Harris, W. K. 1975. The vegetation of Tertiary islands on the Ninetyeast Ridge. Nature, 258, 303– 7. Lawver, L. A., Coffin, M. F., Gahagan, L. M., Campbell, D. A. & Royer, J.-Y. 2000. Opening of the Indian Ocean. University of Texas Institute for Geophysics, Houston, TX. Lelono, E. B. 2007. Gondwanan Palynomorphs from the Paleogene Sediment of East Java:? The Evidence of Earlier Arrival. In: Proceeding of Indonesian Geologist Association (IAGI), 36th Annual Convention, Bali. Leupold, W. & van der Vlerk, I. M. 1931. The tertiary. In: Escher, B. G. et al. (eds) Stratigraphie van Nederlandsch Oost-Indie. Feestbundel Martin, Leidsche Geologische Mededelingen, 5, 611– 648. Martin, H. A. 1974. The identification of some Tertiary pollen belongong to the family Euphorbiaceae. Australian Journal of Botany, 22, 271–291. Martini, E. 1971. Standard Tertiary and Quaternary calcareous nannoplankton zonation. In: Farinacinni, A. (ed.) Proceedings 2nd Plankton Conference, Rome, 1969. 739– 785. Morley, R. J. 1978. Palynology of Tertiary and Quaternary sediments in southeast Asia. In: Proceedings of the 6th Annual Convention, Indonesian Petroleum Association, May 1977, 255– 276. Morley, R. J. 1991. Tertiary stratigraphic palynology in South-East Asia; current status and new directions. Proceedings of the Geological Society of Malaysia, 28, 1– 36. Morley, R. J. 1998. Palynological evidence for Tertiary plant dispersals in the Southeast Asian region in relation to plate tectonics and climate. In: Hall, R. & Holloway, J. D. (eds) Biogeography and Geological Evolution of SE Asia. Backhuys, Leiden, 211–234.

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Morley, R. J. 2000. Origin and Evolution of Tropical Rain Forests. Wiley & Sons, London. Morley, R. J. 2011. Palaeoecology and dispersal of tropical Podocarps. In: Turner, B. L. & Cernusak, L. M (eds) Ecology of the Podocarpaceae in Tropical Forests. Smithsonian Institute Scholarly Press, Washington, DC, Smithsonian Contribution to Botany, Vol. X. Morley, R. J. & Shamsuddin, J. 2006. The Sequence Biostratigraphy and Chronostratigraphy of the Malay Basin. In: PGCE Malaysia Proceedings, 77–78. Morley, R. J., Morley, H. P. & Restrepo-Pace, P. 2003. Unravelling the tectonically controlled stratigraphy of the West Natuna Basin by means of palaeoderived Mid Tertiary climate changes. In: 29th Indonesian Petroleum Association Proceedings, volume, 1– 24. Morley, R. J., Salvador, P., Challis, M. L., Morris, W. R. & Adyaksawan, I. R. 2007. Sequence biostratigraphic evaluation of the North Belut Field West Natuna Basin. In: Proceedings Indonesian Petroleum Association, 31st Annual Convention, Jakarta, IPA07-G-120. Perch-Nielsen, K. 1985. Cenozoic calcareous nannofossils. In: Bolli, H. M., Saunders, J. B. & PerchNielsen, K. (eds) Plankton Stratigraphy I. Cambridge University Press, Cambridge, 427–554. Pickett, E. J., Harrison, S. P. et al. 2004. Pollen-based reconstructions of biome distributions for Australia, Southeast Asia and the Pacific (SEAPAC region) at 0, 6000 and 18,000 14C yr BP. Journal of Biogeography, 31, 1381– 1444. Pusoko, S., Sofyan, S., Rahardjo, K. & Endarto, M. 2005. Hydrocarbon Evaluation of North East Java (Madura). Report of Lemigas In-house Research, Lemigas, Jakarta. Richards, P. W. 1996. The Tropical Rain Forest, 2nd edn. Cambridge University Press, Cambridge. Stover, L. E. & Evans, P. R. 1973. Upper Cretaceous – Eocene Spore– Pollen Zonation, Offshore Gippsland Basin, Australia. Geological Society of Australia, Sydney, Special Publication, 4, 55– 72. Sun, X. I., Li, M. X. et al. 1981. Palynology Section in: Tertiary of the Northern Continental Shelf of the South China Sea. Guandong Technology Press, China. Vail, P. R., Mitchum Jr., R. M. & Thompson III, S. 1977. Global cycles of relative changes of sea level. In: Payton, C. E. (ed.) Seismic Stratigraphy: Applications to Hydrocarbon Exploration. American Association of Petroleum Geologists Memoir, 26, 83–98. Van der Vlerk, I. M. & Umbgrove, J. H. F. 1927. Tertiaire idsforaminiferen van Nederlandsch Oost Indie. Wetensch. Meded. Dienst Mijnbouw Nederl. Oost-Indie, 6, 1– 45. Whitmore, T. C. 1998. An Introduction to Tropical Rain Forests, 2nd edn. Oxford University Press, Oxford. Yulianto, Martodjojo, & Zaim, 2000. The geology of Java. In: Herman, D. & Hasan, S. (eds) An Outline of the Geology of Indonesia. Ikatan Ahli Geologi, Indonesia, 192, 45– 68. Zachos, J. C., Pagani, M., Sloan, L., Thomas, E. & Billups, K. 2001. Trends, rhythms and aberrations in global climate, 65 Ma to present. Science, 292, 686– 693.

SE Asian carbonates: tools for evaluating environmental and climatic change in equatorial tropics over the last 50 million years MOYRA E. J. WILSON Department of Applied Geology, Curtin University, GPO Box U1987, Perth, Western Australia, 6845 (e-mail: [email protected]) Abstract: This study reviews how shallow water carbonates are revealing environmental and climatic changes on all scales through the last 50 million years in SE Asia. Marine biodiversity reaches a global maximum in the region, yet the environmental conditions are at odds with the traditional view of ‘blue-water’ reefal development. The region is characterized by complex tectonics, major volcanism, high terrestrial runoff, nutrient influx, everwet and monsoonal climates, low salinities, major currents and ENSO (El Nin˜o Southern Oscillation) fluctuations. Terrestrial runoff, nutrient upwelling, tectonics, volcanism and recent human activities are major influences on the modern development of carbonate systems. Coral sclerochronology is revealing how these factors vary locally over annual and decadal scales. The strong impact of vertical tectonic movements and the interplay with eustasy is evaluated from Quaternary and Pleistocene coral reef terraces. Isotopic data from terrace deposits indicates that interglacials may have been up to 3– 6 8C warmer than glacials, consistent with the region’s record from terrestrial and deep marine deposits. Study of outcrop and subsurface carbonate deposits reveals the impact of tectonics, siliciclastic, nutrient influx, eustasy and oceanography on individual systems over millennial timescales. Major changes in oceanography, plate tectonics, climate change and perhaps fluctuating CO2 levels impacted Cenozoic regional carbonate development. Results of studies from terrestrial and deep marine realms are comparable with those from the carbonates, but have yielded higher resolution records of changing currents, precipitation and the monsoons. There is considerable scope for further research, however, SE Asian carbonates are powerful tools in evaluating past environmental change in the equatorial tropics.

There is intense debate over the role of the equatorial tropics as a participant in natural or human-influenced environmental change, and/or as a major driver of global change (Kerr 2001; Pearson et al. 2001). SE Asia is renowned as the region of highest modern global reefal biodiversity, with an extensive and continuous geological record of reefal and non-reefal carbonates spanning the last 50 million years. Carbonate rocks form through the accumulation of calcium carbonate (CaCO3) sediments, most commonly in shallow waters from the skeletons of marine organisms, such as corals or foraminifera. The biota, textures and geochemical signatures within these deposits respond directly to variations in environmental and climatic conditions (Montaggioni & MacIntyre 1991). The hypothesis here is that SE Asian carbonates can be used as tools in evaluating the unique but poorly understood conditions of the equatorial tropics and its past environmental changes together with the response of marine systems. This paper reviews how carbonates have been, and may continue to be, used to evaluate environmental change, on annual to millennial scales, over the last 50 million years in SE Asia (Fig. 1). The mention of coral reefs and other tropical carbonates generally conjures up images of a

myriad of brightly coloured creatures shimmering and darting through crystal-clear, warm, shallow waters. One can be mesmerized by just such a scene on days of ‘glass-flat’ calm in some parts of SE Asia. However, due to the unique regional conditions it is not uncommon to find yourself shivering in a wet suit, peering through turbid waters or ‘plankton-soup’ trying to make out the inhabitants of the reef. The equatorial tropics are characterized by a range of conditions that are commonly underappreciated and may at first appear incompatible with the high diversity of carbonate biota and systems found in the region (Figs 2 & 3; Fulthorpe & Schlanger 1989; Tomascik et al. 1997; Wilson 2002, 2008; Park et al. 2010). SE Asia has been arguably the most tectonically active area of the world throughout much of the Cenozoic (Hall 1996, 2002). This complex tectonic setting together with high rainfall and lush tropical vegetation results in common influx of volcaniclastics, siliciclastics, fresh water and nutrients into the coastal water of the region (Figs 2 & 3; Tomascik et al. 1997; Wilson & Lokier 2002). A locally monsoonal climate causes strongly seasonal terrestrial runoff together with shifts in wind and current patterns (Umbgrove 1947; Park et al. 2010). The region lies outside the cyclone belt and strong

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonics of the Australia– Asia Collision. Geological Society, London, Special Publications, 355, 347–372. DOI: 10.1144/SP355.18 0305-8719/11/$15.00 # The Geological Society of London 2011.

General SE Asia Modern

Modern Makassar Straits

Burollet et al. (1986) Roberts et al. (1988), Tomascik et al. (1997), Roberts & Sydow (1996) Renema & Troelsta (2001), Wilson & Vecsei (2005)

Java & Sumatra modern deposits

Umbgrove (1939), Scoffin et al. (1989), Gagan et al. (1998), Park et al. (1992, 2010), Risk et al. (2003), Abram et al. (2003), Baird et al. (2005) Campbell et al. (2007), Hagan et al. (2007);

Huon Peninsula & offshore N. Papua New Guinea

Chappell (1974), Chappell & Polach (1976), Chappell et al. (1996), Galewsky et al. (1996), McCulloch et al. (1996, 1999) , Webster et al. (2004a, b), & others in text

Papua New Guinea modern deposits

Eastern Indonesia coral reef terraces Pulau Seribu, NW Java & Java Sea Carbonates, Batu Putih Limestone & Mahakam Delta associated carbonates, E. Borneo

Atmospheric CO2

Biotic evolution

Monsoons Indonesian Throughflow and global ocean currents ENSO-like fluctuations

Eustasy

Glacial-interglacial or greenhouse-icehouse Tectonics & volcanism

Currents

Temperature

Siliciclastic or volcaniclastic influx

Salinity & Precipitation

Tectonics

Relative sea-level change

Volcanic/hydrothermal activity

Regional or global changes:

Dating

U/Th, C14

Chappell & Veeh (1978), Sumosusastro et al. (1989), Pirazzoli et al. (1993), Hantoro et al. (1994), Bard et al. (1996) Burbury (1977) , Park et al. (1992, 2010), Carter & Hutabarat (1994) Roberts & Sydow (1996), Wilson & Lokier (2002), Wilson (2005), Lokier et al. (2009), Saller et al. (2010)

Nutrients

Local environmental factors: Subsurface–Seismic (S), Wireline (W)

Sedimentology

Geochemistry

Biota

Basin

Type of study:

Platform

Strata

Biota

Cenozoic

Palaeogene

Neogene

Recent (Holocene & Quaternary)

Studied:

M. E. J. WILSON

Banda Sea modern deposits

Brown & Holley (1982), Tudhope & Scoffin (1994), Brown (2005), Phongsuwan & Brown (2007); Scoffin et al. (1989), Tudhope et al. (1995), Fallon et al. (2002), Aycliffe et al. (2004) Pichler et al. (2000) Wallace (1869), McManus & Wenno (1981), Best et al. (1989), Heikoop et al. (1996), Tomascik et al. (1997)

Thailand modern deposits

Cenozoic scales (107–108 years)

Key References Kuenen (1933), Umbgrove (1947), Tomascik et al. (1997), Kleypas et al. 1999, Latief (2000), Montaggioni (2005), Gordon, 2005, Stoddart, 2007, Peñaflor et al. (2009)

Annual to centennial (1–102 years) Thousand to hundred thousand years (103–105 years) Millennial scales (106 years)

Study (stratigraphic formation and/or area)

Period:

348

Timescales:

S

U/Th, ESR C14, (Seismic picks) Seismic picks, Bio: LBF, Pk, Nanno, U/Th

Fig. 1. Study areas and key references in which local environmental change and regional or global controls have been inferred from carbonate deposits in SE Asia. Also shown are the timescales of change, the period of study, what elements of the carbonate system have been studied and the main types of analyses. Key references and examples are given in the text, and all reference details are provided (see Wilson 2002, 2008 for further details). Abbreviations for the dating techniques are: bio, biostratigraphy; LBF, Larger benthic foraminifera; pk, planktonic foraminifera; nanno, nannofossils; Sr, strontium isotope stratigraphy; C14, carbon 14; U/Th, uranium thorium; ESR, electron spin resonance.

