Channel flow, ductile extrusion and exhumation in continental collision zones

Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones

The Geological Society of London Books Editorial Committee B. PANKHURST (UK) (CHIEF EDITOR)

Society Books Editors J. GREGORY ( U K ) J. GRIFFITHS ( U K ) J. HOWE (UK) P. LEAT (UK) N. ROBINS (UK) J. TURNER (UK)

Society Books Advisors M. BROWN ( U S A ) E. BUFFETAUT (France) R. GIERE ( G e r m a n y ) J. GLUYAS (UK) D. STEAD (Canada) R. STEPHENSON (Netherlands)

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: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) 2006. Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268. BEAUMONT, C. NGUYEN, U. JAMIESON, R. & ELLIS, S. 2006. Crustal flow in large hot orogens. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 91-145.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 268

Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones EDITED BY

R. D. LAW Virginia Tech, USA

M. P. SEARLE University of Oxford, UK and L . GODIN Queen's University, Canada

2006 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 9000. 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. +44 (0)20 7434 9944; Fax +44 (0)20 7439 8975; E-mail: enquiries @geolsoc.org.uk. 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. +44 (0)1225 445046, Fax +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. 9 The Geological Society of London 2006. 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, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/06/$15.00.

British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 1-86239-209-9 ISBN 13 978-1-86239-209-0 Typeset by Techset Composition, Salisbury, UK Printed by Cromwell Press, Trowbridge, 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 +1 800-972-9892 Fax +1 802-864-7626 Email gsl.orders @aidcvt.com

For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel +1 918-584-2555 Fax § 1 918-560-2652 Email [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 !10 002, India Orders: Tel. +91 11 2327-9113/2326-4180 Fax +91 11 2326-0538 E-mail affiliat@ vsnl.com

Contents

Dedication

vii

Memorial for Doug Nelson

ix

Acknowledgements

x

Introduction

GODIN, L., GRUJIC,D., LAW, R. D. & SEARLE,M. P. Channel flow, ductile extrusion and exhumation in continental collision zones: an introduction GRUJIC, D. Channel flow and continental collision tectonics: an overview

l 25

Evolution of ideas on channel flow and ductile extrusion in the H i m a l a y a - T i b e t a n Plateau system

KLEMPERER, S. L. Crustal flow in Tibet: geophysical evidence for the physical state of Tibetan lithosphere, and inferred patterns of active flow

39

HODGES, K. V. A synthesis of the Channel Flow-Extrusion hypothesis as developed for the Himalayan-Tibetan orogenic system

71

Modeling channel flow and ductile extrusion processes

BEAUMONT, C., NGUYEN, M. H., JAMIESON, R. A. & ELLIS, S. Crustal flow modes in large hot orogens

91

MEDVEDEV, S. & BEAUMONT, C. Growth of continental plateaus by channel injection: models designed to address constraints and thermomechanical consistency

147

JAMIESON, R. A., BEAUMONT, C., NGUYEN, M. H. & GRUJIC, D. Provenance of the Greater Himalayan Sequence and associated rocks: predictions of channel flow models

165

GRASEMANN, B., EDWARDS, M. A. & WIESMAYR, G. Kinematic dilatancy effects on orogenic extrusion

183

JONES, R. R., HOLDSWORTH, R. E., HAND, M. & GOSCOMBE, B. Ductile extrusion in continental collision zones: ambiguities in the definition of channel flow and its identification in ancient orogens

201

WILLIAMS, P. F., JIANG, D. & LIN, S. Interpretation of deformation fabrics of infrastructure zone rocks in the context of channel flow and other tectonic models

221

Geological constraints on channel flow and ductile extrusion as an important orogenic process

Himalaya- Tibetan Plateau HARRISON, T. M. Did the Himalayan Crystallines extrude partially molten from beneath the Tibetan Plateau?

237

ROBINSON, D. M. & PEARSON, O. N. Exhumation of Greater Himalayan rock along the Main Central Thrust in Nepal: implications for channel flow

255

vi

CONTENTS

GODIN, L., GLEESON,T. P., SEARLE,M. P., ULLRICH,T. D. & PARRISH,R. R. Locking of southward extrusion in favour of rapid crustal-scale buckling of the Greater Himalayan sequence, Nar valley, central Nepal

269

SCAILLET,B. & SEARLE,M. P. Mechanisms and timescales of felsic magma

293

segregation, ascent and emplacement in the Himalaya

ANNEN, C. & SCAILLET,B. Thermal evolution of leucogranites in extensional faults:

309

implications for Miocene denudation rates in the Himalaya

WANG, Y., LI, Q. & GUOSHENG,Q. 4~

thermochronological constraints on the cooling and exhumation history of the South Tibetan Detachment System, Nyalam area, southern Tibet

327

SEARLE, M. P., LAW, R. D. & JESSUP, M. J. Crustal structure, restoration and evolution

355

of the Greater Himalaya in Nepal-South Tibet: implications for channel flow and ductile extrusion of the middle crust

JESSUP, M. J., LAW,R. D., SEARLE,M. P. & HUBBARD,M. S. Structural evolution and

379

vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/Nepal: implications for orogen-scale flow partitioning

HOLLISTER, L. S. & GRUJIC,D. Pulsed channel flow in Bhutan

415

CAROSI, R., MONTOMOLI,C., RUBATTO,D. & VISON.~,D. Normal-sense shear zones in the core of the Higher Himalayan Crystallines (Bhutan Himalaya): evidence for extrusion?

425

LEE, J., MCCLELLAND,W., WANG,Y., BLYTHE,A. & MCWILLIAMS,M. Oligocene-Miocene middle crustal flow in southern Tibet: geochronology of Mabja Dome

445

AOYA, M., WALLIS, S. R., KAWAKAMI, T., LEE, J., WANG, Y. & MAEDA, H. The Malashan gneiss dome in south Tibet: comparative study with the Kangmar dome with special reference to kinematics of deformation and origin of associated granites

471

Hellenides and Appalachians

XYPOLIAS, P. & KOKKALAS,S. Heterogeneous ductile deformation along a mid-crustal extruding shear zone: an example from the External Hellenides (Greece)

497

HATCHER, R. D. JR. & MERSCHAT,A. J. The Appalachian Inner Piedmont:

517

an exhumed strike-parallel, tectonically forced orogenic channel

Canadian Cordillera BROWN, R. L. & GIBSON H. D. An argument for channel flow in the southern Canadian Cordillera and comparison with Himalayan tectonics

543

CARR, S. D. & SIMONY,P. S. Ductile thrusting versus channel flow in the

561

southeastern Canadian Cordillera: evolution of a coherent crystalline thrust sheet KUIPER, Y. D., WILLIAMS,P. F. & KRUSE, S. Possibility of channel flow in the southern Canadian Cordillera: a new approach to explain existing data

589

Index

613

THIS VOLUME IS DEDICATED TO THE W O R K OF KARL DOUGLAS NELSON 26 March 1953-17 August 2002

Doug supervising INDEPTH-III field operations from the running-board of his field vehicle, on the banks of Siling Tso, central Tibet, summer 1998.

Memorial for Doug Nelson

Doug Nelson, the Jessie Page Heroy Professor of Earth Sciences and Department Chair at Syracuse University, died as he was reaching new heights in an increasingly distinguished career. His sudden and untimely death from heart failure robbed us all of many insights and papers that would have been forthcoming in decades still to come. His most visible legacy is a new understanding of Tibet, resulting in large part from the work that he led and supervised as the intellectual leader of the INDEPTH (International Deep Profiling of Tibet and the Himalaya) program. Doug graduated from Cornell University with a BS in 1975. He received his PhD as a structural geologist working on the Newfoundland Appalachians from SUNY Albany in 1979, at a time when that department stressed the continuum from field observations to plate-tectonic synthesis. After a brief post-doctorate at Otago University, New Zealand, Doug returned to Cornell University to join COCORP (Consortium for Continental Reflection Profiling). There he learned to interpret deep seismic reflection data, and to value geophysics for the study of large-scale processes in mountain belts. Doug became a proponent of taking the COCORP methodology to the greatest of all mountain belts, the Himalaya. When Doug first went to Tibet in the 1980s, even the basic crustal architecture was uncertain; for example, whether the plateau crust was thick because two normal crusts had been vertically stacked, or because a single crust had been thickened by pure shear. Doug used the pilot 1992 INDEPTH reflection profile across the Himalaya to show that the Indian foreland was subducting beneath southern Tibet along an active master detachment-named by Doug the Main Himalayan Thrust-to depths from which it could confidently be extrapolated to underthrust the Indus-Tsangpo suture. This result only fuelled speculation on the ultimate northward limit of penetration of Indian crust beneath Tibet, and the fate of the subducting continental crust in the suture zone. Doug had already addressed this mass-balance problem for the overthickened crust of other continent-continent collisions, arguing from reflection profiles for delamination in the Appalachians, but for a phase change at the Moho in the Trans-Hudson orogen.

The second and third INDEPTH field campaigns in 1994 and 1998 progressed into interior Tibet, and added to the original reflection profiling additional scientific techniques: wide-angle and refraction seismology, broadband teleseismic recording, magnetotelluric observations and field geology. Doug actively participated in all these separate programmes, and more than anyone was the enthusiastic integrator in the large multi-national group of investigators (from the USA, China, Canada and Germany), best able to synthesize seemingly disparate observations from all the techniques. Doug's intellectual legacy includes a generation of students and colleagues, at Syracuse, Cornell and the other INDEPTH institutions, who now regard such broad interdisciplinary science as the norm.

The most serendipitous discovery of the INDEPTH project, and the most consequential, was the crustal melts in southern Tibet, recognized independently by all the geophysical techniques employed. Doug, by now professor at Syracuse University, pushed forward studies on the structure of the suture zone that, he believed, could be regarded as a region of mixing of the crusts of India and Tibet, which then extruded towards India, crystallizing as leucogranites now exposed at the erosional front of the High Himalaya. Doug's recognition of partial melting in interior Tibet implied at once that this region is hot, and hence mobile. Much of the recent popularity of models of continent-continent collisions in which material transport is dominated by middle and lower crustal flow must be attributed to observations such as these that bear directly on crustal viscosity. Though such flow is now widely accepted, this is only a recent paradigm shift. Doug was a master of his trade, able to integrate his training as a field geologist with the big picture drawn from his regional geophysical surveys. Although he did not live to write a final synthesis of the INDEPTH results, our picture of Tibet and hence of all continent-continent collisions has changed and grown far richer, a legacy that enriches the Earth Sciences community.

Simon Klemperer

Acknowledgments

The papers in this volume arise from a conference held at the Geological Society of London on 6-7 December 2004. The conference was attended by 102 participants from 13 countries. We thank all the conference participants for a fascinating series of talks and posters and for lively discussion. We would also like to thank all contributors for dealing with editorial decisions courteously and promptly, and the Geological Society Publishing House for their help and advice. We gratefully acknowledge the following colleagues who helped with the reviewing of manuscripts submitted for this volume. Tom Argles, Open University, UK Jean-Philippe Avouac, California Institute of Technology, USA Chuck Bailey, College of William & Mary, USA Rebecca Bendick, Cambridge University, UK Andy Bobyarchick, University of North Carolina at Charlotte, USA Dick Brown, Carleton University, Canada Roger Buck, Columbia University, USA Jean-Pierre Burg, ETH-Zurich, Switzerland Mike Cosca, Universit6 de Lausanne, Switzerland Alexander Cruden, University of Toronto, Canada Nick Culshaw, Dalhousie University, Canada Peter DeCelles, University of Arizona, USA Declan De Poar, Boston University, USA John Dewey, University of California at Davis, USA Mike Edwards, University of Vienna, Austria Taras Gerya, ETH-Zurich, Switzerland Dan Gibson, Simon Fraser University, Canada Bernhard Grasseman, University of Vienna, Austria John Grocott, Kingston University, UK Djordje Grujic, Dalhousie University, Canada St6phane Guillot, Universit6 de Lyon, France Brad Hacker, University of California at Santa Barbara, USA Nigel Harris, Open University, UK Mark Harrison, Australian National University, Australia Kip Hodges, Massachusetts Institute of Technology, USA

Greg Houseman, Leeds University, UK Mary Hubbard, Kansas State University, USA Becky Jamieson, Dalhousie University, Canada Mike Johnson, University of Edinburgh, UK Richard Jones, University of Durham, UK Bill Kid, State University of New York at Albany, USA Simon Klemperer, Stanford University, USA Jeff Lee, Central Washington University, USA Sergei Medvedev, Freie Universit~it Berlin, Germany Mike Murphy, University of Houston, USA Steve Noble, British Geological Survey, UK Randy Parrish, British Geological Survey, UK Nick Petford, Kingston University, UK John Platt, University of Southern California, USA Uwe Ring, Johannes Gutenberg Universitat, Germany Martin Robyr, Universit6 de Lausanne, Switzerland Philip Simony, University of Calgary, Canada Christian Teyssier, University of Minnesota, USA Basil Tikoff, University of Wisconsin- Madison, USA Peter Treloar, Kingston University, UK Jonathan Turner, University of Birmingham, UK Simon Wallis, Nagoya University, Japan Donna Whitney, University of Minnesota, USA Paul Williams, University of New Brunswick, Canada Paris Xypolias, University of Patras, Greece

Channel flow, ductile extrusion and exhumation in continental collision zones: an introduction L. G O D I N 1, D. G R U J I C 2, R. D. L A W 3 & M. P. S E A R L E 4

1Department of Geological Sciences & Geological Engineering, Queen's University, Kingston, Ontario, K7L 3N6, Canada (e-mail: [email protected]) 2Department of Earth Sciences, Dalhousie University, Halifax, Nova Scotia, B3H 4J1, Canada 3Department of Geological Sciences, Virginia Tech., Blacksburg, VA 24061, USA 4Department of Earth Sciences, Oxford University, Oxford, OX1 3PR, UK Abstract: The channel flow model aims to explain features common to metamorphic hinterlands

of some collisional orogens, notably along the Himalaya-Tibet system. Channel flow describes a protracted flow of a weak, viscous crustal layer between relatively rigid yet deformable bounding crustal slabs. Once a critical low viscosity is attained (due to partial melting), the weak layer flows laterally due to a horizontal gradient in lithostatic pressure. In the Himalaya-Tibet system, this lithostatic pressure gradient is created by the high crustal thicknesses beneath the Tibetan Plateau and 'normal' crustal thickness in the foreland. Focused denudation can result in exhumation of the channel material within a narrow, nearly symmetric zone. If channel flow is operating at the same time as focused denudation, this can result in extrusion of the mid-crust between an upper normal-sense boundary and a lower thrust-sense boundary. The bounding shear zones of the extruding channel may have opposite shear sense; the sole shear zone is always a thrust, while the roof shear zone may display normal or ttuqast sense, depending on the relative velocity between the upper crust and the underlying extruding material. This introductory chapter addresses the historical, theoretical, geological and modelling aspects of channel flow, emphasizing its applicability to the Himalaya-Tibet orogen. Critical tests for channel flow in the Himalaya, and possible applications to other orogenic belts, are also presented.

The hinterlands of collisional orogens are often characterized by highly strained, high-grade metamorphic rocks that c o m m o n l y display features consistent with lateral crustal flow and extrusion of material from mid-crustal depths towards the orogenic foreland. A recent model for lateral flow of such weak mid-crustal layers has b e c o m e widely known as the 'channel flow' model. The channel flow model has matured through efforts by several research groups and has also been applied to a variety of geodynamic settings. Thermal-mechanical modelling of collision zones, including the H i m a l a y a n - T i b e t a n system, has brought the concept of channel flow to the forefront of orogenic studies. Original contributors to the concept of channel flow initiated an important paradigm shift (Kuhn 1979), from geodynamic models of continental crust with finite rheological layering to the more encompassing channel flow model. This time-dependent mid- to lower crustal flow process, which will be reviewed in this chapter, may progress into foreland fold-and-thrust tectonics in the upper crust, thereby providing a spatial

and temporal link between the early d e v e l o p m e n t of a metamorphic core in the hinterland and the foreland fold-and-thrust belt at shallower structural levels. Outcomes and implications of such a viscous flowing middle to lower crust include a dynamic coupling between mid-crustal and surface processes, and limitations to accurate retro-deformation of orogens (non-restorable orogens, e.g. Jamieson et al. 2006). This Special Publication contains a selection of papers that were presented at the conference 'Channel flow, extrusion, and exhumation of lower to mid-crust in continental collision zones' hosted by the Geological Society of L o n d o n at Burlington House, in December 2004. Because most of the ongoing debate on crustal flow focuses on the Cenozoic age H i m a l a y a - T i b e t collisional system, some of the key questions that are addressed in this v o l u m e include the following. 9 Does the model for channel flow in the Himal a y a - T i b e t system concur with all available geological and geochronological data?

ChannelFlow,DuctileExtrusionand Exhumation in Continental CollisionZones. Geological Society, London, Special Publications, 268, 1-23.

From: LAW,R. D., SEARLE,M. P. & GODIN,L. (eds)

0305-8719/06/$15.00

9 The Geological Society of London 2006.

2

L. GODIN ETAL.

9 How do the pressure-temperature-time (P-T-t) data across the crystalline core of the Himalaya fit with the proposed channel flow? 9 Are the microstructural fabric data (pure shear and simple shear components) compatible with crustal extrusion (thickening or thinning of the slab)? 9 If the channel flow model is viable for the Himalaya-Tibet system, what may have initiated channel flow and ductile extrusion? 9 Why did the extrusion phase of the Himalayan metamorphic core apparently cease during the late Miocene-Pliocene? 9 Are some of the bounding faults of the potential channel still active, or were they recently active? 9 Is the Himalayan channel flow model exportable to other mountain ranges? This introductory paper addresses the historical, theoretical, geological and modelling aspects of crustal flow in the Himalaya-Tibet orogen. Critical tests for crustal flow in the Himalaya, and possible applications to other orogenic belts, are presented and difficulties associated with applying these tests are discussed. Personal communication citations (pers. comm. 2004) identify comments expressed during the conference.

The Himalaya-Tibetan plateau system The Himalaya-Tibet system initiated in Early Eocene times, following collision of the Indian and Eurasian plates (see Hodges (2000) and Yin & Harrison (2000) for reviews). The collision resulted in closure of the Tethyan Ocean, southward imbrication of the Indian crust, and northward continental subduction of Indian lower crust and mantle beneath Asia. The collision thickened the southern edge of the Asian crust to 70 kin, and created the Tibetan Plateau, the largest uplifted part of the Earth's surface with an average elevation of 5000 m (Fielding et al. 1994). The Himalayan orogen coincides with the 2500kin-long topographic front at the southern limit of the Tibetan Plateau. It consists of five broadly parallel lithotectonic belts, separated by mostly north-dipping faults (Fig. 1). The Himalayan metamorphic core, termed the Greater Himalayan sequence (GHS), is bounded by two parallel and opposite-sense shear zones that were both broadly active during the Miocene (Hubbard & Harrison 1989; Searle & Rex 1989; Hodges et al. 1992, 1996). The Main Central thrust (MCT) zone marks the lower boundary of the GHS, juxtaposing the metamorphic core above the underlying Lesser Himalayan sequence. The South Tibetan detachment (STD) system defines the upper boundary

roof fault of the GHS, marking the contact with the overlying unmetamorphosed Tethyan sedimentary sequence. The apparent coeval movement of the MCT and STD, combined with the presence of highly sheared rocks and high grade to migmatitic rocks within the GHS, has led many workers to view the GHS as a north-dipping, southward-extruding slab of mid-crustal material flowing away from the thick southern edge of the Tibetan Plateau, towards the thinner foreland fold-thrust belt.

Dynamics of channel flow The concepts of crustal extrusion and channel flow originated in the continental tectonics literature in the early 1990s. Unfortunately, these two processes are often referred to interchangeably without justification. One of the main points that emerged from the Burlington House conference was that a distinction between channel flow and crustal extrusion must be made. Parallel versus tapering bounding walls on channel flow and/or extrusion processes, and how these processes may replenish over time, are two resolvable parameters that are critical for distinguishing channel flow from extrusion. Brief definitions and overviews of the two processes are presented below. A more detailed overview of the mechanics of the related processes is provided by Grujic (2006). Channel flow

Channel flow involves a viscous fluid-filled channel lying between two rigid sheets. The viscous fluid between the sheets is deformed through induced shear and pressure (or mean stress) gradients within the fluid channel (Fig. 2; e.g. Batchelor 2000; Turcotte & Schubert 2002). The weak layer flows laterally due to a horizontal gradient in lithostatic pressure; gravity is therefore the driving force. The finite displacement depends on the geometry of the channel, viscosity, and displacement rate of the bounding plates. In situations where the channel walls are non-parallel, the (non-lithostatic) pressure gradient may cause high rates of buoyant return flow of the channel material, provided that the viscosity is low enough (Mancktelow 1995; Gerya & Strckhert 2002). The simplest qualitative characteristic of the channel flow model is that the velocity field consists of a hybrid between two end-members: (1) Couette flow (simple shear) between moving plates where the induced shear across the channel produces a uniform vorticity across the channel (Fig. 2A, left); and (2) Poiseuille flow (also known as 'pipe-flow' effect) between

INTRODUCTION

3

Fig. 1. Simplified geological map of the Himalayan orogen, with general physiographic features of the HimalayaTibet system (inset). The Greater Himalayan sequence is bounded above by the north-dipping top-to-the-north South Tibetan detachment system (STDs), and below by the north-dipping top-to-the-south Main Central thrust zone (MCTz). The Himalaya is arbitrarily divided into four sections to facilitate age compilations (see Fig. 4). ITSZ: Indus-Tsangpo Suture Zone.

stationary plates in which the induced pressure gradient produces highest velocities in the centre of the channel and opposite shear senses for the top and bottom of the channel (e.g. Mancktelow 1995; Fig. 2A, right). Quantitative analyses of channel flow between moving or stationary boundaries (e.g. Batchelor 2000; Turcotte & Schubert 2002) have been applied to a wide range of geodynamic processes including: (1) asthenospheric counterflow (Chase 1979; Turcotte & Schubert 2002); (2) mechanics of continental extension (Kusznir & Matthews 1988; Block & Royden 1990; Birger 1991; Kruse et al. 1991; McKenzie et al. 2000; McKenzie & Jackson 2002); (3) continental plateau formation and evolution, both in extension and compression (Zhao & Morgan 1987; Block & Royden 1990; Wernicke 1990; Bird 1991; Fielding et al. 1994; Clark & Royden 2000; McQuarrie & Chase 2000;

Hodges et al. 2001; Shen et al. 2001; Husson & Sempere 2003; Clark et al. 2005; Gerbault et al. 2005; Medvedev & Beaumont 2006); (4) tectonics of large continent-continent collision orogens (Johnston et al. 2000; Beaumont et al. 2001, 2004, 2006; Grujic et al. 2002, 2004; Williams & Jiang 2005; (5) metamorphic histories in large, hot, collisional orogens (Jamieson et al. 2002, 2004, 2006); (6) subduction zone flow regimes under both lithostatic and overpressured conditions (Bird 1978; England & Holland 1979; Shreve & Cloos 1986; Peacock 1992; Mancktelow 1995; Gerya et al. 2002; Gerya & St6ckhert 2002); and (7) deformation along passive continental margins in the presence of salt layers (Gemmer et al. 2004; Ings et al. 2004). For all these examples, with the exception of salt tectonics, the most likely cause for weakening in the channel is partial melting. In the case of salt tectonics, flow

4

L. GODIN E T A L .

Fig. 2. Schematic diagram of the flow pattern in a viscous channel of width h. The viscosity of channel material is lower than the viscosity of rocks in the hanging wall and in the footwall (/,Zh > /-s < /.Lj~). Velocity distributions are shown relative to a reference frame attached to the hanging wall. The vorticity values (rotational component of the flow profile) are schematically indicated by the width of the black bar: the wide bar segment indicates a high simple shear component; the narrow bar segment indicates a high pure shear component. Only the absolute value of the vorticity is indicated regardless of whether it is positive (sinistral simple shear) or negative (dextral simple shear). (A) End-members of flow in a channel: left, Couette flow with velocity profile caused by shearing (toe: vorticity in pure Couette flow); right, Poiseuille flow with velocity profile caused by pressure gradient within the channel (top: vorticity in pure Poiseuille flow). (B) For a given velocity of the subducting plate and channel width there is a critical viscosity of the channel material below which the Poiseuille flow will counteract the shear forces and cause return flow (negative velocity) and therefore exhumation of that part of the channel material. The part of the channel that remains dominated by the induced shear (positive velocities) will continue being underplated (w~: vorticity in a hybrid channel flow). From Grujic et al. (2002), after Mancktelow (1995) and Turcotte & Schubert (2002).