Ardila (1983), Rose (1983), Collins et al. (1996), Park et al. (1995) May & Eyles (1985), Rudolph & Lehmann (1989), Bachtel et al. (2004), Paterson et al. (2006), Epting (1980), Ali & Abolins (2000), Vahrenkamp et al. (2004), Zampetti et al. (2003, 2004) Lokier (2000), Wilson & Lokier (2002), Lokier et al. (2009)

Fig. 1. Continued.

Bio: LBF, Pk, Nanno

Seismic picks, Sr, Bio: LBF

Bio: LBF, Pk, Nanno

Francis (1988), Stewart & Sandy (1988), Wilson et al. (1993)

Bio: LBF, Nannos Seismic picks, Sr, Bio: LBF, Pk, Nannos Bio: LBF, Pk, Nanno Seismic picks, Sr, Bio: LBF, Nannos Bio: LBF, Pk, Nanno Bio: LBF, Pk, Nanno

Seismic picks, Sr, Bio: LBF, Pk

Seismic picks Seismic picks

Seismic picks, Sr Bio: LBF, Pk, Nanno Bio: LBF, Pk, Nanno Seismic picks, Bio: Pk, Nannos

Seismic picks

Bio: LBF, Pk, Nanno

S

S

S

Seismic picks, Bio: Pk

Wilson & Bosence (1996), Wilson (1999, 2000), Wilson et al. (2000) Jurgan & Domingo (1989), Muller et al. (1989), Porth et al. (1989) Vincelette (1973), Redmond & Koesoemadinata (1976), GibsonRobinson & Soedirdja (1986), Brash et al. (1991)

et al. (2004, 2005)

Adams et al. (1965), Wannier (2009) Grötsch & Mercadier (1999), Fournier

Saller et al. (1993), Burollet et al. (1986), Saller & Vijaya (2002)

Wilson et al. (1999), Wilson & Evans (2002)

Kenyon (1977), Cucci & Clark (1993, 1995), Kusumastuti et al. (2002), Adhyaksawan (2003), Johansen (2003), Carter et al. 2005, Sharaf et al. (2005), Hughes et al. (2008), Ruf et al. (2008)

Pieters et al. (1983), Leamon & Parsons (1986), Pigram et al. (1990), Sarg et al. (1995), Eisenberg New Guinea/Darai Limestone & equivalents, et al. (1996), Allan et al. (2000), Tcherepanov et al. (2008a, b) Papua & Gulf of Papua Area Fulthorpe & Schlanger (1989), Wilson & Rosen (1998), Wilson (2002, , 2008), Wilson & Vecsei (2005), Wilson & Hall (2010), Park et al. (2010) Regional SE Asia

New Guinea Limestone & equivalents, Irian Jaya New Guinea Limestone equivalents in northern Papua and islands to NE of New Guinea

Visayas, Central Philippines

Berai Limestone, Paternoster Platform, SE Borneo Melinau, & other, Limestone(s), Onshore N Borneo Nido Limstone, Offshore Palawan Philippines Tonasa Formation, South Sulawesi, Indonesia

NE Java & E Java Sea Carbonates, Ngimbang, Prupuh, Rancak & Kujung Carbonates NE Kalimantan, Taballar & Kedango Limestones, Indonesia

Tacipi Formation, South Sulawesi, Indonesia Mayall & Cox (1988), Ascaria (1997) Matthews et al. (1997), Mayall et al. (1997) Vietnam Carbonates Ziujiang Carbonates, Liuhua & northern S. Erlich et al. (1990, 1993), Tyrrel & Christian (1992), Sattler et al. (2004) China Sea Carbonates Bali Flores Sea Tyrrel et al. (1986)

Wonosari Limestone, S. Java, Indonesia

Luconia Carbonates, S. China Sea

Terumbu Formation, Natuna, S. China Sea

Backarc Carbonates of Sumatra, Peutu, Batu Raju Formations

CARBONATES AND ENVIRONMENTAL CHANGE 349

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M. E. J. WILSON

Fig. 2. Map of SE Asia showing present-day plate tectonic context, locations of volcanoes and sediment discharge from six major SE Asian islands (Sumatra, Java, Borneo, Sulawesi, Timor, New Guinea; from Milliman et al. 1999 with base map after Wilson & Lokier 2002).

cyclonic winds and waves are rare (Umbgrove 1947; Tomascik et al. 1997). Tectonic faulting, subsidence and uplift combined with glacioeustasy control localized relative sea level changes that influence carbonate deposition, coral reef growth, their subaerial exposure and flooding (Wilson 2002, 2008; Park et al. 2010; Wilson & Hall 2010). Within the region volcanism, seismic activity and associated tsunami’s cause major environmental change to both land-, and seascapes. On the short term these wreak devastation on communities, but longer term may bring ecological opportunities (Wilson & Lokier 2002; Stoddart 2007; Satyana 2005; Pandolfi et al. 2006). SE Asia is now the last remaining equatorial ‘oceanic gateway’ allowing interchange of oceanic waters between the Pacific and Indian Oceans via the major Indonesian Throughflow Current (Fig. 3; Gordon 2005). The region’s climate and current systems are influenced by, and/or interact, with global ocean and atmospheric phenomena including the El Nin˜o Southern Oscillation (ENSO), Indian Ocean Dipole (IOD),

fluctuations in the monsoons and the Inter-Tropical Convergence Zone (ITCZ; Tudhope et al. 2001; Kuhnt et al. 2004; Wang et al. 2005; Abram et al. 2009). These factors are influential on annual to millennial scales in changing sea surface temperatures to both locally warmer and cooler than ambient (Gagan et al. 1998; Pen˜aflor et al. 2009). Nutrient influx and areas of upwelling are also affected, and in turn cause changes in water clarity associated with plankton blooms (Fig. 3; Gagan et al. 1998; Wilson & Vecsei 2005). Longer term oceanographic (temperature, acidity and compositional changes) and atmospheric (CO2) changes over the scale of the Cenozoic (Zachos et al. 2001; Jia et al. 2003; Pagani et al. 2005) during the switch from greenhouse to icehouse climatic states are also major influences on the marine biota and systems of the region (Wilson 2008). Both the long and shortscale changes as well as influencing marine systems are now known to be major drivers in global climate change (Gordon et al. 2003; Visser et al. 2004).

Fig. 3. Maps of SE Asia showing present day environmental conditions. (a) Satellite (AQUA-Modis) derived sea surface temperatures during the SE Monsoon (from Gordon 2005). Wind directions are shown with arrows and the position of the Inter-Tropical Convergence Zone as a dashed line. (b) Sea surface chlorophyll-a (a nutrient proxy) from satellite-derived SeaWIFs data during the SE Monsoon (from Gordon 2005). Areas of wind induced upwelling are associated with elevated levels of chlorophyll-a. (c) Indonesian throughflow pathways and estimates of total volume transport (in Sverdrups Sv ¼ 106 m3 s21; from Gordon 2005). (d) Annual mean rainfall across the globe with data derived from Special Sensor Microwave Imager (SSM/I). Satellite data collection began in 1987 NOAA data from Sidi et al. 2003). (e) Present-day sea surface aragonite saturation at 380 ppm atmospheric CO2 (from Hoegh-Guldberg et al. 2007). The minimum aragonite saturation that coral reefs are associated with today is 3.25. Aragonite saturation in

CARBONATES AND ENVIRONMENTAL CHANGE

351

Fig. 3. (Continued) seawater (Varg) is the ion product of the concentrations of calcium and carbonate ions, at the in situ temperature, salinity,     and pressure, divided by the stoichiometric solubility product (K*sp) for those conditions: Ksparg . Varg ¼ Ca2þ CO2 3

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Fig. 4. Examples of annual to centennial scale environmental changes from modern deposits. (a) Comparison between PNG rainfall records (Port Moresby) with fluorescence banding pattern from offshore slabbed coral (from Scoffin et al. 1989). Fluorescent banding correlates with periods of runoff. (b) Modern reefal deposit from turbid, sedimentinfluenced area in 1–2 m depth, east coast of Borneo near Balikpapan (photograph by M. Wilson). (c) Paired

CARBONATES AND ENVIRONMENTAL CHANGE

The combination of major tectonism, frequent relative sea level changes, low marine salinities, clastic and nutrient influx, changing oceanographic and temperature conditions all strongly influence regional and local carbonate development. These conditions may be at odds with the perceived view of ideal conditions for coral reefs and tropical carbonate production principally developed from studies in warm, more arid, subtropical regions such as the Bahamas, Red Sea or Persian Gulf (Wilson 2002). Although no individual factor is mutually exclusive to SE Asia, it is the unique combination of factors that results in the distinctiveness of equatorial carbonates. This paper assesses to what extent the carbonate record from SE Asia has been, and may continue to be, used to evaluate environmental conditions and changes in the equatorial tropics.

Annual to centennial-scale changes The majority of annual to centennial-scale changes in SE Asian shallow carbonates are inferred from geochemical proxies obtained from modern to sub-recent unaltered marine skeletons. The most widely used record is from corals, where unaltered samples typically extend back tens, hundreds or perhaps thousands of years. Coral sclerochronology studies in the region have mainly concentrated on Porites (particularly Porites lutea and lobata), which have dense skeletal structures and for which fractionation effects in skeletal precipitation are studied (i.e. if there are systematic variations in skeleton chemistry compared with ocean water chemistry). Annual and decadal variations have been compared with results of recent monitoring of factors including temperatures, terrestrial runoff, currents and nutrients (Fig. 4; Scoffin et al. 1989; Gagan et al. 1998; Fallon et al. 2002). In turn these have been variously linked to changes in upwelling, strength of the monsoons, ENSO fluctuations as well as volcanic and anthropogenicinduced events (Fig. 4; Tudhope et al. 1995; Heikoop et al. 1996; Tomascik et al. 1997; Gordon 2005; Pen˜aflor et al. 2009). In an early study, rainfall data and distance from the coastline were compared with density of skeletal growth and fluorescence bands from corals off

353

north Java (Pulau Seribu) and Papua New Guinea (Fig. 4a; Scoffin et al. 1989). Heavy wet season deluges (.150 mm/month) correlate with less dense growth bands and distinctive bright fluorescence banding in offshore corals. In nearshore areas (,2–5 km for the coast) sufficient influx of organic compounds appears to swamp any seasonal effects (Scoffin et al. 1989). Fluorescence intensity decreases offshore, presumably reflecting decreasing freshwater and nutrient influence (Scoffin et al. 1989). Seaward spread of freshwater plumes, was more localized off Java than Papua New Guinea, and may be modified by marine currents or local runoff from islands. Strong monsoonal driven currents are widely known to result in east –west elongation of reefal buildups in the Java Sea at the present day (Park et al. 1992, 2010) and in the past (Carter & Hutabarat 1994). The shape of emergent coral reef islands strongly affects monsoonally influenced circulation on carbonate buildups and the dynamics of reef island shorelines (with circular islands most heavily impacted; Kench et al. 2009). Increased growth (low density skeleton) during freshwater and/or nutrient influx and no or low growth during exceptionally high runoff events (Scoffin et al. 1989) is seen in other corals from the region (Tomascik et al. 1997), but is at odds with some studies from outside the equatorial tropics (Isdale 1984; Tomascik & Sanders 1985). Although high sedimentation does limit coral and carbonate production within the region, there are many examples where diverse modern communities are present in areas of sedimentation, turbidity, upwelling and nutrient influx (Fig. 4b; e.g. Banten Bay (Java), Seram, Borneo, Ambon; Tomascik et al. 1997; Rosen et al. 2002; Wilson & Lokier 2002; Wilson 2005). In these areas the depth of abundant coral or larger benthic foraminifera may be strongly depth and light limited (Titlyanov & Latypov 1991; Renema & Troelsta 2001). In addition many organisms are sediment or nutrient adapted, and filter feeders or deposit feeders are common and diverse (Tomascik et al. 1997). The Makassar Straits has a wide range of examples of carbonate systems influenced by clastic sedimentation and nutrients. In the north, patch reefs with abundant soft corals, detritivores and foliaceous hard corals develop to 5–8 m water depth in highly turbid waters just 8 km from the

Fig. 4. (Continued) X-radiograph (A) and slabbed coral (B) from offshore Banda Api. Dashed line and arrow show the position of iron-rich precipitate related to volcanic and hydrothermal activity associated with the 1988 eruption of Banda Api Volcano (from Heikoop et al. 1996). (d) Comparison between coral Sr/Ca in Porites lutea and blended ship and satellite-derived sea surface temperature data from SE Java (from Gagan et al. 1998). Strong upwelling along the south coast of Java produces cooler SSTs recorded in the Sr/Ca ratios, but not recorded by the satellite data due to spatial smoothing of data (Gagan et al. 1998). (e) Stable oxygen isotope record from coral core from PNG (Madang) compared with climatic indices. Years with strong El Nin˜o events are shown stippled (from Tudhope et al. 1995).

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Berau Delta (Tomascik et al. 1997). At Pulau Sangalaki 60 km east of the Berau Delta strong tidal-currents keep siliciclastic sedimentation to a minimum but bring nutrient-rich coastal waters that support a highly diverse reef community (Tomascik et al. 1997). Seaward of the Mahakam Delta Plio-Pleistocene carbonate buildups and reefs are more common to the north where a less active delta lobe and strong south-directed currents from the Indonesian Throughflow Current result in less turbid-water and sedimentation than to the south (Roberts & Sydow 1996). Halimeda buildups form in nutrient-rich, current swept or deltainfluenced areas of the southern Makassar Straits (Roberts et al. 1988; Roberts & Sydow 1996). Also in the south, large land attached platforms influenced by strong currents, nutrients & clastics are dominated by larger benthic foraminifera below 5–20 m with shallow coral rimmed margins or patch reefs (Burollet et al. 1986; Renema & Troelsta 2001; Wilson & Vecsei 2005). In general the marine systems of western SE Asia are strongly runoff influenced and those of the eastern archipelago have a major upwelling influence (Tomascik et al. 1997). Rather than falling into the traditional view of ‘blue water’ oceanic reefs, most living Indo-Pacific reefs (53% and much of the 20% from the Indian Ocean) are concentrated on the shallow continental shelves of SE Asia, Australia and the Indian Ocean where they are affected to a greater or lesser extent by adjacent land masses (Potts 1983; Tomascik et al. 1997). Despite the co-occurrence of high diversity coral reefs and carbonate production in areas of runoff and nutrient upwelling there are signs that some systems may be operating close to their tolerance limits (Tomascik et al. 1997). In some areas this may be a natural state (Tomascik et al. 1997) with examples of natural mass mortality and demise or subsequent recovery of reefs from the region’s geological record associated with volcanism or sedimentation and/or nutrients (Wilson & Lokier 2002; Wilson 2005; Pandolfi et al. 2006; Lokier et al. 2009). It has been inferred that human activities (including destructive fishing practices, pollution and coastal development) have significantly increased the frequency of disturbance to reefs and have pushed reefs closer to their tolerance limit (Hughes et al. 2003; Pandolfi et al. 2003, 2006; Wilkinson 2008). Umbgrove (1939) described the ‘unrivalled splendour and wealth of . . . reef animals’ after a visit in 1928 to an island 6.6 km away from Jakarta. These coral reefs that dazzled Umbgrove are now considered ‘functionally dead’ replaced by muddy deposits (Tomascik et al. 1997). In a comparable example Wallace (1869) portrayed the ‘most astonishing and beautiful . . . continuous series of corals . . . and other marine

productions, of magnificent dimensions, varied forms and brilliant colours’ from Ambon Harbour. Today most coral communities in the vicinity of Ambon City have been destroyed through a combination of pollution, siltation and destructive fishing practices (McManus & Wenno 1981; Best et al. 1989; Tomascik et al. 1997). A range of studies have shown corals can be excellent proxy recorders of localized natural or human activity, including volcanism, hydrothermal venting, mining activity and forest fires. Within the growth bands of corals from the flanks of the volcanic island of Banda Api are hydrothermalassociated iron-rich lamina, tuffaceous material and death surfaces that correlate with the 1988 eruption of the volcano (Fig. 4c; Heikoop et al. 1996). Significantly negative d13C and d18O isotopic records in coral growth bands from areas of hydrothermal venting off Papua New Guinea reflect higher than background temperatures and fractionation effects (Pichler et al. 2000). Major fires across Indonesia in the El Nin˜o drought year of 1997 are associated with negative shifts in d13C records from coral growth bands that may reflect a shift to more heterotrophic feeding mechanisms (Risk et al. 2003). Iron fertilization from the 1997 Indonesian wildfires together with anomalous Indian Ocean Dipole upwelling and a giant red tide are inferred to have caused demise of a coral reef in the Mentawai Islands south of Sumatra (Abram et al. 2003). Growth bands from corals in the area that survived this event record evidence for cooling associated with upwelling (Sr/Ca) followed by a sharp increase in d13C associated with increased Mn, La and Y, the later interpreted as evidence for a large phytoplankton bloom (Abram et al. 2003). The start of open-cast gold mining and associated sediment runoff from Misima Island (Papau New Guinea – PNG) correlate with dramatic increases in the elements Y, La and Ce in coral growth bands. Although fluxes of the above elements ceased after mining activities finished, zinc and lead continued to be transported to the offshore reef via sulphate-rich waters (Fallon et al. 2002). Adjacent to a tin smelting site in Thailand heavy metal concentrations were noted in bivalves and alga, but not corals, perhaps because the metals are mainly detrital (Brown & Holley 1982). The examples above infer individual causal mechanisms for changes in coral development and their skeletons. However, in practice, the task of attributing changes to specific causal mechanisms is not always straightforward (B. R. Rosen pers. comm. 2010). The enormity of the human tragedy and coastal devastation of the 2004 Boxing Day Tsunami brought into sharp global focus the destructive force of tsunamis (Borrero 2005; Liu et al. 2005; Stoddart 2007; Spencer 2007). However, in the

CARBONATES AND ENVIRONMENTAL CHANGE

marine environment, from Sumatra and around the Indian Ocean margins, .50% of reefs in tsunami affected areas showed minimal impact and ,15% showed major impact (in which .50% of hard corals were affected; Brown 2005; Hagan et al. 2007; Phongsuwan & Brown 2007). Most impacted reefs were likely to recover within 5 –10 years (Brown 2005; Hagan et al. 2007) and broad sand sheets spread by tsunami waves were partially eroded and bioturbated within 1–2 years (Kench et al. 2007; Nichol & Kench 2008). It has been stated that ‘the damage caused to coral reefs by the December 26 earthquake and tsunami was rarely of ecological significance and in northern Aceh (Sumatra), tsunami damage was trivial when compared with that from chronic human misuse (including bombing, cyanide fishing and anchor damage)’ (Baird et al. 2005; Campbell et al. 2007; Hagan et al. 2007). Close to the epicentre of the earthquake generating the tsunami the most lasting associated effects were: (i) ‘uprooting’ of massive coral colonies from unstable substrate at depths .2 m (Baird et al. 2005; Campbell et al. 2007), (ii) boulder fields (with less than 7% directed onshore; Paris et al. 2010), (iii) post-tsunami burial of reefs by sediment (Baird et al. 2005), and (iv) localized uplift, emergence (by 1–2 m) or subsidence of reef flats (Zachariasen et al. 1999; Borrero 2005; Hagan et al. 2007). Although Indonesia with its significant tectonic instability has at least 105 historical records of tsunamis between 1699 and 1990 (Latief 2000), it seems likely that evidence for these or past tsunami activity may be difficult to discern from the geological record. Regional and long-term monitoring and sclerochronology studies are beginning to pick up repeated oceanographic or climatic events and trends across the SE Asian seas. A thirty year study (1985–2006) of satellite-derived sea surface temperatures (SST) reveals an average warming rate of 0.2 8C/decade for the region (Pen˜aflor et al. 2009). Warming is regionally non-uniform and was greatest in the north and east, with shortterm warming and coral bleaching events tied to La Nin˜a years. There was .1 8C of cooling during the 1991 Mt Pinatubo eruption (Pen˜aflor et al. 2009). The inner seas of Indonesia appear to afford some natural protection from warming events, perhaps related to complex geometries and strong current flushing (Pen˜aflor et al. 2009). Areas of upwelling experience cooler SST with seasonal temperature fluctuations recorded in Sr/Ca ratios in corals (such as offshore South Java; Fig. 4d; Gagan et al. 1998). Decreased freshwater runoff (inferred from coral d18O, d13C and UV fluorescence) and slightly decreased SST (0.5–1 8C from Sr/Ca and or d18O) are associated with El Nin˜o events in PNG (Fig. 4e; Tudhope et al. 1995;

355

Aycliffe et al. 2004). Sudden cooling of SST by 1 8C offshore PNG (from coral Sr/Ca) correlate with seasonal establishment of the NW monsoon and enhanced mixing of the water column (Aycliffe et al. 2004). There is a suggestion on the basis of isotopic studies that the western equatorial Pacific may have been less important in modulating inter-annual climatic variability between the 1920s and 1950s than subsequently (Tudhope et al. 1995). Due to reduced marine salinities, SE Asian waters typically show aragonite saturations of c. 0.5 –1 less than those for the Caribbean or Red Sea, translating to c. 10% reduction in calcification rates (Fig. 3e; Kleypas et al. 1999; Hoegh-Guldberg et al. 2007). Simulations of ocean chemistry related to postulated rising CO2 levels over the next century show that the aragonite saturation in Australasia will drop below those suitable for reef development (V-aragonite .3.3) this century and earlier than other areas such as the Caribbean (Kleypas et al. 1999; Hoegh-Guldberg et al. 2007).

Thousand to hundred thousand year scale changes Uplifted coral reef terraces provide the most common record of effects of environmental change on shallow marine carbonates on timescales of thousand to hundreds of thousands of years in SE Asia (Fig. 5). Spectacular ‘flights’ of uplifted terraces dated via U/Th, C14 or Electron spin resonance (ESR) techniques reveal significant recent local uplift (Fig. 5a–c & Table 1). On the Huon Peninsula, PNG, uplift rates of 1–3 m/ka over the last 300 000 years are associated with arc –continent collision (Bloom et al. 1974; Chappell 1974; Ota & Chappell 1996; Cutler et al. 2003). In Eastern Indonesia collision related uplift varies from 0.4 to 1.8 m/ka (Table 1; Chappell & Veeh 1978; de Smet et al. 1989; Sumosusastro et al. 1989; Pirazzoli et al. 1993; Hantoro et al. 1994; Bard et al. 1996). Arc uplift associated with oceanic subduction and sediment underplating is 1.6 + 0.4 m/ ka on New Britain (Riker-Coleman et al. 2006). In all areas prominent terrace levels mainly correlate with major eustatic sea level highstands associated with interglacial stages and substages (Fig. 5c, d; e.g. isotopic stages 1, 5a, 5c, 5e and 9; Chappell 1974; Sumosusastro et al. 1989; Pirazzoli et al. 1993; Hantoro et al. 1994; Bard et al. 1996). Patchy sub-terrace development with regressive configurations may be due to metre-scale discrete uplift events or possibly rapid falls in sea level (Yokoyama et al. 2001). Periodic iceberg discharges from partial breakup of northern Hemisphere icesheets (‘Heinrich’ events) or Antarctic icesheet surges that precede times of cold climates,

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Fig. 5. Examples of centennial to thousand year scale environmental changes from coral reef terrace deposits. (a) Geomorphic block diagram from the Luwuk area of East Sulawesi showing three groups of uplifted reefal terraces (from Sumosusastro et al. 1989). Vertical exaggeration is approximately 2. The low terrace offshore Luwuk town is attached by a recurved spit to the mainland, the middle series above Luwuk town consists of a moderately seaward slanting surface and an upper group is dated at c. 229 ka. (b) Photograph of uplifted coral reef terraces with a stepped appearance from the Tukang Besi Islands, SE Sulawesi (by M. Wilson). (c) Age-height plot of terraces from the Huon Peninsula, based on extrapolation of uplift rate estimated from 120 000 a old terrace, tentatively correlated with major transgressions (from Chappell & Polach 1976). (d) Stratigraphic section of the raised coral reef terraces of the Huon Peninsula, with age constraints from U– Th dating (from McCulloch et al. 1999). (e) Sea level heights derived from U-series dating of corals together with planktonic foraminifera d18O variations for the Sulu Sea ODP site 769, and temperature data from the Huon Peninsula for the period prior to, and including, the Last Interglacial (from McCulloch et al. 1999).