is due to the inherent low viscosity of salt under upper crustal conditions. The application and significance of channel flow in continent-continent collisional settings is becoming progressively more refined, yet remains controversial. Evidence for channel flow and/or ductile extrusion of mid-crustal rocks from the geologically recent Himalaya-Tibet orogen (Grujic et al. 1996, 2002; Searle & Szulc 2005; Carosi et al. 2006; Godin et al. 2006; Hollister & Grujic 2006; Jessup et al. 2006; Searle et al. 2006) and related geodynamic models (Beaumont et al. 2001, 2004, 2006; Jamieson et al. 2004, 2006) are vigorously disputed (e.g. Hilley et al. 2005; Harrison 2006; Williams et al. 2006), while there are still few documented examples from older orogens (Jamieson et al. 2004; White et al. 2004; Williams & Jiang 2005; Brown & Gibson 2006; Carr & Simony 2006; Hatcher & Merschat 2006; Kuiper et al. 2006; Xypolias & Kokkalas 2006). Extrusion

Extrusion is defined as the exhumation process of a channel (or the shallower part of it) operating at a localized denudation front. Channel flow and extrusion can operate simultaneously, with lateral tunnelling occurring at depth, while extrusion occurs at the front of the system, at progressively shallower crustal levels (Fig. 3). The focused denudation

results in exhumation of the channel material within a narrow, nearly symmetric zone; the extruded channel is characterized by an upper normal-sense boundary, and a lower thrust-sense boundary. We present brief reviews of the four major channel flow and extrusion models. Figure 3A presents a schematic overview of the kinematic relationships between channel flow and extrusion processes. The weak crustal channel flow (Fig. 3A, no. 5) is localized structurally below the 750~ isotherm, where melting starts (Fig. 3A, no. 8). Material points affected by a Poiseuille flow within the channel (Fig. 3A, no. 7) traverse the rheological boundary at the tip of the channel (Fig. 3A, no. 9), and will continue their exhumation by extrusion of the palaeo channel (Fig. 3A, no. 10). It follows that the ductile outward channel flow converts to ductile extrusion where rock motion is balanced by surface erosion (Fig. 3A, no. 13). The extrusion is aided by focused erosion along the mountain front (e.g. Grujic et al. 2002; Vannay et al. 2004), the rate of extrusion/exhumation being proportional to the rate of flow in the mid-crustal channel. Extrusion of the crustal layer is assumed to occur along two bounding shear zones (Fig. 3A, nos 11-12). The bounding shear zones may have opposite shear sense; the sole shear zone is always a thrust, while the roof shear zone may display normal or thrust sense, depending on the relative velocity

INTRODUCTION

5

Fig. 3. (A) Schematic diagram of kinematic relationship between channel flow and extrusion of a palaeo-channel. All material depicted belongs to the underthrusting plate. 1, Lithospheric mantle; 2, lower crust; 3, mid-crust; 4, upper crust; 5, weak crustal channel; 6, isotherms (taken from Beaumont e t al. 2004), 7, schematic velocity profile during return channel flow; 8, 750 ~ C isotherm structurally below which partial melt starts; 9, theological tip of the channel: at lower temperatures (for a given channel width and pressure gradient) Couette flow will dominate and all the material will be underthrust; 10, extruding crustal block (palaeo-channel): if the theological tip is at steady state, material points may move through this tip and pass from the weak crustal channel into the extruding block; 11, lower shear zone of the extruding crustal block (dominantly thrust-sense kinematics); 12, upper shear zone of the extruding crustal block (dominantly normal-sense kinematics); 13, focused surface denudation (controlled by surface slope and, as in the Himalaya, by orographic precipitation at the topographic front). (B)-(D) Possible strain distribution in an extruding crustal block; possible end-members (inspired by Grujic et al. 1996; Grasemann et al. 1999). (B) Rigid block with high concentration of strain along the boundaries (Hodges et al. 1996). (C) Ductile block deforming by pervasive simple shear. (D) Ductile block deforming by general shear with a pure shear component increasing towards the bottom of the wedge, as well as with time following a 'decelerating strain path' (Grasemann et al. 1999; Vannay & Grasemann 2001).

between the upper crust (Fig. 3A, no. 4) and the underlying extruding material. This is similar to the asymmetric thrust exhumation/extrusion mode described by Beaumont et al. (2004). The extruding mid-crustal layer can be slab- or wedge-shaped, depending on the parallelism of the bounding shear zones, and advances towards the foreland. As the channel material is extruded, deformation is pervasively distributed within, or at the boundaries of the crustal layer. A concentration of deformation along the boundaries results in extrusion of a rigid crustal wedge (Fig. 3B; e.g. Burchfiel & Royden 1985; Hodges et al. 1992). This type of extrusion cannot be a long-lived

geological process, but rather is probably a transient event (cf. Williams et al. 2006). Alternatively, deformation that is distributed throughout the wedge results in ductile extrusion (Fig. 3C; Grujic et al. 1996). The vorticity of flow within the extruded crust may be a perfect simple shear (Fig. 3C), or more likely a general shear combining components of simple shear and pure shear (Fig. 3D; Grujic e t a l . 1996; Grasemann etal. 1999; Vannay & Grasemann 2001; Law et al. 2004; Jessup et al. 2006; but cf. Williams et al. 2006). During extrusion, the crustal slab or wedge cools from mid-crustal ductile flow to upper crustal brittle

6

L. GODIN ETAL.

conditions where deformation is partitioned into discrete faults. This concept resembles the buoyancydriven extrusion of a crustal slab within a subduction zone (Chemenda e t a l . 1995). The differences between these two models include the size of the extruding wedge and nature of the primary driving forces. Analogue models of Chemenda et al. (1995) demonstrate that syncollisional exhumation of previously subducted (underthrust) crustal material can occur due to failure of the subducting slab. In this model, erosional unloading causes the buoyant upper crust to be exhumed (at a rate comparable to the subduction rate), producing a normal-sense movement along the upper surface of the slab. This model, however, regards the exhumed crustal slice as a rigid slab bounded below and above by well-defined thrust and normal faults. One of the key Himalayan problems is whether the GHS represents extrusion of a complete section of the mid-crust, with the S T D - M C T surfaces representing potential channel-bounding structures, or whether the GHS is simply an extruded segment of a cooling channel, with the S T D - M C T surfaces being more akin to roof and sole faults bounding a thrust duplex (Yin 2002). Furthermore, the GHS-bounding faults exposed at the topographic surface could be associated with late-stage exhumation of the GHS, rather than the original channel formed at depth beneath the Tibetan Plateau (Jessup et al. 2006). Related problems also concern the origin of fabrics within the GHS; are they related to flow during channelling, extrusion, or could they pre-date the Himalayan event? Another unresolved question is whether 'the currently exposed GHS more closely resembles an exhumed plugged channel, with little extrusion during exhumation, or whether there was (and perhaps still is) active extrusion at the surface (e.g. Hodges et al. 2001; Wobus et al. 2003)' (Beaumont et al. 2004, p. 26). Exhumation

Exhumation is defined as the displacement of rocks with respect to the topographic surface (England & Molnar 1990), and requires either removal of the overburden (e.g. by erosion, normal faulting, vertical lithospheric thinning) or transport of material through the overburden (e.g. by diapirism, buoyancy-driven return flow in subduction zones) (see reviews by Platt 1993; Ring et al. 1999). In the context of channel flow, exhumation of the channel occurs by a balance between orographically and topographically enhanced focused erosion and extrusion on the southern slopes of the Himalaya. The channel's tunnelling capacity may be dramatically reduced as it is deflected upward during

exhumation and cooling (Fig. 3A; Beaumont et al. 2004). Although exhumation of the GHS in the Himalaya may be associated with southward extrusion along coeval S T D - M C T bounding faults, it may also be locally enhanced by post-MCT warping of the GHS and localized erosion following cessation of extrusion (Thiede et al. 2004; Vannay et al. 2004; Godin et al. 2006), and growth of duplexes in the footwall of the MCT during the middle Miocene (Robinson & Pearson 2006). Exhumation of tunnelling material (not to be confused with extruded palaeo-channel material) will only occur if the active channel flow breaks through to the topographic surface. In the numerical models of channel flow, this is determined by the rheological properties of the upper-middle crust, frictional strength, and degree of advective thinning of the surface boundary layer (Beaumont et al. 2004). Conceptually, a threshold exhumation rate must be achieved to keep the channel sufficiently hot so that the material does not 'freeze' until it is close to the topographic surface (Beaumont et al. 2004). This situation may have been achieved in the two Himalayan 'syntaxes': the Nanga ParbatHaramosh massif (e.g. Craw et al. 1994; Zeitler et al. 2001; Butler et al. 2002; Koons et al. 2002; Jones et al. 2006) and the Namche Barwa massif (e.g. Burg et al. 1997, 1998; Burg 2001; Ding et al. 2001).

Requirements and characteristics of channel flow The following is a list of geological characteristics of channel flow, based on field observation and geodynamic modelling, and field criteria for recognizing an exhumed channel from the geological past (Searle & Szulc 2005; Searle et al. 2003, 2006). The criteria used to identify an active channel are based on geological and geophysical data (Nelson et al. 1996; S. Klemperer pers. comm. 2004; Klemperer 2006) and landscape analyses (Fielding et al. 1994; Clark et al. 2005). (1) (2)

(3)

(4)

A crustal package of lower viscosity material bounded by higher viscosity rocks. A plateau with well-defined margins (or a significant contrast in crustal thickness) to produce a horizontal gradient in lithostatic pressure. Coeval movement on shear zones with thrust and normal-fault geometry that bound the channel flow zone. Kinematic inversion along the roof shear zone: earlier reverse-sense motion resulting from underthrusting (Couette flow) followed by normal-sense shearing resulting from back flow (by dominant Poiseuille flow) in

INTRODUCTION

(5)

(6)

the channel, and/or by normal-sense motion on shear zones and brittle faults during extrusion and exhumation of the palaeo-channel. Pervasive shearing throughout the channel and extruded crustal block, although strain is predicted to be concentrated along its boundaries due to the flow geometry and deformation history. Inverted and right-way-up metamorphic sequences at the base and top of the extruding channel, respectively.

Modelling of the channel flow predicts the following tectonic consequences. (1)

(2)

(3)

(4) (5)

The incubation period necessary for midcrustal temperatures to rise, thereby increasing the melt content for commencement of channel flow, is typically between 10 and 20 million years from the time of onset of crustal thickening. This incubation period is judged necessary to increase the mid-crustal temperature sufficiently to produce the low viscosity necessary for initiation of channel flow. Melts (leucosomes) coeval with ductile channel flow must be younger than shortening structures in overlying rocks (upper crust). When active, the channel is predicted to be 10-20 km thick (Royden et al. 1997; Clark & Royden 2000; Beaumont et al. 2004; Jamieson et al. 2004, 2006). There is more lateral transport of material in the channel than vertical. Pre-existing structures cannot be traced through the channel (from the upper crust, through the channel and into channel footwall rocks). This has direct consequences on correlation possibilities between rock units and structures from the upper crust to the lower crust.

These conditions and consequences are reviewed for the Himalayan belt, and compared with available field and geochronological data. Viscosity

A small percentage of partial melt significantly reduces the effective viscosity of rocks (Rosenberg & Handy 2005, and references therein). The GHS includes a significant percentage of migmatites and synorogenic leucogranites, while evidence for partial melting is absent in both the Lesser Himalayan sequence and the Tethyan sedimentary sequence. At the time of protracted peak temperature metamorphism (up to granulite facies) and melt generation, the GHS was therefore weaker than the overlying and underlying rocks by at least one order of magnitude (Beaumont et al. 2004, 2006; Hollister & Grujic 2006; Medvedev & Beaumont 2006).

7

The Lesser Himalayan sequence consists of a thick package of metasediments that were deformed under greenschist facies or lower conditions (< c. 300~ Although these temperatures allow for ductile flow of quartz-dominated rocks, the expected viscosities are higher than for rocks with partial melt (see Medvedev & Beaumont 2006). The Tethyan sedimentary sequence is generally unmetamorphosed and only experienced greenschist-facies metamorphism in a narrow zone at its base (Garzanti et al. 1994; Godin 2003), although contact metamorphic aureoles have been reported associated with young granites emplaced in the Tethyan sedimentary sequence of southern Tibet (Lee et al. 2000, 2006). Plateau formation

The Himalayan orogen is genetically linked to growth of the Tibetan Plateau (Hodges 2000; Yin & Harrison 2000). Various palaeo-elevation data suggest that the southern Tibetan Plateau has existed since at least the mid-Miocene (Blisniuk et al. 2001; Rowley et al. 2001; Williams et al. 2001; Spicer et al. 2003), attaining high elevations similar to the present day by 18-12 Ma, and possibly by 35 Ma (Rowley & Currie 2006). Geochronological and structural data suggest that east-west extension in the Tibetan Plateau which is believed to be linked to crustal overthickening - was well underway by 14 Ma (Coleman & Hodges 1995; Williams et al. 2001). Contrasts in crustal thicknesses (between the Indian foreland and the Tibetan Plateau) that produced the necessary gravitational potential energy for channel flow (Bird 1991) may have existed at least since the Miocene, and perhaps earlier. Creation of high topography by 'inflational' thickening is a potential consequence of crustal flow (Royden 1996; Burchfiel 2004). In the Longmen Shan belt of eastern Tibet, Clark & Royden (2000) and Clark et al. (2005) suggest that topography is generated when lower crustal flow buttresses against cold, stronger crust (e.g. Sichuan basin). The resistance to lateral flow inflates the lower weak crustal zone, and supports a high topography above, and possibly generates ramping up (extrusion) and eventual exhumation of lower crust. This process could partly address concerns about the 'support' of the plateau's high elevation, if Tibet is underlain by a weak, low-viscosity mid-crust (S. Lamb, pers. comm. 2004). A similar lithospheric strength contrast to the Longmen Shah could exist on the southern edge of the Himalaya, where the southflowing weak mid-crust buttresses against cold, strong Indian lithosphere, favouring extrusion and 'inflational' support of the southern Tibetan Plateau (e.g. Hodges et al. 2001).

8 Coeval channel-bounding

L. GODIN ETAL. structures

The MCT and STD zones include multiple fault strands that operated at different times and under different mechanical conditions (ductile to brittle). The broadly coeval activity of the MCT and the STD over extended geological time (from c. 25 Ma to 5 Ma) is documented by various sets of geochronological data. Figure 4 presents a compilation of interpreted age(s) of motion on the various strands of the MCT and STD, along the length of the Himalayan belt as taken directly from the available literature. We refer to these various strands as lower MCT (MCTI) and upper MCT (MCTu), and lower STD (STD1) and upper STD (STDu) to avoid confusion with past terminology. A major limitation to compiling such diverse data (Table 1) is the range of different approaches utilized by different authors to constrain either a maximum or minimum age of motion, on either the upper or lower strand of each fault system (from indirect geochronological tools - monazite crystal ages or peak metamorphic ages - to field relationships, e.g. preor post-kinematic intrusions). We emphasize that no attempt has been made in this compilation to critically assess the validity of different approaches taken by different authors in different areas. The compiled data indicate that the MCTu and STD~ were mostly active between 25-14 Ma and 24-12 Ma, respectively (Fig. 4). The activity along the higher STDu apparently started later and lasted longer (c. 19 Ma to 14 Ma, and perhaps is still active today, e.g. Hurtado et al. 2001) than along the more ductile lower STD,. The structurally lowest MCT, appears as the youngest structure (c. 15 Ma to 0.7 Ma). Combined, the available data indicate simultaneous or overlapping periods of thrust- and normal-sense ductile shearing between c. 24 Ma and 12 Ma. However, with time, the position of the active faults moved towards upper and lower structural levels, and became more diachronous and possibly less dynamically linked (e.g. Godin et al. 2006). Since the early recognition of the STD, the coeval activity on the two bounding (and innermost) shear zones (MCTu and STD0, and its implication for exhumation of the metamorphic core of the Himalaya, has been suggested (Burchfiel & Royden, 1985; Hubbard & Harrison 1989; Searle & Rex, 1989; Burchfiel et al. 1992; Hodges et al. 1992, 1996; Grujic et al. 1996; Grasemann et al. 1999); however, the proposed driving forces and kinematic details vary between authors. The late Miocene to recent activity along the MCT1 overlaps with activity along the two insequence external thrust faults, the Main Boundary thrust (MBT) and Main Frontal thrust (MFT), and may represent on-going exhumation of the

modern cryptic (hypothetical) channel (Hodges et al. 2004). In the context of proposed exhumation by combined channel flow and extrusion, a corresponding active zone of normal faulting at a higher structural level is required. The data are scarce but there are indications of neotectonic faulting along the northern boundary of the GHS (Hodges et al. 2001, 2004; Hurtado et al. 2001; Wiesmayr et al. 2002). The younging of structures away from the core of the orogen may suggest progressive widening of the channel as it passes from the channel flow to extrusion mode of exhumation (Searle & Godin 2003; Searle et al. 2003, 2006). Kinematic

inversions

Studies have shown that deformation along the STD is distributed in the adjacent footwall and/or hanging wall for up to 3 - 4 km, rather than being restricted to a single fault plane. Most of these studies indicate an overprint of top-to-the-north (normal sense) shearing on an older top-to-thesouth thrusting within the STD system (Burg et al. 1984; B r u n e t al. 1985; Ktindig 1989; Burchfiel et al. 1992; Vannay & Hodges 1996; Carosi et al. 1998; Godin et al. 1999a, 2001; Grujic et al. 2002; Wiesmayr & Grasemann 2002). Based on field evidence for a reversal in shear sense during motion along the STD, it seems likely that the return flow of the metamorphic core (relative to the underthrusting Indian plate) developed late in the channel flow history. The kinematic history of the STD is further complicated by overprinting top-to-the-south shearing (e.g. Godin et al. 1999a; Godin 2003). Some of this late stage overprinting may relate to north-dipping thrust faults in the Tethyan sedimentary sequence, between the suture zone and the STD (e.g. Ratschbacher et al. 1994). Geodynamic modelling (Beaumont et al. 2004) supports this possibility if the weak channel overburden fails and glides towards the foreland causing relative thrusting along the upper boundary of the channel. I n t e r n a l d e f o r m a t i o n within the c h a n n e l

On regional cross-sections and maps, the MCT and STD are often depicted as sharp boundaries; however, both are broad ductile shear zones. Although most field- and laboratory-based investigations agree that there is a broad zone of deformation adjacent to the MCT and STD, pervasively distributed ductile shear throughout the GHS is also documented (e.g. Jain & Manickavasagam 1993; Grujic et al. 1996; Grasemann et al. 1999; Jessup et al. 2006). The kinematics of deformation consistently indicate top-to-the-south shearing in the Lesser Himalayan sequence and in most of the

INTRODUCTION

9

~ 9 +5 o ~ " ~

~ o

~

~

9

~

2e~500 mW m -2 (Shanker et al. 1976), later interpreted as caused by a Quatemary intrusion at the c. 7 km depth of local seismicity (Gupta et al. 1983). Better known because collected in the framework of a large international project are later Sino-French heat-flow measurements in two lakes in southernmost Tibet (close to Kangmar Dome, Fig. 1) of 91 and 146 mW m -2 (Francheteau et al. 1984). The large variation over < 25 km between the two lakes

CRUSTAL FLOW IN TIBET requires an anomalous shallow heat source, presumably magmatic, intruded within the past 1 million years at less than c. 10 km depth beneath Yamdrok Tso, the northern of the two lakes that extends from 15-50 km south of the Indus-Yarlung suture (Jaupart et al. 1985). Additional studies of the 'Himalayan geothermal belt' have documented at least 600 geothermal systems in the Lhasa terrane and northern Tethyan Himalaya, mostly concentrated east of 84~ (Hochstein & Regenauer-Lieb 1998). After correction for convective heat flow, southern Tibet has a regional heat flow of c. 82 mW m -2 (Wang 2001), equivalent to 'characteristic Basin & Range' heat flow of Lachenbruch & Sass (1977). Regional heat flow in northern Tibet (Qiangtang terrane) is estimated, based on only two measurements, at c. 45 m W m -2 (Wang 2001), intermediate between the 'Sierra Nevada' and 'Stable reference crust' geotherms of Lachenbruch & Sass (1977). Standard conductive geotherms predict Moho temperatures of c. 700~ (i.e. in excess of the pelitic wet solidus (Thompson & Connolly 1995) even in northern Tibet (Wang 2001), due to the > 60 km crustal thickness. Analogous geotherms predict such extremely high temperatures beneath southern Tibet that the only reasonable scenario is a more isothermal crust than standard, maintained by underthrusting of a relatively cold Indian crust (Wang 2001).

Seismicity cut-off at depth locates b r i t t l e - d u c t i l e and ductile-brittle transitions Crustal seismicity, both from teleseismic observations (Molnar & Lyon-Caen 1989; Zhao & Helmberger 1991) and local network studies (PASSCAL: Randall et al. 1995; IN-3: Langin et al. 2003), is effectively limited to the upper crust. Both teleseismic studies cited, which necessarily focus on larger earthquakes, place 95% of relocated earthquakes above 18 km depth, and both local studies place 95% of located earthquakes above 2 0 k m depth. In particular the Langin et al. (2003) study (IN-3) located 267 earthquakes from the Indus-Yarlung suture to the Jinsha River suture, and across 6 ~ of longitude over a nine-month period, thereby providing a comprehensive snapshot of crustal seismicity in east-central Tibet. The downward cut-off of crustal seismicity is widely accepted as marking the brittle-ductile transition at a temperature of 250-450~ (Meissner & Strehlau 1982; Chen & Molnar 1983). By itself, the existence of a ductile lower crust does not demonstrate crust-mantle decoupling, because the mantle could in some circumstances be even weaker than the overlying lower crust (Jackson 2003; Afonso & Ranalli 2004). However, in contrast to the absence of earthquakes beneath the Tibetan Plateau with well-constrained focal depths of 30-65 km, it is

45

well-documented that earthquakes occur beneath southern Tibet at depths of 7 0 - 1 1 0 k m , some clearly below the Moho. These sub-Moho earthquakes are known from both teleseismic recordings (Molnar & Chen 1983) and temporary local network recordings (PASSCAL: Zhu & Helmberger 1996). Chert & Yang (2004) review evidence for eight subcrustal earthquakes ( 9 0 - 1 1 0 k m depth) in western Tibet (between the IndusYarlung suture and the Altyn Tagh-Karakax fault, are all clearly distinct from the Hindu Kush intra-continental subduction intermediate-depth seismic belt), and for an additional five presumed snb-Moho earthquakes (80-90 km depth) beneath the Tethyan Himalaya and southern Lhasa terrane (all east of 86~ and south of 30~ hence south of or close to the Indian mantle suture inferred from gravity modelling by Jin et al. 1996). A further two earthquakes at 70 km depth in the Himalaya/southeastern Tibet may be in the lower crust or the upper mantle. Some have suggested that all earthquakes in southern Tibet lie above the Moho (Jackson 2003; Mitra et al. 2005), pointing to evidence for crustal thicknesses as great as 85-90 km near Lhasa (Yuan et al. 1997; Mitra et al. 2005), but in doing so ignore significant west-east variations in crustal structure (SF-2: Zhang & Klemperer 2005) including equivalent evidence for thinner crust (70-80 km) closer to the location of the deep earthquakes (IN-2: Yuan et al. 1997; Chen & Yang 2004). Hence it seems probable that these deep earthquakes of magnitude 4.3-6.0 are evidence for sub-Moho brittle behaviour in southern Tibet (Chert & Yang 2004), and presumably require temperatures no more than 600-800~ (Chen & Molnar 1983). The evidence for subcrustal earthquakes is even clearer for the eastern Nepal segment of the High and Tethyan Himalaya, where over 100 small earthquakes (magnitude 10 km layer of partial melt, or a c. 1 km layer containing > 1 0 % saline fluid. The seismic constraints suggest an interpretation combining both likely causes of enhanced conductivity, with c. 200 m of c. 10% saline fluid above > 10 km of c. 10% partial melt (Li et al. 2003). Gaillard et al. (2004) experimentally confirm that hydrous granitic melts can have conductivities as high as those observed beneath the Yangbajain graben, but

point out that the presence of crystallizing leucogranites demands some local accumulation of hydrous fluids upon magma solidification, while the occurrence of a large amount of fluids in the middle crust requires that molten granite be present immediately beneath levels that are rich in fluids. Thus a combination of both magmatic and aqueous fluids is required by the joint interpretation of the seismic and MT data and our understanding of the geological evolution of crust-derived melts. The high conductivity beneath the Yangbajain graben, southern Lhasa terrane, collocated with the seismic bright spots, is therefore fully consistent with the Nelson et al. (1996) model in which molten analogues of the Miocene High Himalayan leucogranites cause the anomalous conductivity and reflectivity. It is important to note that even though the highest conductivities in southern Tibet are observed in an active hydrothermal area within the Yangbajain graben, abnormal conductance extends outside the graben into unrifted crust: profile in-200 was recorded 8 0 k i n (one crustal thickness) NW and SE beyond the limits of the graben, but shows conductivities along its entire length that are anomalous by global standards (Chen et al. 1996), as does in-500, also acquired outside the graben system (Wei et al. 2001) (Fig. 4). In contrast, the seismic bright spots are

CRUSTAL FLOW IN TIBET

51

Fig. 5. Evidence for crustal seismic anisotropy. (a) Averaged fast directions and delay times based on SKS measurements, from Huang et al. (2000). Result at AMDO from McNamara et aL (1994). BNS, Banggong-Nujiang suture; KIF, Karakoram-Jiali fault system. Very rapid change in splitting over 60 km) means that the pressure dependence of dihedral angle must be considered. Though water forms dihedral angles >60 ~ with typical crustal and mantle minerals in the lower crust and uppermost mantle of normal lithospheres (crust c. 40 km thick), at pressure (p) > 1 GPa and temperature (T) > 800~ the dihedral angle is < 6 0 ~ for water in plagioclase (Yoshino et al. 2002) and quartz (Holness 1993) aggregates, and for water in olivine at p > 1.5 GPa and T > 1000~ (Mibe et al. 1999). These are pT conditions presumably relevant to northern Tibet. Melts are also likely to form interconnected conductive pathways in the upper mantle, as dihedral angles for melt with olivine are low (Waft & Bulau 1979). The situation is less clear in the lower crust, as measured melt-pyroxene dihedral angles (Toramaru & Fujii 1986) and water-pyroxene dihedral angles (Watson & Lupulescu 1993) are > 6 0 ~ Thus the high observed conductivities imply fluids are present throughout the lower crust/uppermost mantle of northern Tibet, but do not necessarily require partial melt.