Table 1. Studies of uplifted coral reefs terraces in SE Asia, uplift rates, dating techniques used and inferred reasons for uplift are shown together with main references Uplift rate (& timeframe)

Dating technique 14

Huon Peninsula, PNG

0.5–3.0 m/ka (300 ka)

Th – U, C

SE New Britain

1.6+0.4 m/ka (9 ka and less)

230

Th

Alor Island

1.0–1.2 m/ka (500 ka)

230

Th/234U, C14, ESR

Atauro Island & adjacent N Timor

0.47 m/ka–0.5 m/ka (Atauro & Timor, dated to 120 ka extrapolated to 700 ka) and 0.03 m/ka near Dili, Timor 0.2–0.5 m/ka (600 ka extrapolated to 990 ka)

Th230 – Ur234

Sumba Island

Th/U, ESR

Luwuk, East Arm & SE Arm, Sulawesi

0.53–1.84 m/ka variable over time & local area (67 – 350 ka)

U/Th, C14

Banda Arc

Up to or . 0.5 m/ka

Palaeontology

Inferred reason for uplift Arc (Finisterre Volcanic Arc) – continent (New Guinea Highlands) collision Arc uplift associated with oceanic subduction & sediment underplating Uplift associated with Australian continental crust collision Uplift associated with Australian continental crust collision Uplift associated with Australian continental crust collision Tectonic uplift of coast facing Sula Platform, but also block faulting and regions of subsidence Collision processes in Banda Arc area

References Chappell (1974), Bloom et al. (1974), Chappell & Polach (1976), Cutler et al. (2003) & others in text Riker-Coleman et al. (2006) Hantoro et al. (1994) Chappell & Veeh (1978)

Pirazzoli et al. (1993), Bard et al. (1996) Sumosusastro et al. (1989), Fortuin et al. (1990)

CARBONATES AND ENVIRONMENTAL CHANGE

Area

de Smet et al. (1989)

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correlate with rapid sea level rises and have been correlated with a number of the Huon Peninsula terraces (Aharon et al. 1980; Yokoyama et al. 2001). From dating of corals in Aladdin’s Cave on the Huon Peninsula Esat et al. (1999) have suggested that ice sheet breakup may have paused during major deglaciation and that meltwater pulses were non uniform and episodic. Uplift rates along individual terraces may vary in both time and space and some are segmented by faulting, or dip, due to tectonic tilting (Chappell 1974; Chappell & Veeh 1978; Sumosusastro et al. 1989; Hantoro et al. 1994). Cave sediments and speleothems from the Melinau Limestone of Borneo have been used to evaluate base-level change over the last 2 Ma (Farrant et al. 1995). Uplift rates of 0.19 m/ka may be related to isostatic uplift associated with regional denudation and flexure associated with offshore sediment loading (Farrant et al. 1995). A series of modern and fossil corals from Sumatra and PNG have been used to infer oscillations in the extent of the Indo-Pacific Warm Pool since the mid Holocene through Sr/Ca derived sea surface temperatures (Abram et al. 2009). Mid Holocene cooling at the study sites was related to a more northerly position of the Intertropical Convergence Zone (ITCZ) and associated strengthening of the summer monsoon (Abram et al. 2009). Caliche horizons formed during dry climates and subaerial exposure of carbonate platforms in the South China Sea have been related to major Pleistocene sea level lowstands (i.e. glacial periods; Gong et al. 2005). The caliche may indicate reduced extent of the Western Pacific Warm Pool, a southerly shift of the ITCZ and reduced strength of the East Asian Monsoon during glacial periods (Gong et al. 2005). Oxygen isotope data from speleothems in Borneo have been used to infer lower rainfall and weakening western Pacific convection related to a southward shift of the ITCZ during deglaciation 18– 20 ka ago (Partin et al. 2007). Areas of significant subsidence may occur in close proximity to regions of contemporaneous rapid uplift (e.g. Luwuk and Huon Peninsula; Sumosusastro et al. 1989; Galewsky et al. 1996; Webster et al. 2004a, b). In the actively subsiding foreland basin of the Huon Gulf a series of backstepping platforms now at depths of 0.1– 2.5 km have been systematically drowned approaching the trench due to subsidence rates of c. 5.7 m/ka over the last 450 ka (Galewsky et al. 1996; Webster et al. 2004b). Submarine terraces are also tilted along strike, probably due to thrust loading (Webster et al. 2004a). Relative sea rise during early major deglaciation events or interglacials when combined with regional subsidence was .10–15 m/ka and was critical in terrace and platform drowning (Webster et al. 2004b). Tectonic

subsidence and basement substrate morphology influenced overall geometries and tilting over timescales of .100– 500 ka (Webster et al. 2004a). A change in the drowned reefs from high energy to low and moderate energy assemblages around 300–350 ka may be related to tectonic constriction in the Huon Gulf and/or perhaps changes in the Intertropical Convergence Zone (Webster et al. 2004b). In uplifted reef terraces older than 10 000 years from SE Asia, most corals are altered from their original aragonite mineralogy. Rare samples from the Huon Peninsula and Sumba have extended the record of unaltered corals back to 130 and 600 ka, respectively (Pirazzoli et al. 1993). Oxygen isotope results from giant clams linked to uplifted reef terrace data from the Huon Peninsula have been used to suggest SST of tropical oceans during interstadials of isotope stages 5 and 3 were similar to, or up to 3 8C cooler than, the present day (Aharon & Chappell 1986). More recent work on Sr/Ca ratios from unaltered corals from the penultimate deglaciation at 130 + 2 ka reveal that SST were significantly cooler, by c. 68 + 2 8C, than either the last interglacial or present day tropical temperatures (Fig. 5e; 298 + 1 8C; McCulloch et al. 1996, 1999). Sr/Ca corals records also reveal seasonal temperature fluctuations from 8.9 and 7.3 thousand years ago in the range of +1 8C with possible ENSO related fluctuations of +2 8C (McCulloch et al. 1996). ENSO fluctuations have now been shown to operate over the last 130 ka and during glacial time periods from d18O original coral mineralogy from the Huon Peninsula (Tudhope et al. 2001). However, the magnitude of ENSO fluctuations today appears strong compared with previous cool (glacial) and warm (interglacial) periods, with changes related to glacial dampening and orbital driven precession (Tudhope et al. 2001).

Millennial and Cenozoic-scale changes A range of shallow marine SE Asian carbonate deposits, known from outcrop and subsurface data, provide an extensive record spanning much of the Cenozoic (last 50 million years; Fulthorpe & Schlanger 1989; Wilson 2002, 2008). Individual systems variably show the strong influence of clastic and/or nutrient influx, tectonics, eustasy, biotic change and oceanographic factors acting over millennial scales (Fig. 6a, b). Over the Cenozoic timescale a significant change in biota from larger benthic foraminifera to coral-dominated systems occurred around the Oligo-Miocene boundary and had a major impact on carbonate platform development (Wilson & Rosen 1998; Wilson 2008). Longterm changes over the scale of the Cenozoic have been related to changing oceanography, plate

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tectonic configurations, climatic changes and perhaps global fluctuations in CO2 levels (Fig. 6c; Wilson 2008). Regional tectonics via plate tectonic movement, extensional basin formation and uplift was a major control on the location of carbonates during the Cenozoic in SE Asia (Wilson 2008; Wilson & Hall 2010). These processes controlled movement of shallow marine areas into the tropics and their disappearance through emergence or subsidence. Locally, the creation of faulted highs, volcanic edifices, microcontinental blocks, and basins trapping siliciclastics influenced the location of carbonate initiation (Wilson 2008). Individual carbonate platforms are influenced by tectonics through: (1) fault-margin collapse and reworking, (2) fault segmentation, (3) tilting of strata, subsidence, uplift and differential generation of accommodation space and (4) modification of internal sequence character and facies distribution (Fig. 6a; Wilson & Bosence 1996; Wilson 1999, 2000; Wilson et al. 2000; Bachtel et al. 2004; Wannier 2009; Wilson & Hall 2010). Eustasy, particularly during the Miocene, was also a major impact on changing accommodation space influencing sequence development, facies distribution and platform geometries (Epting 1980; Rudolph & Lehmann 1989; Vahrenkamp et al. 2004; Paterson et al. 2006). A subsurface study of Oligo-Miocene carbonates from the Philippines is amongst the first in SE Asia to relate high-frequency, metre-scale platformtop cycles to 4th and 5th order eustatic fluctuations (10–1000 ka) dated using strontium isotopic analysis (Fournier et al. 2004). In the Gulf of Papua it has been suggested that Oligo-Miocene patterns of carbonate stratal aggradation, progradation and backstepping imaged on seismic relate directly to global sea level fluctuations (Fig. 6b; Tcherepanov et al. 2008a, b). These have been correlated with a eustatic-driven global stratigraphic signature (Fig. 6b, c; Tcherepanov et al. 2008a, b). In areas of strong oceanographic or monsoonal driven currents elongation of carbonate buildups, progradation of sediment and windward-leeward facies differentiation are all common (Tyrrel et al. 1986; Carter & Hutabarat 1994; Gro¨tsch & Mercadier 1999; Wilson & Evans 2002; Carter et al. 2005). Although many of the Cenozoic SE Asian carbonate systems are located away from clastic influx, c. 70% formed as land-attached features (Wilson & Hall 2010). Many of these were affected by clastic or volcaniclastic influx, and that .80% of these developed around small-scale islands is probably a reflection of more limited or periodic influx compared with large-scale islands (Wilson & Hall 2010). In volcanic areas carbonate production is often hindered in the local vicinity of active volcanoes, but distal from these or during periods of

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relative quiescence volcanic edifices often form sites for prolific carbonate development (Fulthorpe & Schlanger 1989; Wilson 2000; Wilson & Lokier 2002; Satyana 2005). Studies of Miocene carbonates in fore-arc settings from Java and delta-front areas in Borneo show that a range of carbonate producers occur in regions of near constant or punctuated influx (Wilson & Lokier 2002; Wilson 2005; Lokier et al. 2009). Habitats characterized by frequent or heavy clastic influx favour organisms that can move around or shed sediment (Wilson & Lokier 2002; Wilson 2005). If light dependent organisms are present in these areas their presence is typically restricted to the upper few metres of the water column as a result of reduced water clarity (Wilson & Lokier 2002; Lokier et al. 2009). Nutrient runoff was also a factor controlling biotic change in these systems, with coralline algae and heterotrophic feeders becoming more common (Wilson & Lokier 2002; Wilson 2005). Nutrient runoff and upwelling causing plankton blooms and reduced water clarity is inferred to promote the regional development of low-light level perforate foraminifera assemblages on the deeper parts (.20 m) of many Cenozoic platforms in the region (Wilson & Vecsei 2005). The analysis of factors influencing temporal trends of Cenozoic carbonates in SE Asia given below is from Wilson (2008), and further explanation can be found therein (Fig. 6c). ‘The Early Miocene acme of coral-rich facies extent and abundance in SE Asia lags Oligocene coral development in the Caribbean and Mediterranean, despite local tectonics providing apparently suitable habitable areas (Wilson & Rosen 1998; Perrin 2002; Wilson 2008). Regional and global controls, including changing, oceanography, nutrient input and precipitation patterns are inferred to be the main cause of this lag in equatorial reefs. It is inferred that moderate (although falling) levels of CO2, Ca2þ and Ca/ Mg when combined with the reduced salinities in humid equatorial waters all contributed to reduce aragonite saturation (Stanley & Hardie 1998; Kleypas et al. 1999; Hallock 2005). This hindered reefal development compared with warm more arid regions during the Oligocene. By the Early Miocene, atmospheric CO2 levels had fallen to preindustrial levels (Fig. 6c; Zachos et al. 2001; Pagani et al. 2005). Although this was a relative arid phase globally, in SE Asia palynological evidence indicates the Early Miocene experienced everwet, but more stable and less seasonal conditions than periods before or after (Morley 2000; Morley et al. 2003). Tectonic convergence truncated deep throughflow of cool nutrient-rich currents from the Pacific to Indian Ocean around the beginning of the Miocene (Kuhnt et al. 2004), thereby directly, and perhaps indirectly (though less seasonal conditions)

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Fig. 6. Examples of millennial scale environmental changes inferred from carbonate deposits of Cenozoic age. (a) East–west cross section through the Eocene to Miocene Tonasa Formation from Sulawesi showing the influence of tectonic faulting, differential uplift and subsidence on the development of the carbonate platform and its facies (from Wilson et al. 2000). (b) Seismic section through the Gulf of Papua with interpretation of carbonate packages correlated with global eustatic sea level fluctuations, also shown in red on the eustasy column of Figure 6c. Interpretation is (1) late Oligocene–early Miocene aggradation, backstepping and partial drowning, (2) late Early Miocene–early Middle Miocene vertical growth or aggradation, (3) Middle Miocene downward shift of deposition, (4) late Middle Miocene progradation (systematic lateral shift of sedimentation) and (5) late Miocene and Early Pliocene re-flooding and

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reducing nutrients. It is inferred that aragonitic reefs were promoted during the Miocene where previously the waters had been more acidic, more mesotrophic, more turbid and less aragonite saturated. Extensive reefal development resulted in an order of magnitude expansion of shallow carbonate areas through buildup and pinnacle reef formation in the Early Miocene. Tectonics via basin formation, increased habitat partitioning and reducing distances to other coral-rich regions may also have contributed to enhanced reefal development (Wilson & Rosen 1998; Renema et al. 2008). Declining reefal importance at the end of the Early Miocene resulted from uplift of land areas, enhanced oceanic ventilation, through thermohaline circulation (Halfar & Mutti 2005) and narrowing of oceanic gateways as well as increased seasonal runoff, at least in SE Asia through initiation/intensification of the monsoons (Jia et al. 2003)’ (Wilson 2008).

Discussion Data sources, limitations and further research Tomascik et al.’s (1997) study is a major contribution to our understanding of modern SE Asian coral reefs and seas. This work expands on early studies by those such as Semper (1881), Umbgrove (1946, 1947) and Kuenen (1933) which were amongst the first to evaluate the biology, morphology and controls on development of the regions reefs. Recent decades have also seen the setting up of projects to evaluate the unique oceanography, complex environmental settings, marine biodiversity and impacts of climate change in SE Asia (Wilkinson 2000; Spalding et al. 2001; Gordon 2005; Pen˜aflor et al. 2009). Whilst Tomascik et al. (1997) and others studies do ‘introduce . . . the fascinating marine environment of the Indonesian Seas’ they also point out ‘serious gaps in our knowledge . . . and a lack of basic research’. There is considerable potential to better evaluate the biota, communities, regional conditions and environments of the modern seas of SE Asia. To quote Tomascik et al. (1997), ‘a solid knowledge and understanding of the marine and coastal environments must form the foundation . . . for the region’s most precious asset, its people, to manage and conserve its secondmost valuable asset, the seas’.