Seismic anisotropy: evidence for past or present strain of mantle and crust

In contrast to low seismic velocities and high electrical conductivities that show modern crustal weakness without demonstrating active deformation, the presence of seismic anisotropy requires flow (or possibly only recrystallization while under a non-uniform stress system) but cannot demonstrate that the flow is ongoing as opposed to a fossil record. In Tibet, anisotropy models have been presented at three scales: (1) plateau-wide observations of crustal anisotropy from surface waves (Shapiro et al. 2004); (2) regional variations in single-layer estimates of splitting of SKS (and related phases) often largely attributed to mantle deformation (e.g. SF-1, -3: Him et al. 1995; PASSCAL: McNamara et al. 1994; IN-3: Huang et al. 2000); and (3) multilayer anisotropic receiver-function inversions beneath individual stations for detailed crustal anisotropy (SF-4: Vergne et al. 2003; IN-3: Ozacar & Zandt 2004; PASSCAL: Sherrington et al. 2004). The most ubiquitous measurements of anisotropy are those of split SKS and related teleseismic phases recorded at individual stations, most routinely presented in terms of a single delay time and orientation (e.g. Silver 1996). Such anisotropy is frequently interpreted as arising in the lithospheric

mantle, because abundant olivine in the mantle allows easy creation of strain-induced lattice preferred orientation, hence significant velocity anisotropy. Most interpretations assume a single anisotropic layer, and cannot resolve delay times 5 s) of SKS waves. In Tibet most stations in the Tethyan Himalaya and the southern Lhasa terrane (PASSCAL: McNamara et al. 1994; SF-3: Him et al. 1995; IN-2: Sandvol et al. 1997) show modest or negligible splitting (0-1 s), representing either small or complexly distributed anisotropy. A similar lack of splitting (transverse isotropy) at stations on the Indian craton has been used to bolster the argument that subducting Indian lithosphere underlies southern Tibet (Chen & Ozalaybey 1998). In contrast, stations in the Qiangtang and Songpan-Ganzi terranes show large split times (1-2 s), with orientations typically 090 ___ 30 ~ (PASSCAL: McNamara et al. 1994; SF-4: Herquel et al. 1995; IN-3: Huang et al. 2000). This splitting pattern beneath northern Tibet has been interpreted as representing either asthenospheric flow enabling eastward extrusion of rigid blocks of Tibetan lithosphere (Lav6 et al. 1996) or finite strain in Tibetan lithospheric mantle (Davis et al. 1997) deforming by north-south shortening or east-west extension. Him et al. (1995; profile SF-3) argued that the same splitting orientation was present throughout the Lhasa terrane, albeit with smaller split times. A remarkable observation from the densely instrumented INDEPTH-3 transect is that the transition from negligible to significant splitting occurs over a distance of only 15 km across the Jiali fault, 40 km south of the Banggong-Nujiang suture, requiring that crustal anisotropy is at least locally substantial and distinct between the adjacent stations (IN-3: Huang et al. 2000) (Fig. 5a). Where observed splitting parameters are more spatially coherent, further north and south along the INDEPTH-3 transect, it is harder to demonstrate the depth of anisotropy, but arguments based on seismic resolution (the Fresnel zone), suggest the anisotropy largely resides in the lower crust and upper mantle (Huang et al. 2000). The crustal component of Tibetan anisotropy has been further elucidated by anisotropic receiver-function analysis that suggests dipping and horizontal Vp and Vs anisotropy in multiple layers 2 - 2 5 km thick, that may be coherent across a few tens of kilometres (Ozacar & Zandt 2004) but are far harder to correlate across inter-station distances >100 km (Sherrington et al. 2004). Although individual interpretations are non-unique (significantly different structures were interpreted beneath the same station by Frederiksen et al. (2003), Vergne et al. (2003), and Sherrington et al. (2004)), it seems

CRUSTAL FLOW IN TIBET clear that there is considerable and complex anisotropy in the crust. In part these complex anisotropies must represent a time-integrated strain history: anisotropic layers in the brittle upper crust cannot be ductiMy deforming at present, so their anisotropy may represent either open cracks or old metamorphic fabrics, whereas complex patterns of mid- and lower-crustal anisotropy may represent active deformation superimposed on older strain fabrics (PASSCAL: Sherrington et al. 2004) (Fig. 5b). The clearest relationship yet seen between active deformation and anisotropy is the observation of a layer with 20% anisotropy aligned in the direction of plate motion at the drcollement between subducting India and the Lesser Himalaya (HIMNT: Schulte-Pelkum et al. 2005), and this study shows the clear potential for very dense passive seismometer arrays to detect anisotropic flow channels elsewhere in Tibet, though caution is always required to discriminate the effects of dipping layers from those of anisotropy. The large-scale behaviour of the crest may be better captured by plateau-wide inversions of surface wave anisotropy. Shapiro et aL (2004) averaged surface waves propagating through Tibetan crust on a 1~ grid and showed that seismic velocity in the vertical direction (Rayleigh mode) is slower than in the horizontal direction (Love mode) (note this study did not resolve azimuthal variations in the horizontal wave-speed) (Fig. 5c). This observation is most simply interpreted as due to near-horizontal reorientation of anisotropic minerals - principally micas - during thinning of the anisotropic layer, because pure-shear thickening tends to produce the opposite sense anisotropy to that observed. Depending on the assumed thickness of the anisotropic layer, a vertical flattening strain of 20% (if the entire middle and lower crust below 20 km depth is deforming) to 40% (if only a 25km-thick middle crust deforms) is required to produce the observed radial anisotropy in a rock containing 30% mica (Shapiro et al. 2004). (Note that such thinning may be less than the flux of Indian crust into/beneath Tibet, so may not imply net crustal thinning; cf. Kapp & Guynn 2004.) Few rocks contain as much as 30% mica, and less micaceous rocks would require larger strains to develop the same anisotropy. Although this observed anisotropy could in principle represent effects of deformation before the Himalayan collision and/or younger events, it is clearly present beneath both the Lhasa and Qiangtang terranes, and arguably also beneath the Tethyan Himalaya and Songpan-Ganzi terrane (Shapiro et al. 2004). As with the reflective layering of the lower crust observed both north and south of the Banggong-Nujiang suture (discussed below) (IN-3: Ross et al. 2004), the simplest interpretation is that both the radial anisotropy and the crustal

53

reflectivity represent an ongoing flattening flow in a 25-50 kill thick mid-to-lower crust.

Seismic reflectivity: evidence for fluids and flow

INDEPTH has recorded c. 350 km of multifold seismic reflection data (Fig. 1) that were used to interpret the structural development of the Himalayan thrust belt (IN-l: Zhao et al. 1993; IN-2: Hauck et al. 1998), the Indus-Yarlung suture zone (IN-2: Makovsky et al. 1999) and the southern Lhasa terrane (IN-2: Alsdorf et al. 1998a). Observations particularly germane to crustal rheology are unusually high-amplitude reflections ('bright spots') that may represent magmatic (Brown et al. 1996) or aqueous (Makovsky & Klemperer 1999) fluids, and analyses of crustal reflectivity (Ross et al. 2004) (Fig. 6). The bright-spot observations on the INDEPTH-2 profiles, restricted by lack of data elsewhere to the Yangbajain graben and geothermal province of the southern Lhasa terrane, are: (1) unusually strong amplitudes recorded from a depth of c. 15 km below the surface, which require unusual lithologies or fluids, possiblyin layers thinner than the seismic wavelength (c. 10 m) that act to enhance the amplitudes (Brown et al. 1996) (Fig. 6a); (2) strong P-toS-wave seismic conversions at these reflectors, which require the presence of fluids (Makovsky & Klemperer 1999); (3) a variation of amplitude with incidence angle of these reflections that seems to indicate a fluid that is dominantly aqueous in >15% porosity (Makovsky & Klemperer 1999), though this last measurement is one with large uncertainties. Incontrovertibly, the crust below the Yangbajain graben contains magma and/or free aqueous fluids. If magma is present, the temperature must exceed the minimum granitic melt temperature at 15 km depth (c. 650~ e.g. Thompson & Connolly 1995) given the large percentage volume of melt indicated, unless we have fortuitously imaged very recent intrusions prior to their cooling and freezing. If instead the crust contains aqueous fluid, it does so in volumes too large to be easily maintained over geological time periods without replenishment (Bailey 1990; Warner 1990), thereby implying a fluid source at depth, possibly subducting India. Even small percentages of free water in the crust are likely to trigger partial melting at temperatures above the minimum melt of granite. Whether water or melt is present, the crust is likely to be weakened. These bright spots were interpreted by Nelson et al. (1996) and Wu et al. (1998) as melts of Indian and Asian crust that will in the future be exposed at the surface in a return channel flow as

54

S.L. KLEMPERER

~

~

b, ~ ~ ~.~

~

.,.~

~ix

~N

~c-,l

i~

~i ~

~

"~

~

>

~.~ ; > ~

~ 9~

~ ~

~ ' ~

~ ~ ~

~~

~o.-~

~'~

~

~

~ ~_~ 9

~

9 ~C~

~

~

.m

' "~ ~ ' ,=

o

- ~ o

~~.~

CRUSTAL FLOW IN TIBET High Himalayan leucogranites. This interpretation of leucogranitic melt is permitted but not required by the seismic and MT observations. If a return channel flow exists bounded below by the MHT and above by the STD, then crystallized analogues of the bright spots may be expected to be present in the subsurface beneath the Tethyan Himalaya. It has been suggested that bright reflections with a characteristic upward concavity above the MHT (Alsdorf et al. 1998b) may represent such intrusions (L. D. Brown, pers. comm. 2004) analogous to saucer-shaped sills now routinely mapped on 3D seismic reflection data over volcanic rifted margins (e.g. Thomson 2004). However, modelling suggests that these saucer shapes develop only at shallow depths, when their length exceeds the depth by a factor of two to three (requiring 3 0 45 km long sills for the intrusion depth of 15 km suggested by INDEPTH data) (Malthe-SCrenssen et al. 2004); and the apparent curvature on INDEPTH profiles may be an artifact of 2D seismic migration of the INDEPTH data. Pronounced 'lower-crustal' reflectivity seen elsewhere on Earth is commonly thought to result from tectonically imposed layering, including shear zones (e.g. Green et al. 1990) and/or multiple horizontal planar mafic intrusions ('sills') (e.g. Warner 1990). The characteristic reflectivity is bounded below by the Moho, and in many tectonically active locations dies out upwards at about the brittle-ductile transition (e.g. Klemperer 1987; Holbrook et al. 1991). Ross et aL (2004) show that in the northern Lhasa and southern Qiangtang terranes (IN-3, Fig. 1), strongly reflective crust exists from 25 km to the Moho at c. 65 km depth (Fig. 6a & c). Beneath the Yangbajain graben no lower-crustal reflectivity is observed (except on the northernmost profile); however, amplitude analysis of these data shows that reflected energy is a factor of three to ten below comparable values for INDEPTH-HI, consistent with a hypothesis that the observed high reflectivity of the bright spots, and observed high attenuation of the upper crust (Xie et al. 2004) due to hydrothermal activity there, prevents observations of a lower crust that is in fact reflective (Ross et al. 2004). The lack of high velocities in most of the crust (IN-2: Kola-Ojo & Meissner 2001, IN-3: Zhao et al. 2001) is an argument against widespread mafic intrusions as an explanation of the observed reflectivity, and there is no obvious magmatic event recorded at the surface in either the Qiangtang or the Lhasa terrane to which mafic intrusions may be linked. The most plausible interpretation is that the reflective zone represents a region of active ductile deformation and flattening foliation, such as would be produced in a region undergoing extensional flow, or having undergone such flow as the most recent

55

tectonothermal event (Ross et al. 2004). This reflective zone demonstrably occupies the entire middle and lower crust in the Qiangtang and northern Lhasa terranes in the few locations where profiling has been undertaken; it may occupy the entire middle and lower crust in the southern Lhasa terrane, though obscured by shallow attenuating bodies; and it is demonstrably limited to the upper crust in the Tethyan Himalaya (Ross et al. 2004). Reflectivity is particularly pronounced below the STD and above the MHT, even south of the North Himalayan gneiss domes (Fig. 6b) (IN-l: Makovsky et al. 1996), hence south of MT or geothermal evidence for young intrusive activity. Cooled felsic intrusions would not produce substantial reflectivity within the GHS, and the moderate seismic velocity within the reflective zone _800-900~ at depths of 3 0 - 5 0 k i n before reacting with magma at 1300~ in the lower crust and entrainment in a 3 Ma eruption. Initial estimates of uppermost mantle temperature from Pn tomography were 840-1170~ lower beneath the Lhasa terrane, higher beneath the Qiangtang (McNamara et al. 1997). However, the lateral temperature difference is much better determined than the absolute temperatures. Uncertainties in experimental elastic and anelastic parameters and in seismic velocities do not allow absolute constraints from seismic velocities any tighter than _+ 150~ (Goes & van der Lee 2002), and the Pn velocities are also consistent with temperatures in the seismogenic upper mantle beneath the Lhasa terrane of 10 Ma (e.g. Lachenbruch & Sass 1977) so that the south Tibetan mid-crust cannot be substantially cooler now than it was during melt formation at 7 Ma. Finally one may mention temperature estimates based on two-dimensional finite-element methods, that seek to model effects of the geometry of India underthrusting from the south and frictional heating on faults; convergence rates and erosion rates at the southern margin of Tibet; radiogenic heat production in the crust, and mantle heat flow, etc., either in steady-state kinematic models (Henry et al. 1997; Huerta et al. 1998) or in coupled thermal-mechanical models (Beaumont et al. 2004; Jamieson et al. 2004). These models are of most value in southern Tibet where constrained by geometries indicated by the INDEPTH and other seismic transects. The Henry et al. (1997) model predicts crustal temperatures south of the Indus-Yarlung suture, where they suggest that temperatures of 550-850~ may be attained in the core of the orogenic wedge, just above the MHT at c. 40 km depth, with lower temperatures of 500-600~ at the base of the crust at c. 80 km depth. Approximately doubling the mantle heat flow can increase these temperatures to 950~ and 750~ respectively. An extension of this model into the Lhasa terrane (Cattin et al. 2001) gives Moho temperatures of c. 550-700~ Beaumont et al. (2004) predict higher maximum crustal temperatures of 900~ and 800~ at the same locations, and suggest that Moho temperatures increase to about 900~ in the Lhasa terrane and 1000~ beneath the Qiangtang. Although predicted temperatures vary considerably with assumptions about poorly constrained parameters (e.g. radiogenic heat production and mantle heat flow), sensible parameter choices yield temperatures consistent with the geological and geophysical observations above.

Physical properties inferred from geophysical observations Inferred vertical strength profiles

In southern Tibet, the observation of sub-Moho earthquakes (Chen & Yang 2004; Monsalve et al. 2006) implies the crust is weaker than deeper layers. Even if the earthquakes at 80-90 km are above the Moho (Jackson 2003; Mitra et al. 2005) or in eclogitized crust now below the Moho as possibly recognized seismically (Sapin & Him 1997) or by the adakitic products of their melting (Chung et al. 2003), the relative absence of seismicity from c. 20 to 70 km demonstrates a broad channel in which deformation is accommodated by ductile, not brittle processes (Fig. 7a). This crust has lower-than-normal P-wave velocities (Him et al. 1984a; Owens & Zandt 1997; Rodgers & Schwartz 1998; Haines et al. 2003), lower-thannormal S-wave velocities (Cotte et al. 1999; Rapine et al. 2003), an intra-crustal S-wave low-velocity zone (Kind et al. 1996), and seismic reflectivity (Makovsky & Klemperer 1999) and electromagnetic conductivity (Li et al. 2003) indicative of fluids, all suggesting that the crust is unusually weak. This weakness is consistent also with the high conductive heat flow (Shanker et al. 1976; Francheteau et al. 1984), vigorous geothermal activity (Hochstein & Regenauer-Lieb 1998) and short-wavelength flexural topography (Masek et al. 1994) of southern Tibet. Evidence that magmatic intrusion (Gupta et al. 1983; Jaupart et al. 1985) and maximum concentrations of fluids (Unsworth et al. 2005) are in the upper crust, and that the intra-crustal low-velocity zone may not extend to the lower crust (Kind et al. 1996; but see also Cotte et al. 1999) suggest that the low-strength channel may be localized in the middle crust and may not include the lower-crust (Figs 3 & 7a). This region with a strength minimum in the middle crust includes the Tethyan Himalaya at least as far south as the North Himalayan gneiss domes, based on the conductivity anomalies (Unsworth et al. 2005) and measurements of high heat flow (Gupta et al. 1983; Francheteau et al. 1984). A dramatically weaker mid-crust does not extend significantly further south based on the disappearance of the highest conductivities and heat flow, the rapid decrease in seismic attenuation (Xie et al. 2004), and the southward decrease in crustal temperatures required by any reasonable thermal model (Henry et al. 1997; Huerta et al. 1998). The crustal weak zone presumably includes much of the Lhasa terrane, extending to the northern edge of colder south-Tibetan mantle that is crudely placed beneath the Banggong-Nujiang suture by Pn and Sn velocity tomography

CRUSTAL FLOW IN TIBET

(a)

relative strength

.-

Tethyan Himalaya and Lhasa Terrane south of Indian mantle suture recent intrusions or magma bodies localized in Tethyan Himalaya & Yangbajain graben

'

relative strength

(b)

l

Northern Lhasa Terrane from Indian mantle suture to l n d i a n m a n t l e front

57 (c)

relative strength

Qiangtang Terrane

)

low strength in thermal maximum around H~malayan Thrust

. L,,,,,\ mafic lower crust of Qiangtang Terrane

" " "-4 [ top of subducting " "- ./ Indian craton ~deepens

I25 s seismogenic \ eclogiUzed Indian I crust and/or j

mafic lower crust of Lhasa Terrane

Moho warm Asian mantle

warm Asian lithospheric mantle subducting Indian crust _ _ _._-'-;....... ........ -/......,,.

/

cold ~ubducting Indian mantle deepens northwards

Fig. 7. Cartoons of inferred relative strength-depth profiles for Tibet at about 90~ The competing effects of increasing temperature with depth and compositional variability (generally more mafic at depth) allow multiple weak channels in the lithosphere. (a) In the Tethyan Himalaya and southern Lhasa terrane the greatest lithospheric strength is likely in the seismogenic upper mantle. Dashed lines represent laterally variable conditions. (b) In the northern Lhasa terrane north of the Indian mantle suture it is uncertain whether the upper crust or upper mantle is stronger, but the weakest zone is probably at the base of the middle crust. (e) In the central and eastern Qiangtang terrane the only significantlithospheric strength is probably in the brittle upper crust; west of c. 85~ the mantle may become somewhat colder and stronger.

(McNamara et al. 1995, 1997; Tilmann et al. 2003). A better resolved measure of the change in mantle properties may be the clear change in mantle anisotropy at the Jiali fault system (Huang et al. 2000), which has recently been suggested to reach through the crust to the Moho based on receiverfunction imaging (Galv~ et al. 2002b) and crustal refraction data (Zhang & Klemperer 2005). A plausible suggestion is that this boundary represents the northern limit of Indian cratonic mantle in contact with the Moho (Beghoul et al. 1993; Jin et al. 1996; Tilmann et al. 2003), either the leading edge of Indian lithosphere, or its northernmost extent before it bends sharply down to the north. It is possible that a thin zone of warmer Asian mantle still overlies the Indian lithosphere

(from c. 30~ the 'mantle suture' of Jin et al. (1996) to c. 32~ their 'mantle front'), as indicated by mantle-derived helium emissions in this region (Hoke et al. 2000). The presence of a thin zone of warm Asian mantle above cold Indian mantle would not remove the strength minimum, though could potentially lead to a more complex strength profile in which uppermost (Asian) mantle forms part of the weak zone (Fig. 7b). In northern Tibet the situation is less clear-cut. There is overwhelming evidence that the crust is unusually weak by global standards because it is thick (Owens & Zandt 1997; Vergne et al. 2002), low-velocity (Rodgers & Schwartz 1998; Rapine et al. 2003), seismically attenuating (Rodgers & Schwartz 1998; Fan & Lay 2003), hot

58

S.L. KLEMPERER

(Hacker et al. 2000; Mechie et al. 2004), and contains a few per cent of fluids (Unsworth et al. 2004). However, the upper mantle of northern Tibet is also low-velocity after appropriate temperature correction (Rodgers & Schwartz 1998), is also seismically attenuating (Ni & Barazangi 1983; McNamara et al. 1995), and probably also contains fluids (Unsworth et al. 2004) at least east of c. 85~ (Barazangi & Ni 1982; Dricker & Roecker 2002) (at present we lack detailed observations of the Qiangtang mantle from broadband seismic deployments at 82-88~ Fig. 1). Some have argued for removal of the entire lithospheric mantle of northern Tibet and its replacement by asthenospheric material (Molnar et al. 1993; Willett & Beaumont 1994), perhaps due to subduction of Indian and Asian lithospheres (Kosarev et al. 1999; Kind et al. 2002; Shi et al. 2004) which would limit the region of upwelled asthenosphere to south and north respectively. However, recent estimates of depth to the lithosphere-asthenosphere boundary strongly suggest that > 100 km lithospheric mantle is present everywhere beneath Tibet (Priestley & McKenzie 2006), albeit somewhat thinner beneath the Qiangtang than beneath the Lhasa terrane (Kumar et al. 2006). The clear anisotropy observed in both crust (Shapiro et al. 2004; Sherrington et al. 2004) and mantle (McNamara et al. 1994; Huang et al. 2000) suggests both layers are deforming at the present by distributed strain. It is therefore not certain whether the mantle is weaker than the crust, or vice versa (Fig. 7c). The clear lowercrustal and Moho conversions on receiver functions across northern Tibet (Owens & Zandt 1997; Vergne et al. 2002) demonstrate a significant impedance increase at the top of the mafic lower crust and mantle, requiring rapid changes in lithology. If temperatures are in conductive equilibrium at the Moho, and fluids are present both above and below the Moho (Unsworth et al. 2004), then it seems most likely that there are increases in strength with depth at the top of a mafic lower crust and at the Moho, irrespective of whether the middle-lower crust or upper mantle possesses the larger total strength (Fig. 7c). Inferred

vertical velocity profiles

In southern Tibet, the surface is moving approximately NNE with respect to stable Eurasia (Wang et al. 2001; Chen et al. 2004). The lower lithosphere must have the same sense of motion, because the Indian plate is underthrusting Tibet at the Himalayan front, and converging with Eurasia in the same direction at the same or higher speed as GPS stations located in and north of the Tethyan Himalaya (Wang et al. 2001). Similar focal

mechanisms in the south Tibetan upper crust and upper mantle (Molnar & C h e n 1983) are consistent with identical kinematics at these different depths, whether or not they are decoupled by the relatively aseismic middle and lower crust. However, erosion is continuously removing significant amounts of material from the Himalaya (Johnson 1994; Galy & France-Lanord 2001), in a discontinuous manner highly focused at the topographic front (Wobus et al. 2003). Presuming the mountain front is not rapidly migrating northwards due to erosional retreat (Avouac 2003), the eroded material must be continuously replenished by some form of deformation, whether channel flow (Grujic et al. 2002; Searle & Szulc 2005) or thrust wedging (Grasemann & Vannay 1999), thereby advecting material southwards from Tibet to the Himalayan front (Hodges et al. 2001; Beaumont et al. 2004). Hence for at least southernmost Tibet there should be a reduced northward velocity (with respect to stable Eurasia) in the middle crust with respect to the upper crust and the upper mantle (Fig. 8). This is by definition a channel flow. The flow may be largely of Poiseuille form, driven outwards from the plateau by gravitational potential energy, but presumably also experiences a northward basal drag due to underthrusting of India (a Couette-type component - Fig. 2b) (Beaumont et al. 2004). It is at least plausible that this southward-directed flow (with respect to superjacent and subjacent layers) extends with diminishing vigour as far north as the inferred Indian mantle front at about the position of the Jiali fault. North of the Jiali fault, the eastward velocity of Tibetan crust from GPS measurements reaches as much as 50% of the total convergence between India and Asia in the eastern part of the plateau (Chen et al. 2004; Zhang et al. 2004) (no data are available for the western Qiangtang). The continuity of the strain field over hundreds of kilometres as measured on tens of stations in single transects shows that this eastward motion is not that of a relatively undeforming block bounded by rapidly moving strike-slip faults, as suggested by past extrusion models (Tapponnier et al. 1982, 2001), but instead is driven by more-distributed east-west extension in response to the north-south convergence across Tibet. Clearly northern Tibet is deforming in an essentially continuous manner, and the lower crust and upper mantle must both be deforming by flow because they are weak (Fig. 8). Strong anisotropy in both the crust (Shapiro et al. 2004) and mantle (McNamara et al. 1994; Huang et al. 2000) reflect this modern strain field. The correlation of the surface strain field (from GPS) with the mantle strain field (from SKS splitting) could simply result if the crust and mantle lithosphere are under similar boundary conditions (Holt 2000);

CRUSTAL FLOW IN TIBET

59

Fig. 8. Suggested directions of middle/lower-crustal channel flow in Tibet (open arrows) and regions of no flow (open circles). Gravitational potential energy and orographic exhumation drive a south-directed Poiseuille-type flow between subducting Indian lower lithosphere and brittle upper crust of Tethyan Himalaya and southern Lhasa terrane. Southern boundary of this flow province is the Himalayan topographic front; northern boundary may be the Jiali fault in the east, and the Karakoram-Jiali system or Banggong-Nujiang suture in the west. North-south compression and east-west extension of northern Tibet drives a mixed Poiseuille/Couette-type flow eastwards beneath the Qiangtang terrane, which bifurcates north and south of the rigid Sichuan basin. Northern boundary of this flow province is the KarakaxAltyn Tagh fault system bounding the rigid Tarim basin, and the Qaidam border fault south of the Qaidam basin, which both mark large and abrupt changes in crustal thickness. Arrows cross-cutting faults/sutures represent interpreted flow at depth beneath upper-crustal faults, including Kunlun fault, Jinsha River suture, and Banggong-Nujiang suture. See caption to Figure 1 for abbreviations.

all levels in the lithosphere may be flowing in the same direction, but with different velocities and strain rates. Tikoff et al. (2004) have argued that global evidence for compatible upper-crustal and upper-mantle deformation implies at least partial mechanical coupling between these layers; they suggest that bottom-driven systems must be the general rule in orogens since tectonic movements ultimately result from deep-mantle processes. If applied to northern Tibet this argument would seem to ignore the profound weakness of the lower lithosphere that can hardly drive deformation in the strong upper crust. Flesch et al. (2005) provide a stronger argument for crust-mantle coupling in Tibet: they use dynamic modelling of Eurasian deformation to show that plate-tectonic boundary conditions and topographically induced body forces are roughly equal; but because the body forces are applied to the crust, they can only contribute to mantle deformation if crust-mantle coupling is strong. Flesch et al. (2005) go on to argue that the

vertically coherent deformation implied by equivalence of GPS- and SKS-derived strain fields implies that mantle deformation in Tibet is driven by uppercrustal dynamics through a strong crust-mantle coupling that is 'incompatible with the "jellysandwich" rheology' and 'precludes large-scale lower crustal flow or mantle delamination' as described by Willett & Beaumont (1994), Clark & Royden (2000) and Beaumont et al. (2004). This seems too strong a conclusion and ignores the likelihood that the middle crust is the weakest part of the lithosphere throughout much of Tibet (Fig. 7b & c). Comparison of Rayleigh and Love wave velocities implies a strong flattening anisotropy throughout Tibet that has apparently thinned the middle crust more than the upper crust (Shapiro et al. 2004), while seismic-velocity measurements (above) suggest there is at least some strength increase from the low-velocity middle to the higher-velocity lower crust, and from the mafic lower crust to the mantle. The probable strength

60

S.L. KLEMPERER

increases at depth in the lithosphere, and the likelihood that the observed anisotropy represents active deformation, point to a channel flow with a Couette-type component driven by the strong upper crust, in addition to the Poiseuille flow predicted by Clark & Royden (2000). There could even be total decoupling of the upper crust and upper mantle by a very weak mid-crustal layer (Royden et al. 1997; Shen et al. 2001), with the strong upper crust deforming under the influence of both lateral boundary forces and body forces, and the stronger upper mantle deforming under the influence of lateral boundary forces and basal drag. The northern boundary of this proposed flow regime in northern Tibet with mixed PoiseuilleCouette characteristics is presumably the KarakaxAltyn Tagh-Qaidam border fault, which marks an abrupt change into thinner crust (Zhu & Helmberger 1998; Kao et al. 2001; Vergne et al. 2002; Wittlinger et al. 2004) overlying a seismically faster, colder and stronger mantle (Wittlinger et al. 1998; Zhao et al. 2006). In mirror image of the Himalayan thrust front, the older, colder Tarim and Qaidam lithosphere is underthrusting southwards beneath Tibet along these faults (e.g. Kao et al. 2001; Kind et al. 2002) but the low erosion rates at the northern margin of Tibet prevent the emergence there of a channel flow similar to that at the Himalayan front (Willett 1999). The least studied area of Tibet is the west-central Qiangtang terrane (Fig. 1). If lower-crustal material cannot escape northward across the Karakax-Altyn Tagh fault, either no channel flow is occurring here or, I speculate, flow is generally out to the east to replenish material and provide continuity of the flow interpreted beneath the eastern Qiangtang (Fig. 8). The western Qiangtang has the strongest crustal radial anisotropy of any part of Tibet in the interpretation of Shapiro et al. (2004), and the nearest SKS splitting measurements (that combine crustal and mantle anisotropy) at 80~ (Herquel 2005) and 88~ (Huang et al. 2000) are generally oriented NE to east, consistent with a uniform flow pattern across the Qiangtang terrane. The western limit of this northern Tibetan flow regime (Fig. 8) is unknown. However, high crustal attenuation consistent with at least some partial melt extending west beyond 80~ (Fan & Lay 2002) coupled with possibly higher mantle velocities beneath northwestern Tibet (Dricker & Roecker 2002; but see also McNamara et al. 1997) indicate a pronounced low-strength channel in the midlower crust at least to 80~ Hence mid-lower crustal flow may also be active in far western Tibet where north-south narrowing of the orogen is suggestive of the need for larger magnitudes of extrusion and crustal escape than further east, and where the highest average elevations on the plateau offer additional driving forces.

Evaluation models:

of existing channel-flow

southern

Tibet

For the Himalaya and southern Tibet the geophysical observations and the implied rheological properties suggest flow is presently occurring in the inevitable strength minimum in the mid-to-lower crust, and is likely south-directed to compensate for orographic exhumation (Fig. 8). This is presumably not the precise model of Nelson et al. (1996) or of Beaumont et al. (2004), since newer data cast doubt on some aspects of these models, encouraging their further refinement. For example Nelson et al. (1996) predicted that the protolith for the channel material north of the Indus-Yarlung suture would be Asian (Lhasa block) crust, because the magmas inferred from seismic bright spots at 15 km depth in the Yangbajain graben (Brown et al. 1996) are within Asian crust in all reasonable interpretations of INDEPTH reflection data (IN-2: Alsdorf et al. 1988a; Makovsky et al. 1999). However, a decade of new isotopic data makes it clear that GHS rocks exposed between the MCT and STD have Indian, not Asian affinity (e.g. Myrow et al. 2003; Harrison 2006). Either sufficiently northern parts of the channel are not yet returned to the surface or, ironically, the seismic bright spots that were a significant original motivation for the Nelson et al. (1996) model of channel flow do not participate in the exhumed channel. This need not preclude a conceptually similar channel flow at greater depth: in the Beaumont et al. (2004) and Jamieson et al. (2004) model the return channel flow is confined beneath the Indus-Yarlung suture which is modelled to extend at mid-crustal levels for >200 km north of the surface position of the suture. Even newer data show that Miocene (21-8 Ma) granitoids uplifted from 12 to 15 km depth in the Nyainqentanghla Range, immediately NW of the Yangbajain graben (Fig. 1), lack the Indian isotopic component (Kapp et al. 2005; Harrison 2006) that might be expected from a deeper, partially molten Indian crust; and some 11.5 Ma dykes south of the Indus-Yarlung suture zone show the same isotopic characteristics, i.e. represent Asian-affinity material south of the suture (King et al. 2005). Though these isotopic measurements argue against one specific formulation of the channel-flow model, they are additional evidence for crustal flow, as they require a crustal magmatic system in the Lhasa terrane throughout the Miocene (hot hence weak crust: D'Andrea et al. 2001; Kapp et al. 2005) and require southward transport of Asian material beneath the surface suture (lateral displacement with respect to upper crust: King et al. 2005) These new isotopic data might be satisfied by a geological history in which the Indian mantle

CRUSTAL FLOW IN TIBET suture did not reach north of the Nyainqentanghla Range until after 8 Ma (after the youngest magmas intruded; D'Andrea et al. 2001), and by a channel-flow model incorporating spatially variable material properties (Beaumont et al. 2006), thereby allowing more complex flow geometries and the potential for mixing of material north and south of the Indus-Yarlung suture. Most measurements, geological or geophysical, do not adequately discriminate between different competing models in southern Tibet. For example, shallowing thermal (Jaupart et al. 1985) and conductivity (Unsworth et al. 2005) anomalies at the latitude of the North Himalayan gneiss domes match predictions of channel flow in Beaumont et al. (2004), but do not require it: they could be ascribed to gneiss diapirism (Lee et al. 2004) triggered by an underthrust ramp or overburden extension whether or not a flow channel is present at depth. The concentration of seismic reflectivity above the Main Himalayan Thrust and below the STD (Fig. 6b; Makovsky et al. 1996; Ross et al. 2004) also matches a prediction of the Beaumont model, but could equally be explained as deformation within an extruding thrust wedge or a conventional ductile thrust duplex. Similarly, ductile shear strain and out-of-sequence thrusting that have been considered diagnostic of channel-flow models are equally compatible with conventional thrust behaviour (e.g. Robinson & Pearson 2006). However, the large-scale fold-and-thrust-belt geometry of the Himalaya, and the coherent fightside-up stratigraphic succession of the Greater and Lesser Himalayan Sequences (e.g. Harrison 2006; Robinson & Pearson 2006) seem inconsistent with the wholesale mixing during channel flow shown by current numerical realizations of channel-flow models (Beaumont et al. 2004; Jamieson et al. 2004) in which apparent stratigraphy does not always represent original relationships (Jamieson et al. 2006). Most flow models are created a posteriori to explain existing data, and will continue to be revised to better fit newer data. Nonetheless, one can conceive measurements that might discriminate between some specific models. Continuing coeval motion of the MCT and STD is a requirement of the Nelson et al. (1996) and Beaumont et al. (2004) channel-flow models, and also of the extruding wedge concept (Burchfiel & Royden 1985), but is not required by conventional thrust belt behaviour. It is debated whether the STD and MCT are presently inactive (Searle & Szulc 2005) or are still active today (Hurtado et al. 2001) and hence capable of acting as boundaries of a modern channel flow, ductile at depth but exhumed at the surface by brittle processes. High-resolution seismic reflection data could track the MCT and

61

STD to depth: the extruding wedge hypothesis suggests these two structures should meet at depth but channel-flow models predict that they do not, so that a clear image of the STD merging with the MCT at depth would refute channel flow at that location. Thus far, INDEPTH reflection data suggest a reasonable uniformity of thickness of the 'channel' between the STD and MHT for 100 km from surface outcrop of the STD to the North Himalayan gneiss domes (Kangmar Dome, Fig. 1) (IN-l: Hauck et al. 1998), albeit in an area of substantial along-strike structural complexity (Wu et al. 1998), but do not unequivocally image both the STD and MHT beneath or north of the Indus-Yarlung suture (IN-2: Alsdorf et al. 1998a; Cogan et al. 1998; Makovsky et al. 1999). Channel-flow models require pervasive deformation, hence strong seismic anisotropy, within the channel, whereas the thrust-duplex model permits passive underthrusting of relatively undeformed slices though it is also consistent with ductile flow and transposition of rocks between the STD and MCT. Thus while falsification of the hypothesis of modern channel flow in southern Tibet seems clearly possible (though certainly not yet accomplished), clear proof seems more difficult. Perhaps the closest to a demonstration that is likely to be achieved in the foreseeable future, of a channel flow following the intellectual construct if not the precise details of Nelson et al. (1996), would be the observation of strong anisotropy between and parallel to STD and MCT reflections that continue from outcrop north beyond the Indus-Yarlung suture.

Evaluation of existing channel-flow models: northern and eastern Tibet

For northern and eastern Tibet, the most explicit channel-flow models yet proposed are the pure Poiseuille 2D model of Clark & Royden (2000) and the 3D Newtonian viscous model of Shen et al. (2001) that predicts a Poiseuille-dominated flow. The Clark & Royden model is a simple 2D model that provides an excellent fit to the only data it seeks to match, the low-gradient topographic slope at the NE and SE plateau margins, but suggests unrealistically low viscosities. The Shen et al. (2001) model specifies a simple Newtonian viscous rheology that is horizontally homogeneous in the crust, but incorporates a weak zone within the deep crust of the plateau above a much stronger mantle. Because no equivalent to the Greater Himalayan Sequence crops out on the eastern margin of Tibet, there is no equivalently detailed geohistory to be matched by more complex models. The style of flow in these two models is certainly consistent

62

S.L. KLEMPERER

with that inferred from the mid-crustal radial anisotropy measured from surface waves (Shapiro et al. 2004). Given the large body of evidence for the extreme weakness of the middle and lower crust and upper mantle north of the B a n g g o n g - N u j i a n g suture, the greater part of the strength of the lithosphere surely resides in the upper crust. The upper crust is deforming in a nearly continuous strain field (at least when measured at scales of c. 100 km; Chen et al. 2004; Zhang et al. 2004), and this eastward expulsion of Tibetan upper crust between India and Tarim/Qaidam suggests an additional top-driven Couette component to the mid-lower crustal flow. Conversely, the surface velocity field is far more continuous than the midcrustal flow directions inferred by Clark & Royden (2000) from the topographic slopes of the eastern margin of Tibet (NE-directed north of the Sichuan Basin and SE-directed to the south with no or low flow predicted eastwards beneath the Sichuan Basin; Fig. 8), implying at least some decoupling of flow at different depths in the lithosphere. Rather than the upper crust completely driving upper-mantle deformation (Flesch et al. 2005), to some extent the upper mantle may be deforming independently, though fortuitously in a seemingly vertically coherent fashion because of the imposition of near-identical boundary conditions (Holt 2000).

Summary The now extensive geophysical measurements in Tibet, catalogued above, imply temperatures and rheologies that require crustal flow. Crustal seismic velocities are so low and crustal attenuation so high that Tibetan crust is surely fluid-rich, a conclusion also demanded by the MT data. Great crustal thickness requires temperatures above the pelitic wet solidus even in areas of low heat flow, as is also suggested for Tibet by recent magmatic activity, so that many of these crustal fluids are likely to be partial melts. These fluids weaken the middle and lower crust as shown by the low elastic thickness, low topographic relief and shallow cut-off to crustal seismicity on the Tibetan Plateau. These geophysical measurements represent actual conditions; hence Tibetan crust today is hot, fluid-rich and weak. Though strong crustal anisotropy and reflectivity do not require active deformation, most plausibly they are diagnostic of ongoing flow in today's weak crust. My assessment of specific modes of crustal flow should be considered more tenuous and likely to be changed in significant ways by future modelling and data acquisition. Nonetheless, Indian underthrusting in southern Tibet is most consistent with a south-directed flow of the middle crust with

respect to both the Tibetan upper crust and Indian lower crust/upper mantle - conceptually a Poiseuille-type flow. No geophysical evidence requires the very explicit formulation of the channel-flow model by Beaumont et al. (2004) and Jamieson et al. (2004) for southern Tibet, but all the geophysical data are consistent with it - inevitably, since the model was designed a p o s t e r i o r i to fit the observations. Some geological data (Robinson & Pearson 2006; Harrison 2006) disagree with specific predictions of the preferred Beaumont et al. (2004) model, but it seems probable that the model could be tuned further to match these observations also. In northern and eastern Tibet, flow patterns are still less certain, in large part because of uncertainty about the strength of the mantle lithosphere. A reasonable assessment of sometimes conflicting datasets suggests there may be a weakly coupled flow of less viscous mid-lower crust above more viscous mantle lithosphere, involving elements of both Poiseuille- and Couette-type flow. The current generation of coupled thermal-mechanical models is 2D, but will need to become 3D in order to simultaneously model the clearly different rheological stratification (Fig. 7), hence inferred different flow patterns (Fig. 8), in different parts of Tibet. Metalwhile, field geophysicists need to acquire new transects across the plateau to broaden the perspective gained by intense focus on the most logistically accessible corridor across eastern Tibet, on which most models for the development of Tibet and the Himalaya are based. I was privileged to work on the INDEPTH transect under the leadership of Doug Nelson from 1992 until his untimely death in 2002. Doug and my other INDEPTH colleagues provided my education about Tibet. I also thank those vocal participants at the Channel Flow conference who made me better informed about competing models and challenges to the INDEPTH paradigm of channel flow in southern Tibet. R. Bendick significantly improved my understanding of flow modelling through her review; additional reviews provided by M. Searle and R. Law, and comments and access to preprints from P. DeCelles, P. Kapp, J. Ni, L. Flesch, D. Grujic, C. Beaumont, M. Edwards, M. Harrison, K. Hodges, R. A. Jamieson, D. Robinson, C. Ruppel and A. Sheehan all contributed to making this a better and more comprehensive paper. The INDEPTH transect is funded by the NSF Continental Dynamics Program.

References AFONSO, J. C. & RANALLI, G. 2004. Crustal and mantle strengths in continental lithosphere: is the jelly sandwich model obsolete? Tectonophysics, 394, 221-232. ALSDORF, D. & NELSON, D. 1999. Tibetan satellite magnetic low; evidence for widespread melt in the Tibetan crust? Geology, 27, 943-946.

CRUSTAL FLOW IN TIBET ALSDORF, D., BROWN, L., NELSON, K. D., MAKOVSKY, Y., KLEMPERER,S. L. & ZHAO, W.-J. 1998a. Crustal deformation of the Lhasa terrane, Tibet plateau from INDEPTH deep seismic reflection profiles. Tectonics, 17, 501-519. ALSDORF, D., MAKOVSKY, Y., ZHAO, W.-J. er AL 1998b. INDEPTH (International Deep Profiling of Tibet and the Himalaya) multi-channel seismic reflection data: Description and availability. Journal of Geophysical Research, 103, 26,99326,999. AMMON, C. J., RANDALL,G. E. ~r ZANDT, G. 1990. On the nonuniqueness of receiver function inversions. Journal of Geophysical Research, 95, 15,30315,318. AVOUAC, J.-P. 2003. Mountain building, erosion, and the seismic cycle in the Nepal Himalaya. Advances in Geophysics, 46, 1-80. BAILEY, R. C. 1990 Trapping of aqueous fluids in the deep crust. Geophysical Research Letters, 17, 1129-1132. BARAZANGI, M. & NI, J. 1982. Velocities and propagation characteristics of Pn and Sn beneath the Himalayan arc and Tibetan plateau; possible evidence for underthrusting of Indian continental lithosphere beneath Tibet. Geology, 10, 179-185. BEAUMONT C., JAMIESON, R. A., NGUYEN, M. H. 8z ]~'[EDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06406. DOI: 10.1029/2003JB002809. BEAUMONT C., NGUYEN, M. H., JAMIESON, R. A. & ELLIS, S. 2006. Crustal flow modes in large hot orogens. In: LAW, R. D., SEARLE, M. P. 8z GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 91 - 145. BEGHOUL, N., BARAZANGI,M. & ISACKS, B. L. 1993. Lithospheric structure of Tibet and western North America; mechanisms of uplift and a comparative study. Journal of Geophysical Research, 98, 1997-2016. BENDICK, R. 2004. Topography associated with flow over an indenter in a rigid deformable channel (abstract). In: SEARLE, M. P., LAW, R. D. & GODIN, L. (convenors) Programme and abstracts for conference on: Channel Flow, Ductile Extrusion and Exhumation of Lower-Mid Crust in Continental Collision Zones, 6 - 7 December 2004, Geological Society, London. BIRD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10,27510,286. BRACE, W. F. & KOHLSTEDT, D. L. 1980. Limits on lithospheric stress imposed by laboratory experiments. Journal of Geophysical Research, 85, 6248 -6252. BRAITENBERG, C., WANG, Y., FANG, J. & HSU, H. 2003. Spatial variations of flexure parameters over the Tibet-Quinghai plateau. Earth and Planetary Science Letters, 205, 211-224.