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The handful of sclerochronology studies on modern and Quaternary SE Asian corals highlight the value of recently developed techniques in determining environmental change. Factors such as temperatures, runoff, upwelling are now being related to local, regional or global change, both human-induced and natural (e.g. mining, ENSO, glacials and interglacials; McCulloch et al. 1999; Tudhope et al. 2001; Fallon et al. 2002). Currently, most of these studies are restricted to more easily accessible areas (Java) and relatively well-studied areas such as the Huon region of PNG (Scoffin et al. 1989; Tudhope et al. 1995; Gagan et al. 1998). Due to the region’s high rainfall and abundant vegetation (hence humic acids) conditions for preservation of original aragonite coral mineralogy are not good and many corals show signs of recrystallization or complete alteration within a few thousand years of formation (Riker-Coleman et al. 2006). Recent studies have highlighted the potential for preservation of original aragonite to hundred(s) of thousand(s) of years in strongly seasonal parts of the region or highly local areas, such as caves (Pirazzoli et al. 1993; McCulloch et al. 1999; Tudhope et al. 2001). As yet no studies of growth banding on organisms other than corals or those predating the Quaternary have been undertaken in SE Asia (cf. Purton & Brasier 1997, 1999). There is the opportunity to continue to develop combined growth banding and geochemical studies to evaluate short-term environmental change in the region over Quaternary and deeper timescales. The range of studies on the uplifted Quaternary coral reef terraces of the Huon Peninsula reveal the potential for extracting environmental change data on factors including tectonics, eustasy, temperatures, ENSO, glacials/interglacials and ice sheet dynamics (Chappell 1974; McCulloch et al. 1996, 1999; Esat et al. 1999; Tudhope et al. 2001). The now submerged (up to 2.5 km water depth) time-equivalent platform and reef terraces offshore in the Huon Gulf (Webster et al. 2004a, b) are the only such systems in SE Asia to have undergone detailed subsea imaging, dating, facies and biological mapping. Although spectacular, these Papuan examples are by no means unique in the region (Pirazzoli et al. 1993; Hantoro et al. 1994). There is the potential to undertake considerable further studies of uplifted and submerged coral reef terraces to better understand regional tectonics,

Fig. 6. (Continued) aggradation (from Tcherepanov et al. 2008b). (c) Carbonate biofacies, numbers of platforms/ buildups in SE Asia plotted against regional and global events during the Cenozoic (from Wilson 2008; with equatorial climate after Morley 2000; Morley et al. 2003, global framework reefs after Perrin 2002, oceanic Ca & Mg from Stanley & Hardie 1998, global climatic events, d18O & Atmospheric pCO2 from Zachos et al. 2001; eustasy from Haq et al. 1987, and additional atmospheric pCO2 data from Pagani et al. 2005).

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relative sea level variations and the impact of regional and global environmental change. A major review of reef development since the last glacial maximum (23 ka) across the greater Indo-Pacific region (far Western India to far Eastern Pacific Oceans) is based on drill core data (684 cores) through shallow water carbonates (Montaggioni 2005). Despite the comprehensive review of available data, no cores were available from the central Indo-Pacific biodiversity hotspot (Indonesia and the Philippines). Cores are predominantly from oceanic ‘blue-water’ reefs that are unlikely to be representative of many SE Asian carbonate systems (Tomascik et al. 1997). Montaggioni (2005) defined three main Indo-Pacific reef building periods: 17.5 –14.7, 13.8–11.5 and 10 ka to present. Nutrient levels, hydrodynamic energy and to a lesser extent substrate availability and ocean circulation were major influences on reef accretion at local and regional scales (Montaggioni 2005). There is a need for a programme of shallow water carbonate drilling in SE Asia to better characterize the variability of Indo-Pacific reefs, their history and controlling influences. The growing number of studies on SE Asian Cenozoic carbonates are beginning to reveal factors controlling carbonate development over timescales of millions of years (Fulthorpe & Schlanger 1989; Wilson 2002, 2008; Renema et al. 2008; Wilson & Hall 2010). However, just 10% of the regions 299 carbonate formations documented by Wilson (2002) have detailed biofacies studies. There is little recent systematic palaeontology, (with the exception of the foraminifera; Wilson & Rosen 1998; Renema et al. 2008 and early Dutch studies on corals and molluscs [e.g. von Fritsch 1878; Martin 1879; Gerth 1930, 1931; Umbgrove 1939, 1946], geochemistry has been undertaken on very few deposits, and isotopic dating of sections is limited. There is a need for further detailed sedimentological, biotic, geochemical and rigorous dating studies of outcrop and subsurface Cenozoic carbonates. In short, further baseline data together with targeted studies are required to continue to test hypotheses of environmental change and influences on equatorial carbonates on all timescales and during all time periods.

Environmental change: comparisons from terrestrial, shallow and deep marine systems To investigate environmental change in the equatorial tropics a holistic approach is required integrating data from terrestrial, coastal and deep marine systems as well as the shallow marine systems discussed here. Notwithstanding the need

for further studies in all fields (see above, Oldfield 1998) how do the emerging trends from the shallow marine carbonates compare with those from other realms? Also what can be learnt about the role of the equatorial tropics in global climate change? The terrestrial and deep marine records, similar to shallow water carbonates, show the major influence of local environmental perturbations and climate change on annual to thousand year scales (e.g. fire, volcanism, deforestation, precipitation, ENSO, glacials–interglacials). Palynological research of marine cores is now providing a connection between marine and terrestrial environments (Dam et al. 2001). There is recognition of a correlation between droughts, biomass burning and extreme ENSO events in the equatorial tropics (Goldammer 1999). Cores extending back .20 ka reveal charcoal records showing that in some areas tropical forests may never have been continuously fire-free for long (Hope 2001). Changes in forest structure and replacement by grassland vegetation together with increases in disturbance pollen indicators and charcoal levels have been related to increased human impacts on the landscape particularly during Holocene, but also the late Pleistocene (X. Wang et al. 1999; van der Kaars et al. 2000; Anshari et al. 2001), and possibly development of ENSO climatic variability (Anshari et al. 2001; Haberle et al. 2001). Many palynological and/or deep marine records show evidence for reduced precipitation, reduced temperatures (by up to or greater than 4 8C), grassland expansion and altitudinal lowering of montane floras during glacial periods (Flenley 1979; van der Kaars 1991; van der Kaars & Dam 1995; Dam et al. 2001; Hope 2001; Rosenthal et al. 2003). This may not have been the case for what are now the most humid regions today, and increased strength of the winter monsoon may be a possible cause during recent glacial maxima (Sun et al. 2000; Visser et al. 2004). Over longer time periods everwet conditions have been inferred for SE Asia during the Middle Eocene and Early Miocene (Morley 2000). Although everwet conditions persisted in areas such as Borneo for extended periods, more seasonal conditions were experienced in Java and Malaysia during the late Oligocene and from the middle Miocene (Morley 2000; Lelono & Morley 2011), both periods of ice cap expansion (Zachos et al. 2001). Deep marine cores are revealing high resolution records of changing current, monsoonal as well as ENSO and glacial–interglacial climatic fluctuations in SE Asia over Cenozoic and Quaternary timescales. Geochemical data from the intermittently dysoxic Kau Bay in Halmahera is used to suggest diminished ENSO amplitude or frequency

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during the Medieval Warm Period (1000–750 a BP) and decreasing El Nin˜o activity during, and after, the Little Ice Age (Langton et al. 2008). Oceanic productivity increases are associated with reduced precipitation during glacial periods from a 460 ka record in the Timor Sea (Kawamura et al. 2006). This data supports the hypothesis that the IndoPacific Warm Pool (IPWP) experienced ENSOlike conditions during glacial periods (Stott et al. 2002), with low summer radiation and a weak austral summer monsoon (Kawamura et al. 2006). Regional variations are inferred for the Western Pacific marginal seas from sedimentation rates and geochemistry (Wang 1999). The South China Sea experienced emergence and cooling during glacials and has been linked to aridity in China. Conversely strengthened winter monsoons and intensified seasonality of the marginal seas could bring moisture to the islands of SE Asia during glacial periods (Wang 1999; also seen in the palynological record above). Many of the region’s other shallow shelves, such as the Sunda and Sahul Shelves experienced emergence during glacials (Umbgrove 1938; Pelejero et al. 1999; Wilson & Moss 1999). Emergence and drowning of SE Asian shelves impacted monsoonal transport of moisture, hydrological and geochemical cycles of the western Pacific, impacting global climate with La Nina conditions inferred for glacials (Pelejero et al. 1999; L. Wang et al. 1999). Emergence and drowning also dramatically influenced biogeographical connectivity between areas resulting in a region of dynamic geographical, ecological and biogeographical mosaics in which shallow marine carbonate platform and coastal terrestrial development were influenced (B. R. Rosen, pers. comm. 2010; Potts 1983; Rosen 1984). There is debate on the timing of initiation and intensification of the monsoons using a range of proxy data including lithologies, grain textures and geochemistry of marine sediments (Jia et al. 2003; Wang et al. 2005). There is the possibility of inception of a monsoon around 20 Ma, initial strengthening in the early to middle Miocene and strengthening of the winter monsoon around 7 Ma (Clift et al. 2002; Jia et al. 2003). In the Maldives it is proposed that onset and intensification of the monsoon promoted vigorous bottom currents, triggered nutrient upwelling and caused a switch from aggradation to backstepping sequences and ultimately drowning of carbonate platforms (the latter when associated with periods of eustatic sea level rise; Betzler et al. 2009). Any potential link between monsoons, platform development and drowning in SE Asia requires further testing, but may be important where currents are intensified and/or nutrient influx occurs (Wilson 2008). There is growing consensus that not only are the equatorial tropics involved in global environmental

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change, but that this region may be far more important in controlling climate variability than previously thought (Kerr 2001). Visser et al. (2003) showed that 3.5–4 8C rises in SST in the Makassar Straits during glacial to interglacial transitions are synchronous with atmospheric CO2 rises and Antarctic warming, but pre-date Northern Hemisphere icecap melting. The inference was that the tropical Pacific region was driving glacial–interglacial cycles perhaps through ENSO regulation of transfer of water vapour and heat to the poles (Visser et al. 2003). The theory is that as the IPWP warms, the region supplies additional water vapour and CO2 (reduced solubility in warm waters) to the atmosphere, resulting in atmospheric transfer of heat poleward and enhanced greenhouse effects thereby warming the planet (Kerr 2001; Dunbar 2003; Visser et al. 2003). In addition, changing configurations of land and sea in SE Asia associated with tectonics and glacial cyclicity influenced ocean – atmosphere circulation patterns with impacts on regional and global climates (Dam et al. 2001). The Indonesian Throughflow Current is a major influence on global thermohaline circulation which in turn impacts global climate. Restricting the deep Throughflow due to tectonic convergence around 25 Ma reduced the input of cool deep waters to SE Asia (Kuhnt et al. 2004). The resultant increase in sea surface temperatures is inferred to have promoted evaporation and the development of humid, but stable conditions in the central archipelago (Morley et al. 2003). Change in position of the Throughflow passages, again due to tectonic change in the Pliocene are inferred to have caused reduced atmospheric heat transfer from the tropics to high latitudes (Kuhnt et al. 2004). An unrestricted Throughflow (from benthic foraminifera records) between 1.6 to 0.8 Ma is inferred to influence initiation of the Leeuwin Current along Western Australia and the transfer of heat to southerly latitudes (Gallagher et al. 2008). A range of studies show the complex links between the Indonesian Throughflow, ENSO fluctuations and the Asian Monsoon (Wang et al. 2005; Kawamura et al. 2006). During the boreal winter monsoon there is cooling of the tropical Indian Ocean since the warm surface flow of Indonesian Throughflow is hindered by monsoonal winds. This enhances the contrast in sea surface temperature between the western Pacific and Indian Ocean with potential feedback for ENSO, the Indian Ocean Dipole and weakening monsoonal phenomena (Gordon et al. 2003; Xu et al. 2008).