63

BRANDON, C. 8r ROMANOWICZ, B. 1986. A "no-lid" zone in the center Chang-Thang platform of Tibet: Evidence from pure path phase velocity measurements of long period Rayleigh waves. Journal of Geophysical Research, 91, 6547-6564. BROWN, L. D., ZHAO, W., NELSON, K. D. ETAL. 1996. Bright spots, structure and magmatism in southern Tibet from INDEPTH seismic reflection profiling. Science, 274, 1688-1690. BURCHFIEL, B. C. & ROYDEN, L. H. 1985. North-south extension within the convergent Himalayan region. Geology, 13, 679-682. BUROV, E. B. & WATTS, A. B. 2006. The long-term strength of continental lithosphere: 'jelly sandwich' or 'cr~me brfil6e'. GSA Today, 16(1), 4-10. CATTIN, R., MARTELET, G., HENRY, P., AVOUAC,J. P., DIAMENT, M. & SHAKYA, T. R. 2001. Gravity anomalies, crustal structure and thermo-mechanical support of the Himalaya of central Nepal. Geoph~csical Journal International, 147, 381 - 392. CHEN, L., BOOKER, J., JONES, A., NONG, W., UNSWORTH, M., WEI, W. • TAN, H. 1996. Electromagnetic images of colliding continents: A magnetotelluric survey of the Tsangpo Suture Zone and surrounding regions. Science, 274, 1694-1696. CHEN, Q., FREYMUELLER, J., WANG, Q., YANG, Z., XU, C. & LID, J. 2004. A deforming block model for the present day tectonics of Tibet. Journal of Geophysical Research, 109, B01403. DOI: 10.1029/2002JB002151. CHEN, W.-P. & MOLNAR, P. 1983. Focal depths of intracontinental and intraplate earthquakes and their implications for the thermal and mechanical properties of the lithosphere. Journal of Geophysical Research, 88, 4183-4214. CHEN, W.-P. ~; {~ZALAYBEY, S. 1998. Correlation between seismic anisotropy and Bouguer gravity anomalies in Tibet and its implications for lithospheric structures. Geophysical Journal International, 135, 93-101. CHEN, W.-P. & YANG, Z. H. 2004. Earthquakes beneath the Himalayas and Tibet; evidence for 304, strong lithospheric mantle. Science, 1949-1952. CHRISTENSEN, N . I . 1996. Poisson's ratio and crustal seismology. Journal of Geophysical Research, 101, 3139-3156. CHR~STENSEN, N. I. & MOONEu W. D. 1995. Seismic velocity structure and composition of the continental crust; a global view. Journal of Geophysical Research, 100, 9761-9788. CrIUN, K.-Y. & YOSHII, T. 1977. Crustal structure of the Tibetan Plateau; a surface-wave study by a moving window analysis. Bulletin of the Seismological Society of America, 67, 735-750. CHUNG, S.-L., LIU, D.-Y., Jl, J.-Q. ETAL. 2003. Adakites from continental collision zones: melting of thickened lower crust beneath southern Tibet. Geology, 31, 1021-1024. CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze; building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. CLARK, M. K., SCHOENBOHM, L. M., ROYDEN, L. H., ET AL. 2004. Surface uplift, tectonics, and erosion

64

S.L. KLEMPERER

of eastern Tibet from large-scale drainage patterns. Tectonics, 23, TC1006. DOI:10.1029/2002 TC001402. CLARK, M. K., BUSH, J. W. M. & ROYDEN, L. H. 2005. Dynamic topography produced by lower crustal flow against rheological strength heterogeneities bordering the Tibetan Plateau. Geophysical Journal International, 162, 575-590. COGAN, M. J., NELSON, K. D., KIDD, W. S. F. & Wu, C. 1998. Shallow structure of the YadongGulu rift, southern Tibet, from refraction analysis of Project INDEPTH common midpoint data. Tectonics, 17, 46-61. COTTE, N., PEDERSEN, n., CAMPILLO, M., MARS, J, NI, J. F., KIND, R., SANDVOL, E. & ZHAO, W. 1999. Determination of the crustal structure in southern Tibet by dispersion and amplitude analysis of Rayleigh waves. Geophysical Journal International, 138, 809-819. D'ANDREA, J., HARRISON, T. M., GROVE, M. & L1N, D. 2001. The thermal evolution of the Nyainqentanglha Shan; evidence for a long-lived, lower crnstal magmatic system in southern Tibet. Journal of Asian Earth Sciences, 19, 12-13. DAVIS, P., ENGLAND,P. & HOUSEMAN,G. 1997. Comparison of shear wave splitting and finite strain from the India-Asia collision zone. Journal of Geophysical Research, 102, 27,511-27,522. DECELLES P. G., ROBINSON, D. M. & ZANDT, G. 2002. Implications of shortening in the Himalayan foldthrust belt for uplift of the Tibetan Plateau. Tectonics, 21(6), 1062. DOI: l 0.1029/2001TC001322. DEWEY, J. F. & BURKE, K. C. 1973. Tibetan, Variscan, and Precambrian basement reactivation: products of continental collision. Journal of Geology, 81, 683 -692. DING, L., KAPP, P., ZHONG, D. & DENG, W. 2003. Cenozoic volcanism in Tibet: evidence for a transition from oceanic to continental subduction. Journal of Petrology, 44, 1833-1865. DRICKER I. G. & ROECKER, S. W. 2002. Lateral heterogeneity in the upper mantle beneath the Tibetan plateau and its surroundings from SS-S travel time residuals. Journal of Geophysical Research, 107, 2305. DOI: 10.1029/2001JB000797. ENGLAND, P. C. & HOUSEMAN, G. A. 1986. Finite strain calculations of continental deformation; 2, Comparison with the India-Asia collision zone. Journal of Geophysical Research, 91, 3664-3676. ENGLAND, P. & MOLNAR, P. 1997. Active deformation of Asia: From kinematics to dynamics. Science, 278, 647-650. FAN, G.-W. & LAY, T. 2002. Characteristics of Lg attenuation in the Tibetan Plateau. Journal of Geophysical Research, 107, 2256. DOI: 10.1029/ 2001JB000804. FAN, G.-W. & LAY, T. 2003. Strong Lg attenuation in the Tibetan Plateau. Bulletin of the Seismological Society of America, 93, 2264-2272. FIELDING, E., ISACKS,B., BARAZANGI,M. & DUNCAN, C. C. 1994. How flat is Tibet? Geology, 22, 163-167. FLESCH, L. M., HOLT, W. E., SILVER, P. G., STEPHENSON, M., WANG, C.-Y. & CHAN, W. W.

2005. Constraining the extent of crust-mantle coupling in central Asia using GPS, geologic, and shear-wave splitting data. Earth and Planetary Science Letters, 238, 248-268. FRANCHETEAU, J., JAUPART, C., SHEN, X. J., KANG, W. H., LEE, D. L., BAI, J. C., WEI, H. P. & DENG, H. Y. 1984. High heat flow in southern Tibet. Nature, 307, 32-36. FREDERIKSEN,A. W., FOLSOM, n. • ZANDT, G. 2003. Neighbourhood inversion of teleseismic Ps conversions for anisotropy and layer dip. Geophysical Journal International, 155, 200- 212. GAILLARD, F., SCAILLET, B. 8z PICHAVANT,M. 2004. Evidence for present-day leucogranite pluton growth in Tibet, Geology, 32, 801-804. GALVE, A., HIRN, A., JIANG, M. ETAL. 2002a. Modes of raising northeastern Tibet probed by explosion seismology. Earth and Planetary Science Letters, 203, 35-43. GALVE, A., SAPIN,M., HIRN, A., DIAZ, J., LI~PINE,J.-C., LAIGLE, M., GALLART, J. 8z JIANG, M. 2002b. Complex images of Moho and variation of Vp/ Vs across the Himalaya and South Tibet, from a joint receiver-function and wide-angle-reflection approach. Geophysical Research Letters, 29, 2182. DOI: 10.1029/2002GL015611. GALY, A. & FRANCE-LANORD,C. 2001. Higher erosion rates in the Himalaya; geochemical constraints on riverine fluxes. Geology, 29, 23-26. GAO, R., HUANG, D., Lu, D. ETAL. 2000. Deep seismic reflection profile across the juncture zone between the Tarim Basin and the West Kunlun Mountains. Chinese Science Bulletin, 45, 2281-2286. GOES S. & VANDER LEE, S. 2002. Thermal structure of the North American uppermost mantle inferred from seismic tomography. Journal of Geophysical Research, 107. DOI: 10.1029/2000JB000049. GOKARN, S. G., GUPTA, G., RAO, C. K. 8z SELVARAJ,C. 2002. Electrical structure across the Indus Tsangpo suture and Shyok suture zones in NW Himalaya using magnetotelluric studies. Geophysical Research Letters, 29, 1-4. GRASEMANN, B. & VANNAY, J.-C. 1999. Flow controlled inverted metamorphism in shear zones. Journal of Structural Geology, 21, 743-750. GREEN, A. G., MILKEREIT, B., PERCIVAL, J. A. ETAL. 1990. Origin of deep crustal reflections; seismic profiling across high-grade metamorphic terranes in Canada. Tectonophysics, 173, 627-638. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., K~ND1G, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan Crystalline in Bhutan: evidence from quartz microfabrics, Tectonophysics, 260, 21-43. GRUJIC, D., HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel; insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. GUPTA, M. L., SHARMA, S. R., DROLIA, R. K. & S1NGH, S. B. 1983. Subsurface thermal conditions of Puga Valley hydrothermal field, Himalaya, India. Journal of Geophysics, 54, 51-59. HACKER, B. R., GNOS, E., RATSCHBACHER, L., GROVE, M., MCWILLIAMS, M., SOBOLEV, S. V.,

CRUSTAL FLOW IN TIBET WAN, J. & WU, Z. 2000. Hot and dry deep crustal xenoliths from Tibet. Science, 287, 2463-2466. HAINES, S., KLEMPERER,S. L., BROWN, L. ETAL. 2003. Crustal thickening processes in central Tibet: implications of INDEPTH III seismic data. Tectonics, 22, 1-18. HAMMOND, W. C. & HUMPHREYS, E. D. 2000. Upper mantle seismic wave velocity: Effects of realistic partial melt geometries. Journal of Geophysical Research, 105, 10,975-10,986. HARINARAYANA, T., AZEEZ, K. K. A., NAGANJANEYULU, K., MANOJ, C., VEERASWAMY, K., MURTHY, D. N. & RAO, S. P. E. 2004. Magnetotelluric studies in Puga valley geothermal field, NW Himalaya, Jammu and Kashmir, India. Journal of Volcanology and Geothermal Research, 138, 405-424. HARRISON, T. M. 2006. Did the Himalayan Crystallines extrude from beneath the Tibetan Plateau? In: LAW, R. D., SEARLE, M. P. & GODZN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 237 -254. HAUCK, M. L., NELSON, K. D., BROWN, L. D., ZHAO, W. J. & ROSS, A. R. 1998. Crustal structure of the Himalayan Orogen from Project INDEPTH deep seismic reflection profiles at c. 90~ Tectonics, 17, 481-500. HENRY, P., LE PICHON, X. & GOFFE, B. 1997. Kinematic, thermal and petrological model of the Himalayas; constraints related to metamorphism within the underthrust Indian crust and topographic elevation. Tectonophysics, 273, 31-56. HERQUEL, G. & TAPPONNIER, P. 2005. Seismic anisotropy in western Tibet. Geophysical Research Letters, 32, L17306. DOI: 10.1029/2005GL023561. HERQUEL, G., WITTLINGER, G. & GUILBERT, J. 1995. Anisotropy and crustal thickness of NorthernTibet. New constraints for tectonic modeling. Geophysical Research Letters, 22, 1925-1928. HILLEY G. E., BORGMANN, R., ZHANG, P.-Z. & MOLNAR, P. 2005. Bayesian inference of plastosphere viscosities near the Kunhin Fault, northern Tibet. Geophysical Research Letters, 32, L01302. DOI: 10.1029/2004GL021658. HIRN, A., LEPINE, J.-C., JOBERT, G. er AL. 1984a. Crustal structure and variability of the Himalayan border of Tibet. Nature, 307, 23-25. HIRN, A., NERCESSIAN, A., SAPIN, M., JOBERT, G., Xu Z. X., GAO E. Y., Lu, D. Y. & TENG, J. W. 1984b. Lhasa Block mad bordering sutures; a continuation of a 500-km Moho traverse through Tibet. Nature, 307, 25-27. HIRN, A., JIANG, M, SAPIN, M. ETAL. 1995. Seismic anisotropy as an indicator of mantle flow beneath the Himalayas and Tibet. Nature, 375, 571-574. HOCHSTEIN, M. P. & REGENAUER-LIEB,K. 1998. Heat generation associated with collision of two plates; the Himalayan geothermal belt. Journal of Volcanology and Geothermal Research, 83, 75-92. HODGES, K. V. 2006. The channel tunneling-extrusion hypothesis as developed for the HimalayanTibetan orogenic system. In: LAW, R. D.,

65

SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 71-90. HODGES, K. V., HURTADO, J. M. & WHIPPLE, K. X. 2001. Southward extrusion of Tibetan crust and its effect on Himalayan tectonics. Tectonics, 20, 799-809. HOKE, L., LAMB, S., HILTON, D. R. & POREDA, R. J. 2000. Southern limit of mantle-derived geothermal helium emissions in Tibet; implications for lithospheric structure. Earth and Planetary Science Letters, 180, 297-308. HOLBROOK, W. S., CATCHINGS, R. D. & JARCHOW, C. M. 1991. Origin of deep crustal reflections: implications of coincident seismic refraction and reflection data in Nevada. Geology, 19, 175-179. HOLNESS, M. B. 1993. Temperature and pressure dependence of quartz-aqueous fluid dihedral angles: the control of adsorbed H20 on the permeability of quartzites. Earth and Planetary Science Letters, 117, 363-377. HOLT, W. E. 2000. Correlated crust and mantle strain fields in Tibet. Geology, 28, 67-70. HUANG, W.-C., NI, J. F., TILMANN, F. ET AL 2000. Seismic polarization anisotropy beneath the central Tibetan Plateau. Journal of Geophysical Research, 105, 27,979-27,990. HUERTA A. D., ROYDEN L. H. & HODGES K. V. 1998. The thermal structure of collisional orogens as a response to accretion, erosion, and radiogenic heating. Journal of Geophysical Research, 103, 15,287-15,302. HURTADO, J. M., JR., HODGES, K. V. & WHIPPLE, K. 2001. Neotectonics of the Thakkhola Graben and implications for Recent activity on the South Tibetan fault system in the central Nepal Himalaya. Geological Society of America Bulletin, 113, 222-240. HYNDMAN, R. D. • SHEARER, P. M. 1989. Water in the lower continental crust; modelling magnetotelluric and seismic reflection results, Geophysical Journal of the Royal Astronomical Society, 98, 343-365. JACKSON, J. 2003. Faulting, flow and the strength of the continental lithosphere. In: KLEMPERER, S. L. & ERNST, W. G. (eds) The George A. Thompson Volume: The Lithosphere of Western North America and its Geophysical Characterization. Geological Society of America, International Book Series, 7, 5-27. JAMIESON R. A., BEAUMONT, C., MEDVEDEV, S. c~; NGUYEN, i . H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06407. DOI: 10.1029/2003JB002811. JAMIESON, R. A., BEAUMONT, C., NGUYEN, M. H. & GRUJIC, D. 2006. Provenance of the Greater Himalayan Sequence and associated rocks: predictions of channel flow models. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion And Exhumation In Continental

66

S.L. KLEMPERER

Collision Zones. Geological Society, London, Special Publications, 268, 165-182. JAUPART, C., FRANCHETEAU,J. & SHEN, X.-J. 1985. On the thermal structure of the southern Tibetan crust. Geophysical Journal of the Royal Astronomical Society, 81, 131-155. JIN, Y., MCNUTT, M. K. & ZHU, Y. 1996. Mapping the descent of Indian and Eurasian plates beneath the Tibetan plateau from gravity anomalies. Journal of Geophysical Research, 101, 11,275-11,290. JOHNSON, M. R. W. 1994. Volume balance of erosional loss and sediment deposition related to Himalayan uplifts. Journal of the Geological Society, London, 151, 217-220. JONES, A. G. 1992. Electrical conductivity of the continental lower crust. In: FOUNTAIN, D. M., ARCULUS, R. & KAY, R. W. (eds) Continental Lower Crust. Elsevier, Amsterdam, 81 - 143. JORDAN, T. A. & WATTS, A. B. 2005. Gravity anomalies, flexure and the elastic thickness structure of the India-Eurasia collisional system. Earth and Planetary Science Letters, 236, 732750. KAO, H., GAO, R., RAU, R.-J., SHI, D., CHEN, R.-Y., GUAN, T. & Wu, F. T. 2001. Seismic image of the Tarim basin and its collision with Tibet. Geology, 29, 575-578. KAPP, P. & GUYNN, J. H. 2004. Indian punch rifts Tibet. Geology, 32, 993-996. KAPP, J. L. D., HARRISON,T. M., KAPP, P., GROVE, M., LOVERA, O. M. & DING, L. 2005. Nyainqentanghla Shan: a window into the tectonic, thermal and geochemical evolution of the Lhasa block, southern Tibet. Journal of Geophysical Research, 110, B08413. DOI: 10.1029/2004JB003330. KIND, R., NI, J., ZHAO, W., Wu, J. Er AL. 1996. Evidence from earthquake data for a partially molten crustal layer in Southern Tibet. Science, 274, 1692-1694. KIND, R., YUAN, X, SAUL, J. ET AL. 2002. Seismic images of crust and upper mantle beneath Tibet; evidence for Eurasian Plate subduction. Science, 298, 1219-1221. KING, J., HARRIS, N., ARGLES, T. & PARRISH, R. 2005. Constraints on tectonic models from crustal melt formation in southern Tibet. In: MASCLE, G. & LAVg, J. (eds) 20th Himalaya-Karakoram-Tibet Workshop Special Extended Abstracts Volume. Gdologie Alpine Mdmoire H.S. 44, 104-105. KIRBY, S. H. 1984. Introduction and digest to the special issue on chemical effects of water on the deformation and strengths of rocks. Journal of Geophysical Research, 89, 3991-3995. KLEMPERER, S. L. 1987. A relation between continental heat flow and the seismic reflectivity of the lower crust. Journal of Geophysics, 61, 1- 11. KOLA-OJO, O. & MEISSNER, R. 2001. Southern Tibet; its deep seismic structure and some tectonic implications. Journal of Asian Earth Sciences, 19, 249-256. KONG, X., WANG, Q. & XIONG, S. 1996. Comprehensive geophysics and lithospheric structure in the western Xizang (Tibet) Plateau. Science in China (Series D), 39, 348-358.

KOONS, P. O., ZEITLER, P. K., CHAMBERLAIN, C. P., CRAW, D. & MELTZER, A. S. 2002. Mechanical links between erosion and metamorphism in Nanga Parbat, Pakistan Himalaya. American Journal of Science, 302, 749-773. KOSAREV, G., KIND, R., SOBOLEV, S. V., YUAN, X., HANKA,W. & ORESmN, S. 1999. Seismic evidence for a detached Indian lithospheric mantle beneath Tibet. Science, 283, 1306-1309. KUMAR, P., YUAN, X., KIND, R. & NI, J. 2006. Imaging the colliding Indian and Asia lithospheric plates beneath Tibet. Journal of Geophysical Research, 111, B06308. DOI: 10.1029/ 2005JB003930. KUSZNIR, N. J. & MATTHEWS, D. H. 1988. Deep seismic reflections and the deformational mechanics of the continental lithosphere. In: MENZmS, M. A. & Cox, K. G. (eds) Oceanic and continental lithosphere; similarities and differences. Journal of Petrology, Special Lithosphere issue, 63-87. LACHENBRUCH,A. H. &SASS, J. H. 1977. Heat flow in the United States and the thermal regime of the crust. In: HEACOCK, J. G., KELLER, G. V., OLIVER, J. E. & SIMMONS, G. (eds) The Earth's Crust; its Nature and Physical Properties. American Geophysical Union, Geophysical Monographs, 20, 626-675. LANGIN, W. R., BROWN, L. D. & SANDVOL, E. A. 2003. Seismicity of central Tibet from Project INDEPTH III seismic recordings. Bulletin of the Seismological Society of America, 93, 2146-2159. LAVI~, J., AVOUAC, J. P., LACASSIN, R., TAPPONNIER, P. & MONTAGNER, J. P. 1996. Seismic anisotropy beneath Tibet: evidence for eastward extrusion of the Tibetan lithosphere? Earth and Planetary Science Letters, 140, 83-96. LEE, J., HACKER, B. & WANG, Y. 2004. Evolution of North Himalayan gneiss domes: structural and metamorphic studies in Mabja dome, southern Tibet. Journal of Structural Geology, 26, 2297-2316. LEECH, M. L., S1NGH, S., JAIN, A. K., KLEMPERER, S. L. &. MANICKAVASAGAM, R. M. 2005. The onset of India-Asia continental collision: Early, steep subduction required by the timing of UHP metamorphism in the western Himalaya. Earth and Planetary Science Letters, 234, 83-97 LEMONNIER, C., MARQUIS, G., PERRIER, F. ET AL. 1999. Electrical structure of the Himalaya of Central Nepal: high conductivity around the midcrustal ramp along the MHT. Geophysical Research Letters, 26, 3261-3264. LI, J., MILLER, B. V., NELSON, K. D. & SAMSON, S. D. 1998. Two belts of collisional granite in the Himalaya? EOS Transactions of the American Geophysical Union, 79, Suppl., 350. LI, S., UNSWORTH, M. J., BOOKER, J. R., WEI W., TAN H. & JONES, A. G. 2003. Partial melt or aqueous fluid in the mid-crust of southern Tibet? Constraints from INDEPTH magnetotelluric data. Geophysical Journal International, 153, 289-304. MCKENZIE, D. & FAIRHEAD,D. 1997. Estimates of the effective elastic thickness of the continental lithosphere from Bouguer and free air gravity

CRUSTAL FLOW IN TIBET anomalies. Journal of Geophysical Research, 102, 27,523 -27,552. MCKENZIE, D. & JACKSON, J. 2002. Conditions for flow in the continental crust. Tectonics, 21(6), 1055. DOI: 10.1029/2002TC001394. MCKENZIE, D., NIMMO, F., JACKSON, J. A., GANS, P. B. & MmLER, E. L. 2000. Characteristics and consequences of flow in the lower crust. Journal of Geophysical Research, 105, 11,029-11,046. MCNAMARA, D. E., OWENS, T. J., SILVER, P. G. & Wu, F. T. 1994. Shear wave anisotropy beneath the Tibetan Plateau. Journal of Geophysical Research, 99, 13,655-13,666. MCNAMARA, D. E., OWENS, T. J. & WALTER, W. R. 1995. Observation of regional phase propagation across the Tibetan Plateau. Journal of Geophysical Research, 100, 22,215-22,229. MCNAMARA, D. E., WALTER, W. R., OWENS, T. J. & AMMON, C. J. 1997. Upper mantle velocity structure beneath the Tibetan plateau from Pn travel time tomography. Journal of Geophysical Research, 102, 493-505. MAGGI, A., JACKSON, J. A., PRIESTLEY, K. & BAKER, C. 2000. A reassessment of focal depth distributions in southern Iran, the Tien Shan and northern India: Do earthquakes really occur in the continental mantle? Geophysical Journal International, 143, 629-661. MAKOVSKY, Y. & KLEMeERER, S. L. 1999. Measuring the seismic properties of Tibetan bright spots: Evidence for free aqueous fluids in the Tibetan middle crust. Journal of Geophysical Research, 104, 10,795-10,825. MAKOVSKY, Y., KLEMPERER, S. L., HUANG, L.-Y. Lu, D.-Y. & PROJECT INDEPTH TEAM. 1996. Structural elements of the southern Tethyan Himalaya crust from wide-angle seismic data. Tectonics, 15, 997-1005. MAKOVSKY, Y., KLEMPERER, S. L., RATSCHBACHER, L. & ALSDORF, D. 1999. Midcrustal reflector on INDEPTH wide-angle profiles: An ophiolitic slab beneath the India-Asia suture in southern Tibet? Tectonics, 18, 793-808. MALTHE-SORENSSEN, A., PLANKE, S., SVENSEN,H. & JAMTVEIT, B. 2004. Formation of saucer-shaped sills. In: BREITKREUZ, C. & PETFORD, N. (eds)

Physical Geology of High-Level Magmatic Systems. Geological Society, London, Special Publications, 234, 215-227. MASEK, J. G., ISACKS, B. L., FIELDING, E. J. & BROWAEYS, J. 1994. Rift flank uplift in Tibet; evidence for a viscous lower crust. Tectonics, 13, 659-667. MECH1E, J., SOBOLEV, S. V., RATSCHBACHER,L. ETAL. 2004. Precise temperature estimation in the Tibetan crust from seismic detection of the alpha-beta quartz transition. Geology, 32, 601-604. MEDVEDEV, S. & BEAUMONT, C. 2006. Growth of continental plateaus by channel injection: models designed to address constraints and thermomechanical consistency. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile

Extrusion and Exhumation in Continental Collision

67

Zones. Geological

Society, London, Special Publications, 268, 147-164. MEISSNER, R. & STREHLAU,J. 1982. Limits of stresses in continental crusts and their relation to the depth-frequency distribution of shallow earthquakes. Tectonics, 1, 73-89. MELTZER, A., SARKER, G., BEAUDOIN,B., SEEBER, L. & ARMBRUSTER, J. 2001. Seismic characterization of an active metamorphic massif, Nanga Parbat, Pakistan Himalaya. Geology, 29, 651-654. MIBE, K., FuJII, T. & YASUDA,A. 1999. Control of the location of the volcanic front in island arcs by aqueous fluid connectivity in the mantle wedge. Nature, 401, 259-262. MITRA, S., PRIESTLEY, K., BHATTACHARYYA,A. K. & GAUR, V. K. 2005. Crustal structure and earthquake focal depths beneath northeastern India and southern Tibet. Geophysical Journal International, 160, 227-248. MOLNAR, P. 1988. A review of geophysical constraints on the deep structure of the Tibetan Plateau, the Himalaya and the Karakoram, and their tectonic implications. Philosophical Transactions of the Royal Society of London, Series A, 326, 33-88. MOLNAR, P. 2000. Large-Scale Continental Dynamics.

In: Abstracts of the Second Plate Boundary Observatory (PBO) Workshop, Palm Springs, California, 30 October-1 November 2000. (www.scec.org/ news/0 lnews/es_abstracts/molnar.pdf). MOLNAR, P. & CHEN, W.-P. 1983. Focal depths and fault plane solutions of earthquakes under the Tibetan Plateau. Journal of Geophysical Research, 88, 1180-1196. MOLNAR, P. & LYON-CAEN, H. 1989. Fault plane solutions of earthquakes and active tectonics of the Tibetan plateau and its margins. Geophysical Journal of the Royal Astronomical Society, 99, 123-153. MOLNAR, P., ENGLAND, P. & MARTINOD, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the Indian monsoon. Reviews of Geophysics, 31, 357-396. MONSALVE, G., SHEEHAN, A. F., SCHULTE-PECKUM, V., RAJAURE, S., PANDEY, M. R. & Wu, F. 2006. Seismicity and 1-D velocity structure of the Himalayan collision zone: earthquakes in the crust and upper mantle. Journal of Geophysical Research (in press). MYROW, P. M., HUGHES, N. C., PAULSEN, T. S. ETAL. 2003. Integrated tectonostratigraphic analysis of the Himalaya and implications for its reconstruction. Earth Planetary Science Letters, 212, 433 -441. NELSON, K. D., ZHAO, W., BROWN, L. D. ETAL. 1996. Partially molten middle crust beneath southern Tibet; synthesis of Project INDEPTH results. Science, 274, 1684-1688. NI, J. & BARAZANGI,M. 1983. High-frequency seismic wave propagation beneath the Indian Shield, Himalayan Arc, Tibetan plateau and surrounding regions; high uppermost mantle velocities and efficient Sn propagation beneath Tibet. Geophysical Journal of the Royal Astronomical Society, 72, 665 -689.

68

S.L. KLEMPERER

OWENS, T. J. & ZANDT, G. 1997. Implications of crustal property variations for models of Tibetan plateau evolution. Nature, 387, 37-43. OZACAR, A. A. & ZANDT, G. 2004. Crustal seismic anisotropy in central Tibet: Implications for deformational style and flow in the crust. Geophysical Research Letters, 31, L2360l. DOI: 10.1029/ 2004GL021096. PATINO DOUCE, A. & HARRIS, N. 1998. Experimental constraints on Himalayan anatexis. Journal of Petrology, 39, 689-710. PHAM, V. N., BOYER, D., THERME, P., YUAN, X. C., LI, L. & JIN, G. Y. 1986. Partial melting zones in the crust in southern Tibet from magnetotelluric results. Nature, 319, 310-314. PRIESTLEY, K. & MCI~NZIE, D. 2006. The thermal structure of the lithosphere from shear wave velocities. Earth and Planetary Science Letters, 244, 285-301. RAI, S. S., PRIESTLEY, K., GAUR, V. K., SINGH, M. P. & SEARLE, M. 2006. Configuration of the Indian Moho beneath the NW Himalaya and Ladakh. Geophysical Research Letters, 33, L15308. DOI: 10.1029/2006GL026076. RANDALL, G. E., AMMON, C. J. & OWENS, T. J. 1995. Moment-tensor estimation using regional seismograms from portable network deployments. Geophysical Research Letters, 22, 1665-1668. RAPINE R., TILMANN, F., WEST, M., NI, J. & RODGERS, A. 2003. Crustal structure of northern and southern Tibet from surface wave dispersion analysis. Journal of Geophysical Research, 108, 2120. DOI: 10.1029/2001JB000445. ROBINSON, D. M. & PEARSON, O. N. 2006. Exhumation of Greater Himalayan rocks along the Main Central Thrust in Nepal: implications for channel flow. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 255-267. RODGERS, A. J. & SCHWARTZ, S. Y. 1998. Lithospheric structure of the Qiangtang Terrane, northern Tibetan Plateau, from complete regional waveform modeling: Evidence for partial melt. Journal of Geophysical Research, 103, 71377152. ROSENBERG, C. L. & HANDY, M. R. 2005. Experimental deformation of partially melted granite revisited; implications for the continental crust. Journal of Metamorphic Geology, 23, 19-28. Ross, A. R., BROWN, L. D., PANANONT, P., NELSON, K. D., KLEMPERER, S. L., HAINES, S., ZHAO, W. & Guo, J. 2004. Deep reflection surveying in central Tibet: lower-crustal layering and crustal flow. Geophysical Journal International, 156, 115-128. ROYDEN, L. H. 1996. Coupling and decoupling of crust and mantle in convergent orogens; implications for strain partitioning in the crust. Journal of Geophysical Research, 101, 17,679-17,705. ROYDEN, L. H., BURCHFIEL, B. C., KING, R. W., WANG, E., CHEN Z., SHEN F. & LIU Y. 1997.

Surface deformation and lower crustal flow in eastern Tibet. Science, 276, 788-790. RUPPEL, C. & MCNAMARA, D. E. 1997. Seismic and rheological constraints on the thermal state of Tibetan plateau upper mantle: implications for melt production, mantle delamination and largescale tectonics. EOS Transactions of the American Geophysical Union, 78, 650. RUZAIKIN, A. I., NERSESOV, I. L., KHALTURIN,V. I. & MOLNAR, P. 1977. Propagation of Lg and lateral variations in crustal structure in Asia. Journal of Geophysical Research, 82, 307-316. SANDVOL, E., NI, J., KaND, R. & ZHAO, W. 1997. Seismic anisotropy beneath the southern Himalayas-Tibet collision zone. Journal of Geophysical Research, 102, 17,813-17,824. SAPIN, M. & HIRN, A. 1997. Seismic structure and evidence for eclogitization during the Himalayan convergence. Tectonophysics, 273, 1-16. SCHULTE-PELKUM, V., GASPAR, M., SHEEHAN, A., PANDEY, M. R., SAPKOTA, S., BILHAM, R. & Wu, F. 2005. Imaging the Indian subcontinent beneath the Himalaya. Nature, 435, 1222-1225. DOI: 10.1038/nature03678. SEARLE, M. P. & REX, A. J. 1989. Thermal model for the Zanskar Himalaya. Journal of Metamorphic Geology, 7, 127-134. SEARLE, M. P. & SZULC, A. G. 2005. Channel flow and ductile extrusion of the High Himalayan slab - the Kangchenjunga- Darjeeling profile, Sikkim Himalaya. Journal of Asian Earth Sciences, 25, 173-185. SHANKER, R., PADHI, R. N., ARORA, C. L., I~AKASH, G., THUSSU, J. L & DUA, K. J. S. 1976. Geothermal exploration of the Puga and Chumathang geothermal fields, Ladakh, India. In: Proceedings, Second United Nations Symposium on the Development and Use of Geothermal Resources, San Francisco, USA, 20-29 May 1975. US Government Printing Office, Washington DC, 1, 245-258. SHAPIRO, N. M., RITZWOLLER, M. H., MOLNAR, P. & LEVIN, V. 2004. Thinning and flow of Tibetan crust constrained by seismic anisotropy. Science, 305, 233 -236. SHEN, F., ROYDEN, L. H. & BURCHFIEL, B. C. 2001. Large-scale crustal deformation of the Tibetan Plateau. Journal of Geophysical Research, 106, 6793-6816. SHERRINGTON, H. F., ZANDT, G. & FREDERIKSEN, A. 2004. Crustal fabric in the Tibetan plateau based on waveform inversions for seismic anisotropy parameters. Journal of Geophysical Research, 109, B02312. DOI: 10.1029/ 2002JB002345. SHI, D, ZHAO, W., BROWN, L. e r AL. 2004. Detection of southward intra-continental subduction of Tibetan lithosphere along the Bangong-Nujiang suture by P-to-S converted waves. Geology, 32, 209-212. SILVER, P. G. 1996. Seismic anisotropy beneath the continents: probing the depths of geology. Annual Reviews of Earth and Planetary Science, 24, 385 -432.

CRUSTAL FLOW IN TIBET SPRATT, J. E., JONES, A. G., NELSON, K. D., UNSWORTH, M. J. & INDEPTH MT TEAM. 2005. Crustal structure of the India-Asia collision zone, southern Tibet, from INDEPTH MT investigations. Physics of the Earth and Planetary Interiors, 150, 227-237. TAPPONNIER,P., PELTZER, G., LEDAIN,A. Y., ARMIJO, R. & COBBOLD, P. 1982. Propagating extrusion tectonics in Asia-new insights from simple experiments with plasticine, Geology 10, 611-616. TAPPONNIER, P., Xu, Z. Q., ROGER, F., MEYER, B., ARNAUD, N., WITTLINGER, G. & YANG, J. S. 2001. Oblique stepwise rise and growth of the Tibet plateau. Science 294, 1671 - 1677. TAYLOR, M., YIN, A., RYERSON, F. J., KAPP, P. & DING, L. 2003. Conjugate strike-slip faulting along the Bangong-Nujiang suture zone accommodates coeval east-west extension and northsouth shortening in the interior of the Tibetan Plateau. Tectonics, 22(4), 1044. DOI: 10.1029/ 2002TC001361. THOMPSON, A. B. & CONNOLLY, J. A. D. 1995. Melting of the continental crust: Some thermal and petrological constraints on anatexis in continental collision zones and other tectonics settings. Journal of Geophysical Research, 100, 15,56515,579. THOMSON, K. 2004. Sill complex geometry and internal architecture: A 3D seismic perspective. In: BREITKREUZ,C. & PETFORD, N. (eds) Physical Geology of High-Level Magmatic Systems. Geological Society, London, Special Publications, 234, 229-232. TIAN, X., Wu, Q., ZHANG, Z., TENG, J. & ZENG, R. 2005. Joint imaging by teleseimic converted and multiple waves and its applications in the INDEPTH-III passive seismic array. Geophysical Research Letters, 32, L21315. DOI: 10.1029/ 2005GL023686. TIKOFF, B., RUSSO, R., TEYSSIER, C. • TOMMASI, A. 2004. Mantle-driven deformation of orogenic zones and clutch tectonics. In: GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227, 41-64. TILMANN, F., NI, J. & INDEPTH III SEISM1C TEAM. 2003. Seismic imaging of the downwelling Indian lithosphere beneath central Tibet. Science, 300, 1424-1427. TORAMARU, A. & FuJII, N. 1986. Connectivity of melt phase in a partially molten peridotite. Journal of Geophysical Research, 91, 9239-9252. TURNER, S., ARNAUD, N., LIU, J. ET AL. 1996. Postcollision, shoshonitic volcanism on the Tibetan Plateau; implications for convective thinning of the lithosphere and the source of ocean island basalts. Journal of Petrology, 37, 45-71. UNSWORTH, M., WEI W., JONES, A. G. ET AL. 2004. Crustal and upper mantle structure of northern Tibet imaged with magnetotelluric data. Journal of Geophysical Research, 109, B02403. DOI: 10.1029/2002JB002305.

69

UNSWORTH, M. J., JONES, A. G., TEl W., MARQUIS, G., GOKARN, S. G., SPRATT, J. E. & THE INDEPTH-MT TEAM. 2005. Crustal rheology of the Himalaya and Southern Tibet inferred from magnetotellufic data. Nature, 483, 78-81. DOI: 10.1038/ nature04154. VERGNE, J., WITTLINGER, G., QIANG H., TAPPONNIER, P., POUPINET, G., JIANG M., HERQUEL, G. & PAUL, A. 2002. Seismic evidence for stepwise thickening of the crust across the NE Tibetan Plateau. Earth and Planetary Science Letters, 203, 25-33. VERGNE J., WITTLINGER, G., FARRA, V. • Su, H. 2003. Evidence for upper crustal anisotropy in the Songpan-Ganze (northeastern Tibet) terrane. Geophysical Research Letters, 30. DOI: 10.1029/ 2002GL016847. WAFE, H. S. & BULAU, J. R. 1979. Equilibrium fluid distribution in an ultramafic partial melt under hydrostatic stress conditions. Journal of Geophysical Research, 84, 6109-6114. WANG Q., ZHANG P., FREYMUELLER,J. T. ETAL. 2001. Present-day crustal deformation in China constrained by Global Positioning System measurements. Science, 294, 574-577. WANG, Q., MCDERMOTT, F., XU, J.-F., HERVI~ BELLON, H. & ZHU, Y.-T. 2005. Cenozoic K-rich adakitic volcanic rocks in the Hohxil area, northern Tibet: Lower-crustal melting in an intracontinental setting. Geology, 33, 465-468. WANG, Y. 2001. Heat flow pattern and lateral variations of lithosphere strength in China mainland; constraints on active deformation. Physics of the Earth and Planetary Interiors, 126, 121-146. WARNER, M. R. 1990. Basalts, water, or shear zones in the lower continental crust? Tectonophysics, 173, 163-174. WATSON, E. B. & LUPULESCU,A. 1993. Aqueous fluid connectivity and chemical transport in clinopyroxene-rich rocks. Earth and Planetary Science Letters, 117, 279-294. WEI, W., UNSWORTH, M., JONES, A. ET AL. 2001. Detection of widespread fluids in the Tibetan crust by magnetotelluric studies. Science, 292, 716-718. WILLETT, S. D. 1999. Orogeny and orography: The effects of erosion on the structure of mountain belts. Journal of Geophysical Research, 104, 28,957-28,982. WILLETT, S. D. & BEAUMONT, C. 1994. Subduction of Asian lithospheric mantle beneath Tibet inferred from models of continental collision. Nature, 369, 642-645. WILLIAMS, H. M., TURNER, S. P., PEARCE, J. A., KELLEY, S. P. & HARRIS, N. B. W. 2004. Nature of the source regions for post-collisional, potassic magmatism in southern and northern Tibet from geochemical variations and inverse trace element modelling. Journal of Petrology, 45, 555-607. WITTLINGER, G., MASSON, F., POUPINET, G. ET AL. 1996. Seismic tomography of northern Tibet and Kunlun; evidence for crustal blocks and mantle velocity contrasts. Earth and Planetary Science Letters, 139, 263-279.

70

S.L. KLEMPERER

WITTLINGER, G., TAPPONNIER, P., POUPINET, G., MEI, J., SHI, D., HERQUEL, G. & MASSON, F. 1998. Tomographic evidence for localized lithosphefic shear along the Altyn Tagh Fault. Science, 281, 74-76. WITTLINGER, G., VERGNE, J., TAPPONNIER, P. ET AL. 2004. Teleseismic imaging of subducting lithosphere and Moho offsets beneath western Tibet. Earth and Planetary Science Letters, 221, 117-130. WOBUS, C. W. HODGES, K. V. & WHIPPLE, K. X. 2003. Has focused denudation sustained active thrusting at the Himalayan topographic front? Geology, 31, 861-864. WU, C., NELSON, K. D., WORTMAN, G., SAMSON, S. D., YUE, Y., LI, J., KIDD, W. S. F. & EDWARDS, M. A. 1998. Yadong cross structure and South Tibetan Detachment in the east central Himalaya Tectonics, 17, 28-45. DOI: (89 ~176 10.1029/97TC03386. XIE, J., GOK, R., NI, J. & AOKI, Y. 2004. Lateral variations of crustal seismic attenuation along the INDEPTH profiles in Tibet from Lg Q inversion. Journal of Geophysical Research, 109, B10308. DOI: 10.1029/2004JB002988. YOSHINO T., MIBE, K., YASUDA, A. & FuJII, T. 2002. Wetting properties of anorthite aggregates: Implications for fluid connectivity in continental lower crust. Journal of Geophysical Research, 107. DOI: 10.1029/2001JB000440. YUAN, X., NI, J., KIND, R., MECHIE, J. & SANDVOL, E. 1997. Lithospheric and upper mantle structure of southern Tibet from a seismological passive source experiment. Journal of Geophysical Research, 102, 27,491 - 27,500. ZEITLER, P. K., KOONS, P. O., BISHOP, M. P. ET AL. 2001. Crustal reworking at Nanga Parbat, Pakistan: Metamorphic consequences of thermal-mechanical coupling facilitated by erosion. Tectonics, 20, 712-728. DOI: 10.1029/2000TC001243. ZHANG, P.-Z., SHEN, Z., WANG, M. ETAL. 2004. Continuous deformation of the Tibetan plateau from global positioning system data. Geology, 32, 809- 812. ZHANG, Z.-J. & KLEMPERER, S. L. 2005. West-east variation in crustal thickness in northern Lhasa

block, central Tibet, from deep seismic sounding data. Journal of Geophysical Research, 110, B09403. DOI: 10.1029/2004JB003139. ZHAO, J., MOONEY, W. D., ZHANG,X., LI, Z., J1N, Z. & OKAYA, N. 2006. Crustal structure across the Altyn Tagh Range at the northern margin of the Tibetan plateau and tectonic implications. Earth and Planetary Science Letters, 241, 804-814. ZHAO, L. S. 8z HELMBERGER,D. V. 1991. Geophysical implications from relocations of Tibetan earthquakes; hot lithosphere. Geophysical Research Letters, 18, 2205-2208. ZHAO, L. S., SEN, M. K., STOFFA, P. & FROHLICH, C. 1996. Application of very fast simulated annealing to the determination of the crustal structure beneath Tibet. Geophysical Journal International, 125, 355-370. ZHAO, W.-J., NELSON, K. D. & PROJECT INDEPTH TEAM. 1993. Deep seismic reflection evidence for continental underthrusting beneath southern Tibet. Nature, 366, 557-559. ZHAO, W., MECHIE, J., BROWN, L. D. ET AL. J. 2001. Crustal structure of central Tibet as derived from project INDEPTH wide-angle seismic data. Geophysical Journal International, 145, 486-498. ZHAO, W.-L. & MORGAN, W. J. 1987. Injection of Indian crust into Tibetan lower crust: A twodimensional finite element model study. Tectonics, 6, 489-504. Znu, L. & HELMBERGER, D. V. 1996. Intermediate depth earthquakes beneath the India-Tibet collision zone. Geophysical Research Letters, 23, 435 -438. ZHU, L. & HELMBERGER, D. V. 1998. Moho offset across the northern margin of the Tibetan Plateau. Science, 281, ll70-1172. ZOBACK, M. D., TOWNEND, J. & GROLLIMUND, B. 2003. Steady-state failure equilibrium and deformation of intraplate lithosphere. In: KLEMPERER, S. L. & ERNST, W. G. (eds) Tile George A. Thompson Volume: The Lithosphere of Western North America and its Geophysical Characterization. Geological Society of America, International Book Series, 7, 50-68.

A synthesis of the Channel Flow-Extrusion hypothesis as developed for the Himalayan-Tibetan orogenic system K. V. H O D G E S

Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02139, USA Present address: School of Earth and Space Exploration, Arizona State University, Tempe, AZ 85287-1404, USA (e-mail: [email protected]) Abstract: Surface and subsurface geological features of the Himalayan-Tibetan orogenic system may be explained by three sets of processes: those related to plate convergence, those related to the gravitational spreading of a fluid middle crust beneath the Tibetan Plateau, and those related to aggressive erosion along the southern margin of the plateau. In this paper, the possible relationships among the last two of these--and their tectonic manifestations--are presented in the form of a 'Channel Flow-Extrusion' hypothesis. This hypothesis, deriving from a series of ideas advanced by many geologists and geophysicists over the past two decades, suggests the definition of three phases in the Early Miocene-Recent history of the orogenic system. During Phase I (Early Miocene), the crust of southern Tibet was sufficiently hot and thick to enable lateral flow of a weak middle crust. To the north and east, this flow resulted in the expansion of the Tibetan Plateau. To the south, erosion at the Himalayan front permitted the mid-crustal channel to breach the surface; this process is recorded in the deformational history of the Himalayan metamorphic core and the Main Central and South Tibetan fault systems that bound it. While the lateral expansion of the plateau by mid-crustal flow has continued throughout Neogene time, some evidence suggests that extrusion across the Himalayan front waned substantially during the Middle Miocene-Early Pliocene interval (Phase II). In Middle Miocene time, large magnitude extension of the decoupled upper crust of southern Tibet led to the development of a subsidiary channel; its extrusion explains the existence of the North Himalayan gneiss domes. Phase III (Late Pliocene-Recent) has involved renewed southward extrusion of the main channel due to climatically induced increases in the erosion rate at the Himalayan range front. Although numerous models have been advanced to explain why orogenic plateaus exist (e.g. James 1971; Dewey & Burke 1973; Powell & Conaghan 1973; Zhao & Morgan 1985; England & Houseman 1988; Isacks 1988; Allmendinger et al. 1997), less attention has been paid to the effect of these features, once developed, on the subsequent evolution of the orogenic systems in which they occur. An important characteristic of plateau regions like the Puna-Altiplano and Tibet is that they have much thicker crust than surrounding regions. The associated imbalance in the gravitational potential energy results in the mechanical tendency for a plateau region to spread outward to the extent permitted by the rheology of its crust and the surrounding crust, and the extent to which the plateau region is supported dynamically by plate convergence (Artyushkov 1973; Gratton 1989; Bird 1991). If significant spreading occurs, we might expect there to be a clear signature of it in the geological record, especially in the form of structural features with kinematics that are better explained by plateau spreading than by plate convergence alone. This appears to be the case in the Himalaya and Tibet (Fig. 1), where a variety of geological and geophysical observations have led to a tectonic model in

which lateral spreading of the Tibetan Plateau has occurred throughout much of late Cenozoic time by a combination of faulting in the brittle upper crust, channelized flow of a more ductile middle crust, and aggressive erosion along the Himalayan range front (Beaumont et al. 2001; Hodges et al. 2001). This paper begins with a review of the history leading up to development of this 'Channel F l o w - E x t r u s i o n ' hypothesis, includes a concise statement of the hypothesis, and ends with suggestions about how the hypothesis might be tested effectively through future studies.

Evolution of the hypothesis The Channel Flow-Extrusion hypothesis is really a convergence of ideas that were developed to explain the physiography of Tibet and the structural architecture of the southern and eastern plateau margins.

Idea 1: The crust beneath Tibet is partially molten Early speculation (e.g. Dewey & Burke 1973) that the middle and lower crust beneath the Tibetan

From: LAW, R. D., SEARLE,M. P. & GODIN, L. (eds) ChannelFlow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 71-90. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

72

K.V. HODGES

Fig. 1. Generalized map of the Himalayan-Tibetan orogenic system. Medium-grey field indicates region with elevations in excess of 3000 m, including the Tibetan Plateau and the Tian Shan. White regions are sedimentary basins. Dark-grey region encompasses the Himalayan range front. Thick solid lines indicate major transcurrent fault systems and thrust systems (barbed). Thick dashed line indicates only the portion of the INDEPTH reflection seismic line referred to in the text; the complete INDEPTH geophysical continued northward across the plateau. Patterned areas show regions where geological observations led to the development of the Channel Flow-Extrusion hypothesis. Specific localities mentioned in the text are: 1, Yadong cross-structure; 2, Kangmar dome; 3, Marsiyandi Valley; 4, Kali Gandaki Valley; 5, Sutlej Valley; 6, Tso Morari dome.