Conclusions This study highlights the use of carbonate producers and deposits in evaluating environmental and

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climatic changes on all scales through the last 50 million years in SE Asia. Major drivers of local and regional changes at a variety of temporal scales are inferred to include oceanography, nutrients, precipitation patterns and terrestrial runoff. At a higher level these are driven by tectonic changes in SE Asian ocean gateway configurations, subsidence versus uplift, eustasy, and global/ regional climatic and ocean–atmosphere changes (e.g. ENSO, monsoons and CO2 fluctuations). Despite conditions that commonly differ significantly from the more traditional view of ‘bluewater’ oceanic reefs, SE Asia has the highest modern shallow marine biodiversity and the most volumetrically extensive and complete record of equatorial carbonates spanning much of the Cenozoic. In general the western SE Asian seas are strongly influenced by terrestrial runoff, the Indonesian Throughflow current and monsoonal effects, whereas the east experiences major nutrient upwelling and ENSO impacts. These factors all influence carbonate producers and the edifices they build, as evidenced from geochemical signatures, facies and carbonate platform studies of modern and Quaternary deposits in SE Asia. Spectacular coral reef terraces reveal significant localized tectonic uplift and coeval subsidence during the Quaternary and Pleistocene with rates (up to 3–6 m/ka) similar to some of the fastest known. Quantification is possible of amounts, or timing, of factors such as interglacial to glacial temperature changes (up to 3–6 8C), ENSO fluctuations (+2 8C extending back at least 130 ka), meltwater pulses associated with ice sheet breakup, and movement of the ITCZ. Over millennial timescales siliciclastics, nutrients, tectonics, eustasy and oceanography all influence the location, evolution and variability of individual carbonate systems. Over the Cenozoic major changes in oceanography, plate tectonics, climate change and perhaps fluctuating atmospheric CO2 influenced significant changes in carbonate producers and the types of platforms that were constructed. Comparison with parallel studies from terrestrial and deep marine deposits reveals many similar trends to those emerging from the carbonate record, but the former are currently yielding more information about variations in precipitation, currents and the monsoons. There is considerable scope for further study in all fields. It is anticipated that study of SE Asian carbonates will continue to aid evaluation of the relative response of tropical marine systems to global or regional change versus their potential to be dominant drivers of climatic change. I thank Robert Hall and the SE Asia Research Group for their continued support, aiding my research of the fascinating carbonate systems of the region, and enabling me to present this study at the SAGE 2009 conference. Funding

through an Internal Curtin University Research Grant to undertake this study is acknowledged. Bob Park and Brian Rosen are thanked for their constructive reviews.

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Index Page numbers in italic refer to Figures. Page numbers in bold refer to Tables. Agulhas Leakage 268 Ailaoshan Suture 9, 15, 17 Aileu Metamorphic Complex 244, 246, 251– 253, 257 Alino Group 41 Alor Island 248, 251 ammonoids, Olenekian 30 amphibolite, Aileu 257 Antarctic Intermediate Water 270 Ar dating, Benteng Intrusives 189, 190–192, 193 Arafura Platform, comparison with East Java Terrane 65, 66, 71, 72 aragonite saturation 351 arc, intra-ocean, Early Cretaceous 41 arc terranes, Carboniferous –Permian 9– 10 Argo Abyssal Plain 29, 84, 112, 115, 127, 129 Argo Block 82 Argoland 12, 14, 29– 30, 77–78 Atauro Island 248, 251 Attaka Well B, palynology 320, 323– 327 Australasian Mediterranean Water 293, 295 Australia, basement structure 85 Australia– SE Asia collision 1, 75, 84–91 Banda volcanic fore-arc 241, 242, 243, 251, 252, 256–257 dating 243, 245–250, 251 Gondwana margin 19, 21 Java margin 111–134, 139 Sulawesi 179– 180 Australian Plate 7, 8, 78 origin of Borneo and Sarawak 77 Bahamian limestones 254, 255 Balangbaru Formation 71 Balantak Fault 198, 218, 220– 221, 222 Balantak Ophiolite 183, 194 Bali, seismic tomography, south-dipping thrust 153– 154 Banda Arc Outer, non-volcanic 243, 245, 262 volcanic fore-arc 237, 238, 239, 245, 251 age 89–90, 241 collision with Australian margin 242, 243, 245, 258–259 Banda Block 14, 29, 79, 80 see also Inner Banda Block Banda Embayment 72, 79, 80, 84, 86, 99 reconstruction 86, 87– 88, 89–91 Banda Ridges 89, 90 Banda Sea, ITF 272, 273, 277 Banda Terrane 90, 243, 244, 246 basement, origin and composition 250–253 cover rock sequences 253– 254 Mt Cablac Range 255 Triassic sequences 254 Banda Trench subduction 237, 241, 243, 251, 258 see also Tectonic Collision Zone Banggai-Sula microcontinent 10, 158, 172, 179– 180, 194

basin floor, stratigraphy 207, 210 basin margin stratigraphy 207, 208–209 structure 215 basin slope, structure 210, 211, 212 collision with Molucca Sea Plate 203– 222, 204 NE area, structure 212 –215, 216 stratigraphy 205 –210 thrust faults 217– 218, 219 western area, structure 210– 212, 213 Bantimala Complex 158, 169 Baoshan Block 10–11 Barito Basin 38, 92 lithostratigraphy 43 Barrique Volcanic Formation 254 basins, sedimentary Neogene, Borneo 92 subsidence and uplift 100 Gorontalo Bay 95–96, 98, 99, 196–198 Sundaland 38 Batu Putih Limestone 245, 247 diachronism 249 Batui Group 187 Batui Thrust 97, 204, 218, 222 Bawean microcontinental block 10, 14 Bayah Formation 45 Benteng Intrusives Ar dating 189, 190– 192, 193 stratigraphy 181, 182, 187, 189–193 volcanism 198 Bentong-Raub Suture 9, 16, 17, 84 Billiton Depression 77 Billiton Lineament 82 biota Gondwanan 9, 11, 22 Tethyan/Cathaysian Province 9, 22 see also Cathaysian Province, fauna Visean 20, 22 see also fauna; flora Bird’s Head 72 Bobonaro Me´lange 245– 248, 254, 256–257 Bobong Formation 207 Bongka Formation Pliocene uplift 194, 196 stratigraphy 181, 182, 185 Borneo Australia– SE Asia collision 78, 81, 92– 93 Neogene, climate change 319– 331 northern, subduction and deformation 92– 93, 100 –101 origin 76– 78, 81 see also SW Borneo Block Boyan Suture 9, 19 Brunei, Neogene 92–93 Buol Beds, carbonate platform 194 Buru-Seram Block 10 Buton-Tukang Besi Block 10, 158, 172, 180 Buya Formation 207, 210

374

INDEX

Cablac Limestone 248, 254, 255 Cambrian –Ordovician, Gondwanan blocks, evolution and palaeogeography 19, 21 Canning Basin 71, 72, 82, 85 carbon isotopes, Timor Sea outflow 287, 293 carbonates 347–364 Banggai-Sula basin 207, 210 clastic and volcaniclastic influx 359 as indicator of environmental change 350, 351, 352, 353–364 annual to centennial-scale 353–355 millennial to Cenozoic-scale 358– 361 thousand to hundred thousand year scale 355–358 Miocene, Sulawesi 194 Sumba 234–235, 357 tectonic control 359 Carboniferous, Gondwanan blocks, evolution and palaeogeography 20, 23, 24 Carboniferous–Permian, arc terranes 9 –10 Casuarina 336, 337, 338, 339, 341, 342, 344 Cathaysialand 9, 23 Cathaysian province, fauna 9, 12, 14, 23, 77, 81 Celebes Molasse 95, 99, 180, 185, 197 Ceno-Tethys ocean 19, 20, 30 Cenozoic carbonates 358– 361 deformation and plate tectonics 91– 92, 97–102 East Java Terrane 70 Central American Seaway 310–311, 316 Central Sumatra Basin 38 lithostratigraphy 43 Changning-Menglian Suture 9, 10, 15, 16, 23 Chanthaburi arc terrane 9, 10 Chanthaburi Suture 9, 16 chert, radiolarian 16, 17, 19 Karangsambung 41, 42 Chiang Mai Suture see Inthanon Suture Christmas Island Seamount Province 112, 114 –115, 121 Chuiella 19, 23 Cibulakan Formation, palynology 343– 344 Ciletuh-Jampang Block 119 Cimandiri fault zone 117 Cimmerian continent 11, 29 circulation, ocean global heat transport 269 mass transport 268– 269 salt transport 269 thermohaline 267– 268, 269 wind-driven 269– 270 regional and climate 270–272 El Nin˜o-Southern Oscillation 271 –272 climate everwet 319, 320, 331, 342– 343, 362 and regional ocean circulation 270– 272 seasonal 338, 340, 362 superwet 336– 338, 343 climate change Neogene, palynology 321–331 Quaternary, palynology 319– 321 CO2 levels, Miocene 308– 311, 312, 316 Comia flora 9, 26 Community Climate System Model 306, 308

conodonts, Late Palaeozoic 20, 21, 22, 23, 24, 26 ‘conveyor belt’, ocean circulation 267– 268, 269 coral as indicator of environmental change 347, 352 annual to centennial scale 353–355, 361 uplifted terraces, thousand to hundred thousand year scale 355 –358 Cretaceous East Java Terrane 59, 60, 61, 70 evolution and palaeogeography 27–30 Late, uplift, Sundaland 37, 44, 46– 48 subduction and collision 40–43 Sundaland assembly 76–78, 79, 80, 81– 82, 84 Cretaceous– Cenozoic, South China-Indochina margin blocks 14 Cretaceous– Paleocene, unconformity 37, 38, 44– 46, 48 Crocker Ranges 92 crust continental Australian margin 237, 238, 241, 251 East Java Terrane 58, 71, 81 east Sundaland 78, 81–82 exhumation 39 Savu Basin 234, 237, 238 thickening 39 Gorontalo Bay 177, 194, 198 thickening, Java margin 123 crustal flow 100 –101 Da Nang-Zeijiang Suture 18 Dacrydium 336, 337, 338, 339, 341, 342, 344 Dangerous Grounds terrane 14, 41, 84, 93, 94 flora and fauna 77 Dartollu Limestone 248, 254 decollements 256– 257 deformation Cenozoic 91– 93, 97–102 Savu Basin 231–233, 236–237, 239 Sundaland, alternative model 99–101 Timor 243, 247 Devonian, Gondwanan blocks 8– 9 evolution and palaeogeography 19 Dewakan sill 275, 284 diachronism, Timor 249–250 Dian-Qiong Suture 15, 18 Dicynodon 25, 26 dinoflagellates, Oligocene 338, 341, 342 Dondo Suite 159, 161, 165 ductile flow, central Sulawesi, metamorphic rocks 157 –173 dynamic topographic low (DTL) 40, 44, 46–48 earthquakes 1977 Sumba 131, 132 East Indonesia 96, 195 Java margin 115, 131, 133– 134 Sunda arc, seismic tomography 140–154 East Java microcontinental block 10 East Java Sea JavaSPAN seismic survey 55 oil and gas potential 53–54, 55 Oligocene palynology 333– 344 superwet elements 336–338

INDEX East Java Terrane 14, 53, 54 basement age 71–72 JavaSPAN seismic survey 53–72 departure and arrival unconformities 68–70 pre-Cenozoic sedimentary keels 61–65 stratigraphy 65, 67, 70–71 well penetrations 67–68 possible origins 71–72 pre-Eocene history 70 previous work continental crust 58, 71, 81 stratigraphy 58– 59, 61 East Java– West Sulawesi Block 53, 54, 55 collision 43, 82 comparison with Western Australia 71–72, 82 see also Argoland East Malaya Block 8, 9, 17, 84 East Sulawesi, subsidence and uplift 94– 96, 98–99 East Sulawesi Block 10, 86 East Sulawesi ophiolite 158, 171–172, 177, 179 eclogite, Karangsambung 41, 42 Ekman transport 274–275 El Nin˜o-Southern Oscillation 362– 363 and ITF 271–272, 275, 276 Miocene 305–317 changes in variability 313– 314 modelling 306 –308 environment, change role of carbonates 351, 352, 353– 364 comparison with records 362– 363 see also carbonates, as indicator of environmental change Eocene East Java Terrane 59, 60 Middle, evolution and palaeogeography 28, 30 sediments, SE Asia Regional Unconformity 45 Equatorial Counter Current 271, 272 Equatorial Undercurrent 270, 313, 316 Eurasian Plate 7, 8 everwet climate East Java Sea 342– 343 Makassar Strait 319, 320, 331 exhumation, 39, 44, 47 Fatu Laob Member 247 faults East Indonesia 97, 99 normal, Sumba Ridge 236 strike-slip, Banggai-Sula area 213–214, 216, 217, 220, 222 thrust Banggai-Sula area 217–218, 219, 222 Savu Basin 231– 233, 236– 237, 239 fauna Cambrian– Ordovician 19, 21, 22 Cathaysian Province 9, 12, 14, 23, 77, 81 Devonian 19, 22 Palaeozoic – Mesozoic 22 Visean 20, 22 Fitzroy Movement 65, 70 flora Lower Permian 25 Palaeozoic – Mesozoic 22 Visean 20, 22