Plateau may be partially molten received strong support from heat flow measurements in the 1980s that indicated an exceptionally high geothermal gradient (Francheteau et al. 1984). Subsequently, Alsdorf & Nelson (1999) showed how satellite magnetic data indicate temperatures in excess of 550~ at 1 0 - 2 0 k m depth across much of the plateau. Deep crustal xenoliths from c. 3 Ma volcanic fields in the central plateau suggest that temperatures of 800-1000~ were attained at depths of 3 0 - 5 0 k m (Hacker et al. 2000). Based on a seismic velocity model derived from active- and passive-source seismic data, Mechie et al. (2004) inferred the presence of the c~-/3 quartz transition, corresponding to a temperature of c. 700~ at depths of 1 8 - 2 0 km beneath the central plateau and roughly 32 km beneath the southern plateau. Comparing this thermal structure with experimental constraints on the conditions required for crustal anatexis (e.g. Holland & Powell 2001; White et al. 2001), partial melting might be taking place across much of the plateau at depths greater than about 15 to 20 km.

Evidence for the existence of such melts comes from the seismic character of the Tibetan crust as revealed by analysis of earthquake data. Kind et al. (1996), Yuan et al. (1997) and Cotte et al. (1999) noted the existence of a mid-crustal, lowvelocity zone in southern Tibet that suggests the presence of melt. Low velocities in the lower crust of northern Tibet were similarly interpreted by Owens & Zandt (1997). Fan & Lay (2003) documented strong attenuation of Lg energy across the plateau, which is readily interpreted as a consequence of partial melting. Some of the strongest indications of the possible presence of significant volumes of melt were found during geophysical surveys undertaken as part of the multinational INDEPTH project in the 1990s (Nelson et al. 1996). Multichannel seismic reflection data collected along a N E - S W transect between 89 and 91 ~ (Fig. 1) distinguished a prominent subhorizontal band of reflectors at a depth of c. 5 - 6 seconds, or 15-18 km assuming reasonable seismic velocities. These reflectors, which are well-defined from c. 29~ at least as far north

CHANNEL FLOW-EXTRUSION HYPOTHESIS as c. 30~ are characterized by extremely high amplitudes compared to other reflectors on the INDEPTH line and are thus commonly referred to as 'bright spots'. Based on the seismic characteristics of the band of bright spots, Brown et al. (1996) concluded that it was most plausibly interpreted as an interface between solid rock above and partially molten rock below. However, Makovsky & Klemperer (1999) noted that some characteristics of wide-angle INDEPTH data most notably the P-wave to S-wave conversion from the bright spots - might be more consistent with the presence of aqueous fluids rather than partial melts. Magnetotelluric data from southern Tibet confirm the regional extent of the features recorded in the active-source seismic data but do not provide a definitive resolution of the debate regarding their origin. In the 1980s, Sino-French reconnaissance magnetotelluric surveys had indicated the presence of unusually high electrical conductivity in the crust of southern Tibet (Van Ngoc et al. 1986). A more detailed study conducted along the INDEPTH transect confirmed the existence of a high-conductivity layer, and showed that it coincided with the bright spots in the seismic profile (Chen et al. 1996). Chen and co-workers agreed with the Sino-French team that magma bodies provided a reasonable explanation for the high conductivity. Wei et al. (2001) demonstrated that the highconductivity layer extended well beyond the INDEPTH footprint, but suggested that its magnetotelluric signature might be explained as easily by regionally distributed saline fluids as by a horizon of partial melt. Li et al. (2003) suggested that the combined seismic and magnetotelluric data-sets might require an explanation that involved both the presence of aqueous fluids and melts. More recently, Gaillard et al. (2004) showed that experimentally determined electrical properties of leucogranitic melts provide an excellent match with the highconductivity zones observed in Tibet. Makovsky & Klemperer (1999) and Gaillard et al. (2004) also pointed out that, even if the bright spots and highconductivity zone are caused by an aqueous fluid, the necessary fluid concentrations are difficult to reconcile with typical volatile contents of midcrustal metamorphic rocks without appealing to the dewatering of leucogranitic magmas during the last stages of their crystallization (Scaillet et al. 1995).

I d e a 2: T h e T i b e t a n f l u i d l a y e r is s u f f i c i e n t l y w e a k to h a v e f l o w e d l a t e r a l l y

Unlike the great mountain ranges that occur along its margins, the interior of Tibet has very limited topographic relief (Fielding et al. 1994). The

73

plateau nevertheless occurs at a very high elevation, thousands of metres above the Indian subcontinent to the south, the Tarim and Qaidam basins to the north, and south Asia to the east. Geophysical surveys have shown that the crust beneath Tibet is very thick compared to the crust underlying surrounding regions, and the high elevation of Tibet is an isostatic effect related to this abnormality (Molnar et al. 1993). But why is it so flat? One simple explanation may be that the middle and lower crust of the plateau is a fluid that is too weak to support large variations in surface topography and it simply flows laterally under the influence of gravity (Bird 1991; Fielding et al. 1994; Masek et al. 1994). If large amounts of partial melting had occurred in the middle and lower Tibetan crust, it seems intuitively obvious that the consequent weakening of the crust could accommodate significant lateral flow. However, the preponderance of evidence suggests the presence of a few weight per cent of melt at most (Hacker et al. 2000; Kind et al. 2002). A more fruitful way to evaluate the potential for flow is to consider the thermal and compositional dependence of the effective viscosities of metamorphic rocks that are likely candidates for Tibetan middle and lower crust (Brace & Kohlstedt 1980; Ranalli 1997). The most recent rheological studies of rocks that are likely analogues for the middle crust suggest that they would have extremely low viscosities given the geothermal gradient beneath Tibet (Kenis et al. 2005). In contrast, the anhydrous xenoliths found in the volcanic fields of north-central Tibet (Hacker et al. 2000) which are the only actual samples we have of Tibetan lower crust - would be very strong and resistant to flow (Jackson et al. 2004). The xenoliths may not constitute a representative sampling, however, and several lines of evidence suggest that effective viscosities are, in reality, low throughout much of the middle and lower crust. The gravity spectrum of Tibet strongly suggests a rheologically layered crustal structure that is best modelled with a weak lower crust (Jin et al. 1994). Crustal flexural rigidity, as calculated from rift flank uplifts in central and southern Tibet, requires viscosities below c. 1022 Pa s in the lower crust (Masek et al. 1994). Virtually all crustal earthquakes beneath the Tibetan Plateau occur at depths of less than 25 km, suggesting that all deformation in the crust below is aseismic and ductile (Langin et al. 2003). More generally, the seismic wave response of the Tibetan crust below about 20 km is consistent with low strength (Owens & Zandt 1997; Yuan et al. 1997; Cotte et al. 1999), and this conclusion is supported by many other types of geophysical evidence (Klemperer 2006). As Meissner & Mooney (1998) put it, the 'Tibetan Plateau shows

74

K.V. HODGES

much lower viscosities in the middle and lower crust than anywhere else on Earth'. Given such a theology, it seems highly probable that at least the middle crust of Tibet might have experienced lateral flow. The earliest quantitative assessment of the implications of this flow on the tectonic evolution of the Himalayan-Tibetan orogenic system was that of Zhao & Morgan (1987), who suggested that injection of Indian crust into a fluid Tibetan lower crust could explain the wholesale uplift of the Tibetan Plateau. While subsequent geological and geophysical research has rendered this model unlikely, several theoretical studies have explored the crustal-scale modes of deformation that might have accompanied flow beneath Tibet (e.g. Bird 1991; Burg et al. 1994; Royden 1996; Royden et al. 1997; Meissner & Mooney 1998; Shen et al. 2001). A key result of all these studies is that brittle deformation in the upper 25 km of Tibetan crust is likely to be almost fully decoupled from ductile flow in the middle crust. One measure of the possible nature and extent of flow may be the radial anisotropy of Rayleigh and Love wave propagation across the plateau. Shapiro et al. (2004) showed that the observed anisotropy was most easily explained by development of a strong preferred alignment of micas, which they interpreted as the result of rotation of micas during flattening and thinning of a laterally extruding channel. The seismic data require that the anisotropy occur in a mid-crustal zone roughly 2 5 - 3 0 k m thick, but Shapiro and co-workers could not rule out the existence of a zone up to 50 km thick that extended into the lower crust. Those authors went on to note that 20-40% thinning of the Tibetan crust was required by the seismic data, and that at least half of this was accommodated by outward flow of a material

from a mid-to-lower crustal channel under the influence of gravity.

I d e a 3: T h e T i b e t a n P l a t e a u h a s g r o w n e a s t w a r d by l o w e r - t o - m i d d l e

crustal flow

The eastern margin of the Tibetan Plateau is marked by elevation and crustal thickness gradients that, along some segments, rival those of the Himalayan front. A variety of geomorphic and geochronologic arguments, many based on the incision histories of major river systems, suggest that this part of the plateau has experienced sustained uplift and eastward growth since Late Miocene-Early Pliocene time (Kirby et al. 2002; Clark et al. 2004; Schoenbohm et al. 2004). However, there is little evidence of this growth in the surface structural geology. Virtually all shortening structures along this margin pre-date plateau development, and the only significant structures that are contemporary with plateau growth are strike-slip and associated normal faults that accommodate block rotation around the eastern Himalayan syntaxis (e.g. Burchfiel et al. 1995; Kirby et al. 2000; Wang & Burchfiel 2000). Moreover, the modern strain field indicated by GPS data implies minimal surface shortening across the margin (King et al. 1997; Chen et al. 2000; Zhang et al. 2004). Such observations have prompted the development of a model in which eastward plateau growth reflects flow in the middle crust (Fig. 2). This idea can be traced to seminal papers by Gratton (1989) and Bird (1991). It was expanded on subsequently by Westaway (1995), Royden et al. (1997) and Clark & Royden (2000). The last of these papers emphasized the progressive development of the plateau by eastward flow of a middle-to-lower

Fig. 2. Schematic cross-section of a tunnelling mid-crustal channel. Flow from left to right, driven by the gravitational potential energy gradient between thicker and thinner crust, results in inflation of the thinner crust and surface uplift that propagates in the direction of flow.

CHANNEL FLOW-EXTRUSION HYPOTHESIS crustal channel since at least Middle Miocene time. An important component of the Clark & Royden (2000) model is that lateral extrusion of the lower crust depends on the rheology of adjacent crustal sections (cf. Westaway 1995). If the adjacent lower crust is weak, lateral extrusion of the channel is relatively easy and the result should be a low topographic gradient at the plateau margin, such as that observed on the southeastern edge of the plateau. If the adjacent lower crust is strong, as appears to be the case in the vicinity of the Sichuan Basin (Fig. 1), lateral flow of the channel is inhibited and an abrupt plateau margin forms.

I d e a 4: M i o c e n e s l i p o n H i m a l a y a n

thrust

faults and normal faults was coeval

One of the most crucial observations to contribute to the Channel Flow-Extrusion hypothesis is an apparent developmental relationship between thrust faulting and normal faulting in the Himalaya (Burchfiel & Royden 1985; Searle & Rex 1989). Since the early twentieth century, the basic structure of the Himalaya has been known to be dominated by a series of south-vergent thrust fault systems (Heim & Gansser 1939; Le Fort 1975; Searle 1986; Hodges 2000). The oldest and northernmost of these - the Main Central thrust (MCT) system - places high-grade metamorphic rocks and leucogranites of the Greater Himalayan Sequence on slightly lower grade PrecambrianPalaeozoic metasedimentary and metaigneous rocks of the Lesser Himalayan Sequence (Fig. 3). The Lesser Himalayan Sequence is itself deformed by a complex set of structures that includes the Dadeldhura and Ramgarh thrust systems, as well as the Lesser Himalayan duplex, of DeCelles et al. (2001) and Pearson & DeCelles (2005). Collectively, these features will be referred to here as the Lesser Himalayan thrust (LHT) system. Farther south, the Main Boundary thrust (MBT) system juxtaposes Lesser Himalayan hanging wall rocks against footwall deposits of the Siwalik Molasse, and the molasse deposits themselves override the Indian foreland along the Main Frontal thrust (MFT) system. All of these thrust systems apparently sole into a basal Main Himalayan thrust (MHT), which serves as the decollement along which India subducts beneath Tibet (Schelling & Arita 1991; Schelling 1992; Zhao et al. 1993; Hauck et al. 1998). In general, these fault systems developed progressively from north to south; earliest slip on the MCT, LHT, MBT and MFT occurred at c. 2 2 - 2 0 M a , 15-5 Ma, c. 5 Ma, and < 3 Ma, respectively (Hodges 2000; DeCelles et al. 2001).

75

It was not until the 1980s that an entirely different kind of fault system was discovered in the Himalaya. Caby et al. (1983), Burg et al. (1984a), Burchfiel & Royden (1985), Searle (1986) and Herren (1987) noted the existence of a shallowly north-dipping, normal-sense shear zone between the high-grade gneisses and leucogranites of the Greater Himalayan Sequence and low-grade and unmetamorphosed sedimentary rocks of the overlying Tibetan Sedimentary Sequence, and Burchfiel et al. (1992) subsequently demonstrated that the shear zone was a regionally important feature that could be traced along much of the length of the Himalaya at the approximate position of the southern margin of the Tibetan Plateau. Several studies have shown that this South Tibetan fault (STF; Fig. 3) system includes strike-slip faults as well as rooted detachments (e.g. P~cher 1991; Coleman 1996). Soon after the first reconnaissance studies of the STF system, it became apparent that the earliest normal-sense displacement on the system was roughly synchronous with early slip on the structurally lower MCT system (e.g. Hodges et al. 1992, 1993; Searle et al. 1992; Searle 1996; Dbzes et al. 1999). Such studies confirmed the earlier speculation of Burchfiel & Royden (1985) that the STF and either the MCT or MBT systems were simultaneously active, and that together they accommodated southward extrusion of the Greater Himalayan metamorphic core of the Himalaya. In their diagrams, Burchfiel & Royden illustrated the metamorphic core as a northward-tapering wedge, but noted that this geometry was purely speculative. They presented a simple elastic model illustrating how steep topographic gradients - such as those that now exist along the southern margin of the Tibetan Plateau - can result in a stress field consistent with synchronous slip on shallowly dipping thrust and normal faults in a continent-continent collisional setting. However, Burchfiel & Royden were quick to point out that their use of an elastic model to illustrate their point did not mean that they regarded the deformation as purely elastic. Since 1985, a variety of workers have focused their efforts on understanding strain patterns within the ductilely deformed metamorphic core and the implications of these patterns for the mechanics of extrusion in Early Miocene time (Grnjic et al. 1996, 2002; Grasemann et al. 1999; Vannay & Grasemann 2001; Law et al. 2004; Jessup et al. 2006). If the STF system developed as the upper boundary of an extruding channel - a 'stretching fault' following the terminology of Means (1989) - then its protracted deformational history has important implications for the longevity of the extrusion process. Robust geochronologic constraints on the

76

K.V. HODGES

Fig. 3. Conceptual cross-sections illustrating the three phases of channel extrusion at the Himalayan front since Early Miocene time. Dark-grey shading designates the intact, downgoing Indian plate. Fields with light-grey shading, random-dash patterning, and no shading are Indian crust that has been accreted to the overriding plate. Unpatterned material corresponds to unmetamorphosed to weakly metamorphosed Tibetan sedimentary series. Material with random dash patterning includes high-grade metamorphic rocks of extruded channels. The actively extruding material in each frame has a light-red overlay pattern. Previously extruded material has no overlay shading. Note the development of a ductile shear zone (dashed heavy line) at the base of the actively extruding material in the frame for Phase II. The dark-red shading in the Phase III frame indicates partially molten material as imaged in the INDEPTH seismic reflection experiment (Nelson et al. 1996). Half-arrows show the slip on various faults. Larger, freeform arrows indicate large-scale kinematics. Circled 'A' indicates the divergence point of the downgoing Indian Plate and the extruding channel. 'B' indicates the hypothetical position of the lower-crustal duplex discussed in the text. Indian plate material accreted to the system across this duplex is continually incorporated into the channel. The blue bars in each frame represent the zone of orographic rainfall, with the colour gradient representing the rainfall gradient in a schematic way - lighter colours indicate lighter rainfall. Note that average rainfall is much lower during Phase II, a speculation based on sedimentological evidence for lower erosion rates. The green bar indicates a zone of extension over the North Himalayan gneiss domes, with the gradient representing intensity of extensional strain in upper crustal material. Acronyms: LHT, Lesser Himalayan thrust system; MBT, Main Boundary thrust system; MCT, Main Central thrust system; MFT, Main Frontal thrust system; MHT, Main Himalayan thrust; STF, South Tibetan fault system.

timing o f ductile shearing on STF structures range f r o m about 22 to 1 2 M a (Hodges et al. 1992, 1996, 1998; Edwards & Harrison 1997; Searle et al. 1997, 2003; C o l e m a n 1998; W u et al. 1998; D6zes et al. 1999). In a single field area, it is

possible to distinguish multiple episodes o f ductile or b r i t t l e - d u c t i l e deformation that span several million years (Burchfiel et al. 1992; Guillot et al. 1994; C o l e m a n 1996; Hodges et al. 1996; Searle et al. 2003; Searle & Godin 2003; Law et al.

CHANNEL FLOW-EXTRUSION HYPOTHESIS 2004). In most well-studied regions, there are no useful timing constraints on the youngest episodes of displacement, beyond the fact that the youngest strands of the STF cut and must post-date EarlyMiddle Miocene leucogranites (Hodges 1998). There are two notable exceptions to this conclusion. First, the Yadong cross-structure (c. 89~ in southern Tibet; Fig. 1), which has been described as a lateral ramp in the Himalayan thrust system (Wu et al. 1998), seems more easily interpreted as a left-slip tear fault in the STF system since: (1) apparently offset STF strands terminate into the principal bounding faults of the Yadong cross-structure (the Chomolhari fault system of Wu et al.); and (2) the Chomolhari fault system has not actually been mapped north or south of its intersection with the apparently offset strands (Wu et al. 1998). Inasmuch as the Chomolhari fault system displays surface evidence for Quaternary slip and is seismically active (Ni & Barazangi 1984; Ekstrom 1987), it stands to reason that the youngest strands of the STF system in the Yadong area have been active recently. Brittle strands of the STF system that may be similarly young have been described by several other researchers in the Tibet-Bhutan border region (e.g. Burchfiel et al. 1992; Edwards et al. 1996). Second, Hurtado et al. (2001) showed that the youngest strand of the STF system in the Kali Gandaki valley of central Nepal (c. 83~ Fig. 1) cuts the Quaternary Dangardzong fault, the principal growth structure for the Thakkhola graben, and thus must have experienced Pleistocene slip. This strand can be mapped westward to the Dhaulagiri region, where it offsets glacial moraines (Nakata 1989; Hurtado et al. 2001).

77

subsurface to a distance of at least 30 km north of its surface trace (Nelson et al. 1996). Projections farther north differ among different publications by the INDEPTH team. Makovsky et al. (1996) interpreted the STF as flattening to subhorizontal and extending northward for at least an additional 30 km at an approximate depth of 20 km beneath southern Tibet. In contrast, Hauck et al. (1998) felt that the STF reflector continued downward at approximately the same dip and could merge with the MHT about 60 km north of its surface trace and at a depth in excess of 30 km. One reason why such varied interpretations are possible is that many strong reflectors in the seismic record, including both the STF and the MHT, begin to lose their definition between 28~ and 28~ and are impossible to trace north of about 29~ This latitude coincides with the southern extent of the seismic bright spots described by Brown et al. (1996). A simple northward projection of the shallow STF reflector of Makovsky et al. (1996) and the MHT reflector of Nelson et al. (1996) defines a channel that is essentially coincident with the mid-crustal fluid zone as defined by an impressive suite of geophysical data (described above). Subsequent papers by Wu et al. (1998), Hodges et al. (2001), Beaumont et al. (2001, 2004), Hurtado et al. (2001), Grujic et al. (2002) and Jamieson et al. (2004) have expounded further on the tectonic implications of a connection between the extruded metamorphic core of the Himalaya and a weak mid-crustal channel beneath Tibet. I d e a 6: T h e r e is a l i n k b e t w e e n m o n s o o n driven erosion and channel extrusion along

I d e a 5: T h e G r e a t e r H i m a l a y a n roots northward

channel

into the T i b e t a n

middle crust

Given widespread evidence for high-temperature ductile deformation in the Himalayan metamorphic core and the northward dip of the STF system, it seems reasonable to postulate that the STF may be the upper surface of the mid-crustal fluid layer beneath southern Tibet, and that the Greater Himalayan Sequence is an exhumed portion of the fluid layer itself. The first articulation of this idea may be found in Nelson et al. (1996) and was based on the early results of Project INDEPTH. In the INDEPTH reflection seismic data-set, both the MHT and the basal STF structure are prominent north-dipping reflectors that appear to converge northward beneath southern Tibet because the STF reflector dips slightly more steeply than the MHT reflector (Zhao et al. 1993; Hauck et al. 1998). The STF reflector can be traced in the

the H i m a l a y a n f r o n t

As noted by Avouac & Burov (1996), focused erosion can have a dramatic influence on the dynamics of middle and lower crustal flow in convergent orogenic systems. Nelson et al (1996), Searle et al. (1997) and Wu et al. (1998) included rainfall along the Himalayan front in their cartoon depictions of Greater Himalayan extrusion, implicitly suggesting a link between climatic processes and channel extrusion. Hodges (1998) explicitly suggested that coordinated extrusion and erosion played an important role in the Miocene evolution of the Himalaya by dissipating excess gravitational potential energy accumulated during crustal thickening. A few years later, Hodges et al. (2001) argued that intense monsoon rainfall focused the flow direction of a mid-crustal channel toward the Himalayan range front. Independently, Beaumont et al. (2001) used thermalmechanical models to demonstrate how such focusing could be a natural consequence of

78

K.V. HODGES

erosion on the flanks of the Tibetan Plateau. Together, these two papers demonstrated how the Channel Flow-Extrusion model could provide an internally consistent explanation for a wide variety of geological and geophysical observations in the Himalaya and southern Tibet.

The Channel Flow-Extrusion hypothesis Out of these six ideas has emerged the hypothesis described in the following paragraphs. I focus on three phases of the evolution of the HimalayanTibetan orogenic system, during which the process of channelized mid-crustal flow had different tectonic manifestations (Fig. 3). P h a s e I: E s t a b l i s h m e n t

of a steady-state

c o n f i g u r a t i o n in the H i m a l a y a

(Early-

Middle Miocene)

Although the Himalayan-Tibetan orogenic system has developed over a period of roughly 50 million years since the initial stages of India-Eurasia collision, there was a fundamental change in its geodynamics in Early Miocene time (Hodges 2000). While some researchers persist in dating the Tibetan Plateau at c. 8Ma - see Molnar (2005) for a recent review of the arguments - I think that the preponderance of evidence suggests that crustal thickening beneath Tibet had progressed sufficiently to enable the development of the southern part of the plateau before 15 Ma (e.g. Turner et al. 1993, Coleman & Hodges 1995; Chung et al. 1998; Garzione et al. 2000a, b; Williams et al. 2001; Garzione et al. 2003; Spicer et al. 2003; Currie et al. 2005). As a consequence, the existence of a weak middle crust beneath the southern plateau can be postulated to date back to the Early Miocene interval, when the MCT system was the active frontal thrust in the Himalaya. At that time, there had developed a sufficient gradient in gravitational potential energy between the plateau and the Indian foreland - and presumably between the southern plateau and points farther north - to enable the lateral flow of the mid-crustal channel (cf. Hodges et al. 2001). Northward and eastward propagation of the channel - a process referred to as 'tunnelling' by Beaumont et al. (2001, 2004) - resulted in growth of the plateau without leaving substantial evidence in the surface geology of Tibet. The channelling process effectively decoupled the upper and middle-lower crust, setting the stage for subsequent independent upper crustal and middle-lower crustal responses to continued convergence between India and Eurasia (Royden et al. 1997; Chen et al. 2000; Clark & Royden 2000; Zhang et al. 2004).