375

Flores back-arc thrust 153 subduction slab ‘necking’ 141, 148, 152 volcanism 249, 250 Flores Sea 272 –273, 273 fold and thrust belts 39, 101 Banggai-Sula area 211, 213, 214 foraminifera Lonsio Formation 187, 188 Neogene Kutei Basin 323–327 Tarakan Basin 327, 328 Oligocene, East Java Sea 334– 335, 341 Pacific Ocean 316, 317 Peladan Formation 185, 186 Timor 244, 245, 246, 247– 248, 255 Timor Sea outflow 287– 289, 291, 293, 294 gateways see ocean gateways Gigantopteris 20 glaciation, Gondwanan 20, 26 Glossopteris 20 gneiss, Palu-Koro fault zone 164, 166, 167 Gondwana Carboniferous– Permian glaciation 20, 26 continental blocks 19 Devonian 8 –9 Jurassic 11– 14 Permian 9, 10–11 palaeogeography 19, 21, 23 terranes, Triassic– Jurassic 9, 77 Gorontalo Bay 177–199 basement rocks 193–194 basin subsidence 196–198 Miocene, carbonate platform 194 Pliocene alluvial fan 197 post-Pliocene tectonics 198 uplift and subsidence 95–96, 98, 99, 194, 196, 197–198 Greyhound Strait fault 97, 221 Gulf of Papua, carbonates 349, 360 Gumbassa Metamorphic Complex 159, 160–161 Halmahera Eddy 270 Halmahera Thrust 205 Haruyan Formation 41 Haulasi Formation 254 heat flow, Sundaland 82– 84 Heinrich Events, influence on ITF 298 human activity, impact on coral 354–355, 362 Huon Peninsula, coral terraces 348, 355–358, 361 Incertus Island Arc 30 Indian Ocean dipole 272, 279 equatorial, monsoon 271, 272 Miocene 308– 313, 316 pressure driven flow 273 –274 Indian Plate 7, 8 Indochina Block 8, 9, 17, 84 flora and fauna 22 Indochina–Pangaea landbridge, Permian 25, 26 Indonesian Passage 310–311, 313, 316 Indonesian Seas 272– 279 bathymetry 272, 273

376

INDEX

Indonesian Throughflow (ITF) 1, 267–279, 283– 284, 351 El Nin˜o-Southern Oscillation 271–272, 275, 276 freshwater flux 269 Indonesian Seas 272– 279 Karimata Strait 275, 276 Lifamatola Passage 275, 276, 277–278, 284 local forcing 273–275 Ekman transport 274–275 pressure-driven flow 273– 274 wind-driven flow 274 –275 Lombok Strait 275, 276, 277, 278 Makassar Strait 275 –276, 277, 284 ocean circulation 268, 272 Ombai Strait 275, 277, 278 outflow variability, last 140 ka 283–300 and SCST 276, 279 Timor Passage 275, 277, 278–279 Timor Sea outflow 285 –300, 286 Mg/Ca palaeothermometry 287, 288 salinity reconstructions 288, 289 stable isotopes 287–288 wind-driven circulation 269–270 Indoralian province, faunas 11, 23 Inner Banda Arc 89 Inner Banda Block 82 Intertropical Convergence Zone 295– 298 Inthanon Suture 9, 10, 15, 16, 18 intrusives, Benteng 187, 189– 193 island rule, wind-driven circulation 269–270 isotopes, stable, Timor Sea 287– 288, 289, 293, 294 Izu-Bonin arc, subduction 152 Java, subduction complexes 41 Java margin 111–134, 112 bathymetry 114, 121, 126, 130 earthquakes 115, 131, 133–134 fore-arc basin 117, 119, 122 geophysical data 113, 115– 116, 118, 119 landslides 129, 132 Lesser Sunda Islands, subduction–collision transition 127, 129– 131, 132 offshore central-eastern Java fore-arc erosion 121, 123, 132 seamount subduction 117, 119, 121, 123, 126, 127, 132 western Java fore-arc sediment accretion 116 –119, 132 Neogene prism 117, 119, 120, 124, 125 Palaeogene prism 117, 119, 120, 122, 124, 125 fore-arc sediment erosion 121, 123, 129 Java Trench 116–117, 232 subduction 86, 89, 111, 233–234 JavaSPAN seismic survey East Java Terrane 53– 72, 55 data and methodology 54–55, 57– 58 departure and arrival unconformities 68–70 interpretation paradigm 57–58 Meratus Suture 56 pre-Cenozoic sedimentary keels 61– 65 stratigraphy 65– 67 well penetrations 67–68 Jinghong Suture 9, 15, 17 Jinshajiang Suture 9, 17

Jurassic East Java Terrane 70 Gondwanan blocks 11–14 evolution and palaeogeography 27– 30 Kalimantan, correlation with Australia 81 Kambayau Granite 44 Kangean Island, JavaSPAN survey 62, 63 Kangean West-2 well 61, 63, 67, 68 Karangsambung, subduction complexes 41, 42 Karimata Strait 267, 273 ITF transport 275, 276 sill 293, 294 Karossa Metamorphic Complex 160 Kelabit-Longbawen Block 10, 14 ‘Kerapah’ type peat swamp, palynology 338, 341, 342 Khorat Group, red beds 44, 45, 46, 48 Kolbano region, micropalaeontological dating 245, 246, 247 Kujung Formation 333, 334 Kutei Basin 38 climate change, Neogene palynology 320, 323– 327, 330 lithostratigraphy 43 La Nin˜a events, and ITF 272 Labani Channel 275 Lake Sentarum, Quaternary palynology 319, 320 Lamusa Formation, stratigraphy 181, 182, 183 Lancangjiang Suture 9 landbridge, Indochina-Pangaea, Permian 25 landslides, Java margin 129, 130, 132 Langkowala Formation 180 Latimojong Formation 160, 161, 169 Leeuwin Current 271, 272, 283 Lesser Sunda Islands, subduction–collision transition 127, 129 –131 Lhasa Block 9, 10, 11 Mesozoic evolution 27 Lifamatola Passage 267, 272, 273 ITF transport 275, 277– 278, 284 limestone reefal, Luwuk Formation 193 Timor 248, 254–255 Lincang terrane 9, 10 lineaments Australia– SE Asia collision 83 Banggai-Sula area 210– 212, 213 Paternoster-Lupar 92 lithosphere, deformation 40, 82– 84, 99 Lok-Ulo Suture 19 Lolotoi Metamorphic Complex 243, 250 –251, 255, 257 Lombok Basin 127, 129, 130, 132 Lombok Strait 267, 272, 273 ITF transport 275, 277, 278 Lonsio Formation stratigraphy 181, 182, 184, 185, 187 biostratigraphy 187, 188 volcanism 198 Lonsuit Turbidites 187 Luconia Block 10, 14, 84 Luwuk, coral terraces 356, 357, 358 Luwuk Formation, stratigraphy 181, 182, 193 Luzon Strait 276

INDEX Macclesfield Bank Block 10, 14 magmatism Neogene, Western Sulawesi 172–173 subduction-related 41 Mahakam Delta, Quaternary palynology 320–321 Makassar Strait 94, 95, 171, 267, 272, 273 climate change Late Quaternary 319– 321, 322 Neogene, palynology 321–331 geostrophic conditions 274 ITF transport 275– 276, 277, 284 sill 275, 284, 298 Malawa Formation 71 Malay Basin 38 lithostratigraphy 43 Neogene, palaeoclimate 329, 330, 331 Malik Formation 183 Malino Metamorphic Complex 70, 158, 160, 171, 172 Manamas volcanic rocks 254–255 Mangkalihat microcontinental block 10, 14 mantle shallow, Java margin 131 Sunda arc, seismic tomography 139, 141, 142 –151 viscosity and uplift 40 Maubisse Formation 257 Mawgi Suture 9 Medial Sumatra Tectonic Zone 12, 18–19 ‘Mega-Lhasa’ Block 11 Meratus Mountains, subduction 42 Meratus Suture 9, 19, 53, 81 assemblage 53, 55 JavaSPAN seismic survey 56, 69 Meso-Tethys ocean 20 sutures 19 Mesozoic continental blocks 11– 14 evolution and palaeogeography 27, 30 SW Borneo 76–77 Mestognathus 20, 24 metamorphism west-central Sulawesi 159–173 non-coaxial strain 161– 168 causes and timing 169–173 origin of basement rocks 169 Metan Formation 254 Mg/Ca palaeothermometry 287, 288, 294 Mindanao Eddy 270 Miocene climate change, Makassar Strait, palynology 326– 327, 328–329, 331 East Java Terrane 59, 60, 86 El Nin˜o-Southern Oscillation 305– 317 changes in variability 313– 314 modelling 306 –308 Pacific and Indian Oceans, climate 308–313 Sulawesi 179– 180 carbonates 194 Mogok Metamorphic Belt 12, 19 Molucca Sea Plate 96, 98, 194 collision with Banggai-Sula Block 203–222 Molucca Sea wedge 96–97 Molucca-Sorong Fault 205 Monodiexodina 26

377

monsoons equatorial Indian Ocean 272 influence on ITF 275, 295 –298 Montognathus 20, 24 Mount Victoria Land Block see West Burma Block Mount Victoria metamorphic rocks 12 Mt Cablac Range, stratigraphy and tectonics 255 Mutis Metamorphic Complex 243, 250–251, 257 Nam Con Son Basin 38 lithostratigraphy 43 Nan-Uttaradit Suture 9, 10, 15, 17 nannofossils Neogene Kutei Basin 323–327 Tarakan Basin 327, 328 Oligocene, East Java Sea 333– 334, 335, 336 Nansha Block see Spratley Islands–Dangerous Grounds Block Natal continental fragment 78 Neogene accretionary prism, Java margin 117– 119, 120, 124, 125 Australia– Asia collision dating 245– 250 Sulawesi 179–180 climate change Makassar Strait 319, 321– 331 palynology 321– 331 magmatism, western Sulawesi 172– 173 northern Borneo 92–93 ophiolite, eastern Sulawesi 171– 172 SE Asian gateway 1– 2 volcanism 198 Neostreptognathus 23, 26 New Britain, coral terraces 355– 358 Ngimbang Formation 59, 333, 334 Noele Marl 247, 249, 261 North Banda Sea 89 North China Block, flora and fauna 22 North Equatorial Counter Current 270, 271, 311– 312, 313, 316 North Equatorial Current 270, 271, 272, 311–312 North Makassar Basin 55 North Moluccas 96– 97 North Pacific Equatorial Water 270 North Pacific Intermediate Water 270 North Palawan Block 10, 14 North Sula-Sorong Fault 96, 204 North Sulawesi 86, 94–95 uplift and subsidence 98– 99, 197–198 NSA-1F well, JavaSPAN survey 63, 64, 67 Obi-Bacan Block 10 ocean gateways Miocene, impact on ENSO 305–317 Palaeozoic–Mesozoic, evolution 29, 30– 31 ocean plate stratigraphy, Inthanon Suture 16, 18 oceanography, Indonesian Throughflow 267– 279, 283–284 Ogasawara Plateau, collision and subduction 152 Oligocene East Java Terrane 59, 60 palynology, East Java Sea 333 –344

378

INDEX

Ombai Strait 267, 272, 273 ITF transport 275, 276, 277, 278 Ombilin Basin 44 Lower Cretaceous, andesites 41 ophiolites Cretaceous, Sundaland 41, 84 Neogene, eastern Sulawesi 158, 171–172, 177, 179 Palaeogene, Sulawesi 86 Outer Banda Arc 243 oxygen isotopes, Timor Sea outflow 287–293 P-waves, Sunda arc 141, 142-149, 152 –153 Pacific Ocean eastern cold tongue 308, 309, 310, 316–317 Miocene 308 –313, 316 western equatorial circulation 270 –271 pressure driven flow 273 –274 western warm pool 272, 309, 311, 316 Pagerungan Field 61, 65, 67 Palaeo-Tethys ocean 8, 9, 12, 19, 20 closure 23 main suture 9, 15, 16 subduction 23 Paleocene East Java Terrane 59 unconformity 37, 38, 44–48 Palaeogene accretionary prism, Java margin 117, 119, 120, 122, 124, 125 western Sulawesi 171 palaeothermometry, Mg/Ca, Timor Sea 287, 288 Palaeozoic, East Java Terrane 70 Palaeozoic– Mesozoic 7– 31 continental blocks 8 –14 evolution and palaeogeography 19– 30 ocean gateway evolution 29, 30–31 sutures 14–19 Palelo Group 254 Palu Metamorphic Complex 159, 160, 161 Palu valley central, non-coaxial strain 162, 164–166, 167 northern, non-coaxial strain 161 –164 Palu-Koro Fault 99, 157, 159 metamorphism 160–173 non-coaxial strain 161–168 causes and timing 168– 173 central Palu valley 162, 164–166, 167 neck area 162, 164– 165, 166, 168, 170, 171 northern Palu valley 161–164, 164–165 palynology Late Quaternary, Makassar Strait 319–321, 322 Neogene, Makassar Strait 322– 331 Oligocene East Java Sea 335 –344 everwet elements 342–343 Kerapah type peat swamp 338, 342–343 riparian and seasonal climate elements 338, 340, 342 superwet elements 336– 338, 343 zonation 338, 341, 342 Pancoran Formation 207 Pangaea– Indochina landbridge, Permian 25, 26 Panthalassa 30

Papalang-10 core, Quaternary palynology 320–321 Papayato Volcanic rocks 179, 194 Paracel Islands Block 10, 14 Paternoster microcontinental block 10, 14, 77 Paternoster Platform 92, 94, 95 Paternoster-Lupar lineaments 92 Peladan Formation carbonate platform 194 stratigraphy 181, 182, 184– 185 biostratigraphy 185, 186 Permian arc terranes 9 –10 evolution and palaeogeography 23, 25, 26 flora and fauna 22, 23 Gondwanan blocks 10–11 petroleum East Java Sea 53–54, 55 well penetrations 65, 67– 68 Philippine Sea Plate 7, 8 Philippine Trench 96, 98 pillow basalts 41, 42 plate tectonics, and Cenozoic deformation 91 Pliocene climate change, Makassar Strait, palynology 326– 327, 328– 329, 331 uplift, Gorontalo Bay 194, 196 plutons, I-type, Sumatra 41 Poh Head deformation 97, 198 stratigraphy 183, 187 strike-slip fault 97, 220– 221 volcanism 198 Poko Suture 15, 18 pollen see palynology Pompangeo Schist Complex 159, 161, 169 pressure gradient, interocean 273– 274 Proto-South China Sea, subduction 19, 86, 92 Pseudosweetognathus 23, 26 Qiangtang Block 9, 10, 11 Quaternary, Late, climate change, Makassar Strait 319 –321, 322 radiolarians Karangsambung 41, 42 Mesozoic sutures 16, 17 rainfall 351, 352 Rajang Group 81 Rajang-Crocker fold belts 92 Red River Fault 15, 17 Reed Bank Block 10, 14 Retziella 19, 21 roll-back slab, Sunda arc 152, 154 subduction hinge Banda embayment 89–90, 91, 99 Banda Trench 241, 243 North Sulawesi Trench 198 Savu Basin 234, 235 Romang Island 248 Roo Rise 112, 115 subduction 119, 121, 123, 132 Roti Island 249 Roti Thrust 231, 234, 236–237

INDEX S-waves, Sunda arc 141, 149– 150, 152 –153 Sabah, Neogene 92–93, 94 Sabaoe Calcarenite 247 Sakmarian, Gondwanan blocks 11 Salo Alindau, non-coaxial strain 166, 168 Salodik Formation 185, 207 Sangihe Thrust 205 Sangkarang-16 core, Quaternary palynology 320– 321, 322 Savu Basin 225–239, 226 cross-section 237 –238 deformation structures 231– 233, 236 –237 evolution 234– 239 subsidence and volcanism 234 –236 seismic stratigraphy 225–231 Savu Island 249, 258 Savu Thrust 231, 233, 236–237 Savu-Roti Ridge 232 thrust faults 233, 236 schist, Palu-Koro fault zone 161 –162, 166 Schwaner Mountains 81 Scott Plateau 112, 115, 127, 131, 232, 237, 239 SE Asia 76 continental blocks 8–14, 76– 78 Cretaceous– Cenozoic 14 Devonian 8–9 Jurassic 11–14 Permian 10–11 deformation 91–99 alternative model 99–101 Mesozoic–Cenozoic 76–78 Palaeozoic – Mesozoic 7– 31 evolution and palaeogeography 19– 30 see also Australia– SE Asia collision; Sundaland SE Asia Regional Unconformity 37, 38, 44–48 SE Asian gateway, Neogene 1 –2 SE Java Basin 38 lithostratigraphy 43 sea surface, interocean pressure gradient 273– 274 sea surface temperature 351 impact on coral 355 Miocene 307, 308– 309, 316 sea-level change and deformation 98 eustatic, Late Cretaceous –Palaeogene 47 and ITF 293, 294, 298–299 seamounts, subduction, Java margin 114– 115, 119, 121, 123, 126, 127 Sebangau, Quaternary palynology 320 sediment accretion, Java margin 116–119, 131–132 seismicity Sulawesi 205 Sunda arc subduction slab 139–154 Semanggol Formation 16, 17 Semitau Block 9, 10, 14 SGP-1 well 67– 68 Shan Boundary Suture 9, 18–19 Shan-Thai, terminology 11, 13 Sibumasu Block 9, 10–11, 12, 17, 19 Late Palaeozoic evolution 20 flora and fauna 20, 22, 23 Mesozoic, evolution 27, 84 Sikuleh continental fragment 78

379

Silurian, Gondwanan blocks, evolution and palaeogeography 19, 21 Simao Block 9, 11 Solonker Suture 9 Somali Current 271, 272 Song Da Suture 16, 18 Song Ma Suture 9, 15, 17, 18 Sorong Fault 98, 204, 221, 222 South Banda Sea 89 South China Block 8, 11 flora and fauna 20, 22 South China Sea, opening 77, 78, 81 South China Sea Throughflow (SCST) 276, 279 South China-Indochina margin blocks 14, 17 South China-Indochina-East Malaya superterrane 10 Carboniferous flora 20 South Equatorial Counter Current 270, 271 South Equatorial Current 270, 271, 272, 283, 311 South Java Current 271, 272 South Makassar Basin 55 South Pacific Equatorial Water 270 South Sula-Sorong Fault 204 South Sumatra Basin 38 lithostratigraphy 43 Spratley Islands–Dangerous Ground Block 10, 14 Sra Kaeo Suture 9, 15, 17 ST Alpha-1 well 68, 69 strain, non-coaxial Palu-Koro fault zone 161– 168 causes and timing 168 –173 subduction Banda Embayment 87–88, 89 Banda volcanic fore-arc 86, 242, 243 Cretaceous Sundaland 40–43, 47– 48, 78, 81–82 DLT 44, 46–48 Java Trench 86, 89, 111, 233– 234 Neogene 99 Pacific, Jurassic–Early Cretaceous 81 Sunda arc slab structure 139– 154 roll-back 152, 154 tear 141, 152, 153 SW Borneo 76 subsidence 100 Gorontalo Bay 196– 198 Savu Basin 234– 235 Sukhothai Arc System 9, 12, 17, 23 Sukhothai Block 9 Sula Spur 72, 84, 86, 89– 90 Sula Thrust 205, 217–218 Sulawesi Australia– SE Asia collision 179– 180 Central Metamorphic Belt 169, 171 geological background 157, 158, 159– 160 metamorphic rocks ductile flow 157–173 origin 169 Miocene, carbonates 194 Neogene, magmatism 172– 173 ophiolites 86, 171– 172 Palaeogene 171 tectonic setting 178–180 see also East Sulawesi; North Sulawesi; West Sulawesi

380

INDEX

Sulawesi Sea 272, 273 Sumatra, subduction complexes, Cretaceous 41 Sumba Australia–SE Asia collision 90, 111 stratigraphy 226–228 subsidence 234– 235 volcanism 249, 235– 236 Sumba Island, coral terraces 348, 357, 358 Sumba Ridge 131, 238 normal faults 236 stratigraphy 228, 229– 231 structures 231 thrust faults 231, 233, 236 Sunda arc, subduction slab seismicity imaging 139–154 slab tear 141, 152, 153 Sunda Basin, palynology, correlation with East Java Sea 343– 344 Sunda folds 61 Sunda Shelf 38, 40, 82 Sunda Strait 111 offshore sediment accretion 116–119 Sundaland 8, 37, 38, 40 Australia–SE Asia collision 78– 84 lineaments 83 Cretaceous–Paleocene unconformity 37, 38, 44– 48 deformation 86, 91–99 alternative model 99– 101 east, origin of crust 78–82 heat flow 82–84 lithosphere, seismic velocity 82–84 Mesozoic– Cenozoic assembly 76– 78 sedimentary basins 38 subduction and collision Cretaceous 40– 43, 47–48 termination 82 tectonic evolution 12 uplift, Late Cretaceous 37, 44, 46–48 see also SE Asia superwet climate, East Java Sea 336–338, 343 sutures, Palaeozoic– Mesozoic 14–19 SW Borneo Block 9, 14, 29– 30, 76 flora and fauna 22 Mesozoic 76– 77 origin 81– 82 SW Monsoon Current 271, 272 Sweetognathus 23, 26 Talang Akar Formation, palynology 336, 339, 343– 344 Taliabu, stratigraphy 205, 207 Tamky-Phuoc Son Suture 15, 18 Tanamu Formation 207 Tanjung Formation 45 Tarakan Basin, Neogene palynology 320, 327, 328– 329, 330 Tarim Block 20 flora and fauna 22 Tectonic Collision Zone 247, 251, 253, 262 Telen River, Devonian limestones 81 Tengchong Block 10–11 Terbat Limestone 14

thermocline, mean, and ENSO variability 309–311 Timor allochthon 244, 250– 251 Australia– SE Asia collision 89–90, 253 buoyancy and slab tear 152, 153 continental crust 89– 90 dating 245–250 bathyal continental terrace 248 Bobonaro Melange 247–248 induration and deformation 247 micropalaeontology 245– 248 stratigraphic diachronism 249 uplift and erosion 249, 261 volcanic activity 248–249 autochthon 244 geology 243, 246 field studies 242–243 para-autochthon 243, 244, 247, 256, 258 stratigraphy 244 Tectonic Collision Zone 247, 251, 253, 262 tectonic post-collision processes 241– 262 Timor Passage 267, 272, 273 ITF transport 275, 276, 277, 278– 279 variability 295, 299 Timor Sea 283 Timor Sea ITF outflow 285–300, 286 chronology 288–289, 290 Mg/Ca palaeothermometry 287, 288, 289, 291 salinity reconstructions 288, 289, 291 stable isotopes 287–289, 293 surface water characteristics 289, 291 surface water temperature 288, 289, 291– 293 variability 291–300 deepwater 299 ITCZ and trade-monsoon wind systems 295–298 seasonal controls 293, 295 Timor Trough 111, 130, 131, 232, 237, 239, 259 –261 Tinombo Formation 160, 179 Togian Islands 177, 178, 182 Gorontalo Bay origin and evolution 177– 199 stratigraphy 180– 181, 183– 193 subsidence 96, 99 tectonic setting 178– 180 volcanism 179, 198–199 Tomini Bay 158 tomography, seismic, Sunda arc 140, 142–152, 153 –154 Tonasa Formation, carbonate platform 349, 360 Top Crocker Unconformity 92 trade winds influence on ITF 295–298 and ocean circulation 271, 272 transpression, western Banggai-Sula area 211– 212, 213, 215, 220, 222 Trembling Group 45, 46 Triassic, evolution and palaeogeography 27, 29, 76 tsunamis Bali 153, 154 impact on coral 354–355 Java margin 115, 132 Tuban Formation 333, 334 tuff, Lonsio Formation 187, 198

INDEX Una-Una volcano 177– 178, 195, 198 unconformities Cretaceous–Paleocene 37, 38, 44–48 departure and arrival, East Java Terrane 68– 70 see also SE Asia Regional Unconformity; Top Crocker Unconformity uplift Bongka Formation 194, 196, 197–198 collisional v. mantle-driven 39–40 Java margin 123 northern Borneo and North Moluccas 97– 98 Sulawesi 98– 99 Sundaland 37, 44, 46–48 exhumation data 44, 47 Timor 249, 261 Vietnam, offshore, Australia–SE Asia collision 78, 81 Viqueque Group 247, 248, 250, 261 Visean biota 20 palaeogeography 24 Vjalovognathus 23, 26 volcaniclastics Lonsio Formation 185, 187, 198 Savu Basin 235 –6, 238 volcanism Cretaceous 40–41 Gorontalo Bay 179, 198– 199 Savu Basin 235 –236 Timor 248– 249, 250 Wai Luli Formation 256, 257 Walea Formation

381

stratigraphy 181, 182, 183 –184 Togian Island basement 194 Wallace Line 1, 10 Wallacea 1, 75, 89 Wana Metamorphic Complex 159, 160– 161 Weber Deep 89, 91 West Burma Block 8, 9, 12, 18–19, 23, 29 origin 77– 78, 84 West Natuna Basin 38 lithostratigraphy 43 Neogene, palaeoclimate 329, 330, 331 West Sulawesi, deformation 93–94, 95, 98, 101 West Sulawesi microcontinental block 10, 14, 157 collision with East Java Terrane 43 West Sulawesi plutono-volcanic arc 158 metamorphic rocks 169, 171 North Sulawesi Province, tectonic setting 178, 179 West Sulawesi Province tectonic setting 178– 179 Togian basement 194 West Sumatra Block 8, 9, 18– 19, 23, 84 Western Australia, comparison with East Java Terrane 71–72 Wetar Island, volcanism 248–249, 250, 251 Wetar Strait 257– 258 Wetar Suture 251, 253, 258–259, 261, 262 wind, and ocean circulation 269–270 island rule 269–270 ITF 274 –275 Woyla Group 41, 42 Woyla Suture 9 Wyrtki jets 272

Collision between Australia and SE Asia began in the Early Miocene and reduced the former wide ocean between them to a complex passage which connects the Pacific and Indian Oceans. Today, the Indonesian Throughflow passes through this gateway and plays an important role in global thermohaline flow. The surrounding region contains the maximum global diversity for many marine and terrestrial organisms. Reconstruction of this geologically complex region is essential for understanding its role in oceanic and atmospheric circulation, climate impacts, and the origin of its biodiversity. The papers in this volume discuss the Palaeozoic to Cenozoic geological background to Australia and SE Asia collision. They provide the background for accounts of the modern Indonesian Throughflow and oceanographic changes since the Neogene, and consider aspects of the region’s climate history.