A critical element of the Channel FlowExtrusion hypothesis is that sufficiently aggressive erosion was taking place along the southern flanks of the Himalaya to initiate the 'attraction' of the channel toward the range front in Early Miocene time. Indirect evidence of this having been the case comes from the sedimentary record in both the Bengal and Indus fans, which suggest a rapid flux of sediment from the Himalayan fiver systems in early Miocene time (e.g. Copeland & Harrison 1990; Amano & Taira 1992; Clift et al. 2004). Whether or not rapid erosion was driven then, as it is now, by the Indian monsoon remains an unanswered question. In any event, the hypothesis assumes that erosional removal of overburden along the range front was sufficient to allow the tunnelling channel to break toward the surface (Beaumont et al. 2004). The lower bound of this emergent channel was the recently developed MCT system whereas a new zone of shearing with opposing vergence - the STF system - developed to accommodate channel exhumation. Following Hodges et al. (1996), such features will be referred to in this paper as 'compensation structures' because they help mitigate the crustal thickness (and thus gravitational potential energy) contrast across the Himalaya. The top frame in Figure 3 illustrates the tectonic architecture of the Himalaya during Phase I. The hypothesis holds that the STF system can be viewed as the surface expression of the decoupling horizon between the upper and middle crust of southern Tibet. The MCT system is the surface expression of the MHT; beneath the Himalaya, the MHT serves as the lower boundary of the emergent channel. North of point A in Figure 3, the MHT and the lower channel boundary diverge as the Indian plate subducts northward. Between the downgoing plate and the channel, a mid-crustal duplex system (point B) forms as some Indian plate material is scraped off. The continued development of this 'accretionary complex' is crucial to the extrusion process. As the duplex system grows, material that was previously accreted moves northward and upward in the overriding plate, feeding material into the mid-crustal channel (cf. fig. 8 of Searle & Szulc 2005). This provides a constant source of material for recycling toward the erosion front and, because these materials have abnormally high concentrations of radioactive heat-producing elements, it ensures the maintenance of sufficiently high temperatures, and thus sufficiently low viscosities, to support channel flow (Huerta et al. 1999; Jamieson et al. 2004). Note that, while most of the accreted material is metasedimentary rock with high fluid contents, some are orthogneisses that are relatively anhydrous and thus can be entrained in the channel

CHANNEL FLOW-EXTRUSION HYPOTHESIS as rigid pods without the development of a significant metamorphic signature of the channel flow process. As will be seen in the subsequent section on Phase II, such material may retain important evidence of the nature of Eohimalayan metamorphism when eventually exhumed. During Phase I, near the southern margin of the plateau, a steady state is established locally in which the rate of accretion is roughly comparable to the rate of extrusion and the rate of erosion along the Himalayan front. In general, these rates can be traced by the rate of sedimentation in foreland deposits. In NW India, where deposits of the appropriate age have been most extensively studied, material from the extruding channel was being eroded and transported rapidly to the foreland until about 17 Ma, when erosion rates dropped dramatically in the hinterland and most of the detritus was instead sourced from the Lesser Himalayan Sequence farther south (White et al. 2002b). Farther east, in western Nepal, this shift occurred somewhat later (c. 12-10 Ma; Huyghe et al. 2001). In general, it seems that the rate of erosion of Greater Himalayan Sequence rocks dropped significantly across much of the Himalayan front in Middle Miocene or earliest Late Miocene time, although the Bengal Fan detrital record suggests that at least some exposures of the Greater Himalayan Sequence have been eroding and supplying sediment to the fan throughout the Middle Miocene-Recent interval (France-Lanord et al. 1993).

P h a s e II: E s t a b l i s h m e n t o f a s e c o n d e m e r g e n t c h a n n e l north o f the H i m a l a y a n crest a n d e a s t w a r d g r o w t h o f the Tibetan P l a t e a u (Middle M i o c e n e - E a r l y Pliocene)

In the context of the Channel Flow-Extrusion hypothesis, the Middle-Late Miocene decrease in erosion rate of the Greater Himalayan Sequence was contemporaneous with a dramatic slowing of channel extrusion. Large-magnitude slip along the STF and MCT systems had effectively ceased by c. 16Ma throughout much of the Himalaya, although research in the region of southern Tibet north of Bhutan suggests that Phase I STF activity persisted there into the Late Miocene (Edwards & Harrison 1997; Wu et al. 1998). Much of the shortening that had taken place on the MCT system was transferred in sequence toward the foreland to the structurally lower LHT system (DeCelles et aL 2001), although some out-of-sequence thrusts have been identified within the Greater Himalayan Sequence (Brun et al. 1985; Grujic et al. 1996, 2002; Hodges et al. 1996; Vannay & Hodges

79

1996; Searle 1999). During the limited extrusion that did occur at the Himalayan front over the Middle Miocene-Early Pliocene interval, the surface expressions of the boundaries of the extruding channel were active structures of the STF system and, successively, the LHT and MBT systems (Fig. 3). But more important extrusion processes were taking place north of the Himalayan range crest. The thermal-mechanical models of Beaumont et al. (2001, 2004) suggest that any mechanism that reduces overburden above a mid-crustal channel, including extensional denudation, will tend to divert channel flow toward the area of denudation. Beaumont and co-workers suggested that upper crustal extension in the region between the Indus-Tsangpo Suture zone and the STF system might be responsible for the emergence of channel material in the cores of the North Himalayan gneiss domes, a belt of metamorphic culminations that stretches across southern Tibet (Burg et al. 1984b; Hodges 2000; Watts et al. 2005). Figure 4 illustrates how this process changed the tectonic architecture of southernmost Tibet during Phase II. I envision upper-crustal thinning occurring along detachment faults similar to those responsible for large-magnitude extension in the Basin and Range province of the North American Cordillera (Wernicke 1985). In this setting, such detachments root into the mid-crustal channel itself, providing an upper boundary for a new subsidiary channel that will be referred to here as the North Himalayan channel. As the channel front ascends toward the surface, the low effective elastic thickness of the Tibetan upper crust (Masek et al. 1994) permits an isostatic response in the detachment footwall (Wernicke & Axen 1988) and the development of a rolling hinge geometry (Wdowinski & Axen 1992; Axen & Bartley 1997). From this perspective, the mylonitic carapaces of the gneiss domes in southern Tibet - as described by Chen et al. (1990) and Lee et al. (2000, 2004) - are surface expressions of compensation structures related to emergence of the North Himalayan channel. The INDEPTH seismic reflection line captured the subsurface reflections of the south-dipping carapace on the south side of one of the gneiss domes near Kangmar (c. 28~ 89~ Fig. 3). As depicted by Nelson et al. (1996) and Hauck et al. (1998), the shallowly dipping carapace reflectors can be traced downward to a depth of slightly less than 10 km before they are lost. Hauck et al. (1998) inferred it to be an old strand of the STF system that had been warped up and over the Kangmar dome as a consequence of the development of an underlying ramp in the MHT. Such a ramp is highly speculative - there is no sign of it on the INDEPTH reflection profiles.

80

K.V. HODGES

Fig, 4. Conceptual model of the development of a North Himalayan gneiss dome by extrusion enabled by uppercrustal extension (cf. Beaumont et al. 2001). Darker-grey shading indicates upper crust. Random dash pattern indicates channel material; light-grey overlay shading designates less active parts of the channel. Half-arrows indicate slip on individual faults. Thin, dashed lines represent mylonite zones. Large, freeform arrows indicate large-scale kinematics; darker arrows designate more rapid movement. (a) Upper crustal extension initiates on a major detachment that roots into the top of the mid-crustal channel. (b) Material flows upward in the footwall of the detachment and a rolling-hinge geometry develops (Buck 1988; Wernicke & Axen 1988), rotating the surface expression of the detachment to a lower dip angle. (c) With progressive extrusion, the mylonitic carapace of the gneiss dome emerges. (d) Ultimately, a new, subsidiary channel forms, and the original channel becomes less active. Moreover, the Hauck e t al. interpretation is inconsistent with two important sets of observations made by Lee et al. (2000) in the Kangmar dome: (1) deformational fabrics in the carapace imply top-to-the-north hanging wall transport on the

northern margin of the dome, top-to-the-south transport on the southern margin, and a large flattening strain c o m p o n e n t at the dome crest; and (2) thermochronologic data from footwall metamorphic rocks indicate exhumation through

CHANNEL FLOW-EXTRUSION HYPOTHESIS medium- and low-temperature closure isotherms during the Middle to Late Miocene interval, significantly later than most STF footwall rocks exposed farther south. In my view, a better interpretation is that these mylonites developed independently of the STF system as shown schematically in Figure 4, and that renewed extrusion of the main channel during Phase III (see below) truncated the older mylonitic fabrics of the North Himalayan channel. In some cases, the cores of the North Himalayan domes preserve an important record of earlier events despite their involvement in Phase I and Phase II deformation. For example, the Tso Morari complex in NW India (c. 33~ 78~ Fig. 3) is unique in that it contains only one of two known exposures of ultrahigh-pressure (coesite-bearing) mineral assemblages in the Himalaya (Sachan et al. 2004). The coesite-bearing mafic eclogites occur as small, isolated lenses in an orthogneiss matrix that represents Indian continental basement, such that the ultrahigh-pressure assemblages record subduction of Indian continental crust during the early stages of collision in Eocene (Eohimalayan) time (de Sigoyer et al. 2004; Leech et al. 2005). Structural and thermochronologic data suggest that these rocks had been exhumed to mid-crustal levels, equivalent to about 15 km depth, by Early Oligocene time (de Sigoyer et al. 2000), but that the final stages of exhumation to the surface did not occur until Middle to Late Miocene time (Schlup et al. 2003). I infer that the ultrahigh-pressure core of the dome represents a rigid pod of Indian basement that was accreted to the overriding Eurasian plate early in the collisional process, extruded to mid-crustal levels by a mechanism related to the buoyant force (e.g. Chemenda et al. 1996), eventually incorporated into the main Tibetan channel, and finally exhumed as part of the North Himalayan channel during Phase II. If this model is correct, the lack of an Early Miocene metamorphic overprint in these rocks, even though they were entrained in a high-temperature channel, may have been a consequence of low fluid contents in the basement orthogneisses. Another important Phase II phenomenon was continued eastward tunnelling of the Tibetan midcrustal channel. By Middle-Late Miocene time, even the eastern margin of the plateau had reached high elevations due to the propagation of the flow front (Kirby et al. 2002; Clark et al. 2005).

P h a s e III: I n t e n s i v e e x t r u s i o n a t t h e Himalayan front (Late Pliocene-Recent)

Medium-temperature thermochronometers in bedrock samples from the Greater Himalayan Sequence

81

typically record Middle Miocene-Early Pliocene cooling ages. Recent efforts to constrain exhumation rates over that interval from the study of detrital micas in central Nepal (Ruhl & Hodges 2005) suggest values of less than 1 mm a -1. However, fission-track and (U-Th)/He data from the Greater Himalayan Sequence (e.g. Sorkhabi et al. 1996; Arita & Ganzawa 1997; Burbank et al. 2003; Theide et al. 2004; Vannay et al. 2004) suggest a transition to much more rapid exhumation rates in Late Pliocene time (Ruhl et al. 2005). Potentially related to climate destabilization (Molnar 2004), this increase in the erosion rate reinvigorated extrusion of the Tibetan mid-crustal channel toward the Himalayan range front and the process continues today (Fig. 3). During this phase, the base of the channel has been the active, surface-breaking strand of the MHT. For much of the time, this may have been the Main Frontal thrust system (Hodges 2000; Lav~ & Avouac 2000). In some places and at some times, however, the active channel base appears to have been out-of sequence thrust faults. Some of these faults are probably reactivated older strands of the MCT or LHT systems (Hodges et al. 2004); others are newly developed structures (Wobus et al. 2003, 2005; Theide et aL 2004; Vannay et al. 2004). The upper boundary of the extruding channel has remained the surface trace of the STF system during Phase Ill, but the specific structures are different from those that were active during Phase I. These youngest faults of the STF system are characterized by brittle fault rocks and, in some cases, show direct evidence of Quaternary activity (Hurtado et al. 2001). In some areas - like the Sutlej Valley of NW India (c. 31~176 78~ Fig. 3) - the upper boundary of the extruded zone is a normalsense shear zone, exposed south of the trace of the STF system, which represents a newly developed compensation structure (Theide et al. 2004; Vannay et al. 2004). Locations of channel boundaries are probably governed by summer monsoon precipitation patterns. In the Marsiyandi Valley (c. 28~ 84~ where Quaternary thrust faults around the traditional MCT trace serve as the lower boundary (Hodges et aL 2004) and the Phu detachment strand of the STF system serves as the upper boundary (Searle & Godin 2003), recent meteorological studies (Barros et al. 2000; Barros & Lang 2003) demonstrate a strong north-south gradient in precipitation across the Himalayan front: the heaviest precipitation falls exclusively between the active thrusts and the Phu detachment (Hodges et al. 2004). In the Sutlej Valley, Theide et al. (2004) have found precisely the same relationship (based on SSM/I passive microwave satellite precipitation data), but the precipitation

82

K.V. HODGES

maximum falls between a Quaternary out-ofsequence thrust fault roughly 30 km south of the trace of the MCT system and zone of Quaternary normal faults approximately coincident with the MCT trace. Elsewhere in the system, geomorphic studies of the evolution of river systems in southeastern Tibet and Yunnan (Clark et al. 2004; Schoenbohm et al. 2004), coupled with cosmogenic nuclide exposure age dating (Ouimet et al. 2005), testify to the continued southeastward growth of the Tibetan Plateau during Phase III.

Testing the hypothesis A hypothesis that cannot be tested is without scientific value, and the Channel Flow-Extrusion hypothesis is no exception. Because, as reviewed here, the hypothesis was developed specifically to explain basic geological and geophysical observations in southern Tibet and the Himalaya, it would be circular reasoning to regard consistency with observations in this particular orogenic system as some sort of 'confirmation' of the viability of the hypothesis. However, three kinds of research might be useful ways to proceed. Theoretical studies

One class of studies might involve attempts to understand whether or not the physics of mountain building actually demand the development of several phenomena that are crucial to the hypothesis. For example, does the thermal structure of overthickened crust in an orogenic system require the existence of a discrete fluid channel in the middle-lower crust, or are other rheologies plausible? Is it inevitable that such a channel would tunnel laterally through more rigid crust? If such tunnelling occurs, under what conditions is it safe to infer that surface erosion will influence the direction of flow, and under what conditions will a channel be forced to flow upward toward the surface? As impressive as numerical experiments such as those of Beaumont et al. (2001, 2004) are, it is important to remember that complex numerical models can be highly idiosyncratic, and cause and effect are not always obvious in them. Many opportunities exist to refine existing theoretical models of channel flow and to study their behaviours, and particularly the sensitivity of those behaviours to alternative choices of starting conditions. The goal of future studies of this sort should be to show that the essential characteristics of the Channel Flow-Extrusion hypothesis are not only viable but are predictable consequences of the building of orogenic plateaus.

Tests o f p r e d i c t i o n s in the H i m a l a y a

The Channel Flow-Extrusion hypothesis makes several testable predictions that should encourage the design of new research campaigns. Most importantly, it requires the existence of a Tibetan Plateau - or at least a 'proto-Plateau' in southern Tibet - as far back as Early Miocene time. Although I am persuaded by the arguments for high elevations - and presumably thick crust - in southern Tibet based on stable isotope and leaf physiognomy records (Garzione et al. 2000a, b; Rowley et al. 2001; Spicer et al. 2003), this issue is far from settled and requires further study. Moreover, the hypothesis requires progressive northward and eastward growth of the Tibetan Plateau from Early Miocene time onward, not wholesale uplift of the plateau at c. 8 Ma (Molnar et al. 1993). This component of the hypothesis should be testable by expanding studies of palaeo-elevation proxies outside of southern Tibet. One weakness in the logic behind the hypothesis is that the field evidence for extrusion at the Himalayan front is strongest for the Early Miocene interval (Phase I), yet the argument that high precipitation at the range front attracts the channel to the surface is based on the modern pattern of monsoon rainfall (Phase III). Thus, one useful test of the hypothesis would be to search for evidence of high range front precipitation in Early Miocene time. Perhaps the most fruitful avenue for such work involves a more comprehensive study of the sedimentary record in the oldest foreland sequences as well as the Bengal and Indus fans. The idea that the North Himalayan gneiss domes are related to a Phase II subsidiary channel is largely speculative. While it is consistent with what is known about the Kangmar and Mabja domes (Lee et al. 2000, 2004), many other North Himalayan domes have not been studied, and the hypothesis makes testable predictions about what the strain patterns and timing relationships should be in these metamorphic culminations. Perhaps the most controversial aspects of the hypothesis concern whether or not Phase III extrusion continues today. Fortunately, these aspects are also the most testable. Since there is little dispute about the existence of active faults at the range front, and since these faults could serve as the lower boundary of the extruding channel, it is not necessary for there to be active, surfacebreaking faults near the base of the Greater Himalayan Sequence for the hypothesis to be correct. However, a close link between precipitation patterns and extrusion - as has been postulated here and in Hodges et al. (2001) and Beaumont et al. (2001) - would favour a more northerly placement of the lower channel boundary than the

CHANNEL FLOW-EXTRUSION HYPOTHESIS MFT trace. Studies are ongoing to determine whether or not active, out-of-sequence thrusts persist along-strike in areas other than central Nepal (Hodges et al. 2004; Wobus et aI. 2005) and the Sutlej Valley (Theide et al. 2004; Vannay et al. 2004). Phase III extrusion also requires recently active slip on the STF system or on some newly formed compensation structures at deeper or shallower structural levels. None of the available geological data are inconsistent with this requirement, but neotectonic studies of candidate structures are few (Hurtado et al. 2001; Vannay et al. 2004). To test this part of the hypothesis, we particularly need more focused studies of the youngest, brittle structures of the STF system. The existence of these features has been known for years (Burchfiel et al. 1992), but they have received far less research attention than the older, more ductile strands of the STF system. Finally, if the extrusion front truly follows zones of high erosion, alongstrike variations in monsoon precipitation intensity (Bookhagen et al. 2005) should correlate with the aggressiveness of the extrusion process. Thus, the patterns of Quaternary deformation should vary in a predictable way along strike in the Himalaya.

Tests o f the applicability o f the h y p o t h e s i s in o t h e r o r o g e n i c s y s t e m s

Even if the Channel Flow-Extrusion hypothesis is valid for the Himalaya and Tibet, its general usefulness for understanding collisional orogenesis remains unclear. In addition to several papers in this volume, the geological literature contains numerous examples of how geological and geophysical observations in other convergent orogenic systems, especially those with continental plateaus (e.g. the Puna-Altiplano and the Colorado Plateau) might be explained by lower or middle crustal flow (e.g. Bird 1991; Hodges & Walker 1992; Lamb & Hoke 1997; McQuarrie & Chase 2000; Gerbault & Willingshofer 2004; Gerbault et al. 2005). In each case, the evidence is consistent with the Channel Flow-Extrusion model but could be equally consistent with other models that do not invoke large-scale crustal flow. The great challenge before us is to identify unambiguous indicators of such flow in the geological record. Some of the best places to search for such indicators may be more ancient orogenic systems that offer regionally extensive exposures of the middle and lower crust. One example that has received some attention already is the Grenville orogenic system of North America (Culshaw et al. 1997; Jamieson et al. 2002), and another - the East Greenland Caledonides-includes especially dramatic exposures of medium-to-high-pressure metasedimentary rocks,

83

with abundant leucogranitic plutons, that may be our best available analogues for the fluid crustal channel beneath Tibet (Hartz et al. 2001; Strachan & Martin 2001; White et al. 2002a; McClelland & Gilotti 2003; Higgins et al. 2004; Gilotti & McClelland 2005).

Conclusions The Channel Flow-Extrusion hypothesis described here offers a useful conceptual framework for understanding the Early Miocene- Recent tectonics of the Himalayan-Tibetan orogenic system. In the hypothesis, the development of the Tibetan Plateau had a strong influence on the evolution of the system, altering it from what it might have been if the only forcing factor had been India-Eurasia plate convergence. Lateral tunnelling of a fluid Tibetan middle crust under the influence of gravity is postulated to have accommodated the eastward growth of the plateau without leaving a structural imprint on the surface geology. Along the southern plateau margin, aggressive erosion is thought to have coordinated with extrusion of the mid-crustal channel to produce a highly effective means of dissipating excess gravitational potential energy stored in overthickened Tibetan crust. According to the hypothesis, Middle Miocene upper crustal extension in southern Tibet enabled the development of a subsidiary channel that resulted in the formation of the North Himalayan gneiss domes. Many components of the hypothesis remain untested, and the design and implementation of such tests should be a top priority for students of Himalayan tectonics. The perspectives on Himalayan geology presented in this paper have benefited greatly from interactions in recent years with D. Burbank, T. Ehlers, A. Heimsath, K. Huntington, J. Hurtado, K. Viskupic, D. Whipp, K. Whipple and C. Wobus. The late Doug Nelson deserves special thanks for insisting that there must be a relationship between a fluid lower crust beneath Tibet and the tectonic evolution of the Himalaya, and for challenging me to figure out what that relationship must be. (I'm working on it, Doug... ). J.-P. Avouac, S. Klemperer, R. Law and M. Searle provided insightful reviews that proved very useful in producing the final version of this manuscript. My work in the Himalaya has been supported most recently by grants from the Tectonics, Petrology and Geochemistry, and Continental Dynamics programs of the National Science Foundation (EAR-9814497, EAR-9909426, EAR-0087508, EAR-0125870).

References ALLMENDINGER, R. W., JORDAN, T. E., KAY, S. M. & ISACKS, B. L. 1997. The evolution of the

84

K.V. HODGES

Altiplano-Puna plateau of the central Andes. Annual Review of Earth and Planetary Sciences, 25, 139-174. ALSDORF, D. & NELSON, D. 1999. Tibetan satellite magnetic low: Evidence for widespread melt in the Tibetan crust? Geology, 27, 943-946. AMANO, K. & TAIRA, A. 1992. A two-phase uplift of Higher Himalayas since 17 Ma. Geology, 20, 391-394. ARITA, K. & GANZAWA,Y. 1997. Thrust tectonics and uplift processes of the Nepal Himalaya revealed from fission-track ages. Journal of Geography (Japan), 106, 156-167. ARTYUSHKOV, E. V. 1973. Stresses in the lithosphere caused by crustal thickness inhomogeneities. Journal of Geophysical Research, 78, 7675-7708. AVOUAC, J. P. & BUROV, E. B. 1996. Erosion as a driving mechanism of intracontinental mountain growth. Journal of Geophysical Research, 101, 17,747-17,769. AXEN, G. J. & BARTLEY, J. M. 1997. Field tests of the rolling hinge hypothesis. Journal of Geophysical Research, 102, 20,515-20,537. BARROS, A. P. & LANG, T. J. 2003. Monitoring the monsoon in the Himalayas: Observations in central Nepal, June 2001. Monthly Weather Review, 131, 1408-1427. BARROS, A. P., JOSHI, M., PUTKONEN, J. & BURBANK, D. W. 2000. A study of the 1999 monsoon rainfall in a mountainous region in Central Nepal using TRMM products and raingauge observations. Geophysical Research Letters, 27, 3683-3686. BEAUMONT, C., JAM1ESON, R. A., NGUYEN, M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & MEDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with implications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109. DOI: 10.1029/ 2003JB002809. BraD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10,27510,286. BOOKHAGEN, B., THEIDE, R. C. & STRECKER, M. R. 2005. Abnormal monsoon years and their control on erosion and sediment flux in the high, arid northwest Himalaya. Earth and Planetary Science Letters, 231, 131-146. BRACE, W. F. 8z KOHLSTEDT, D. L. 1980. Limits on lithosphefic stress imposed by laboratory experiments. Journal of Geophysical Research, 85, 6248 -6252. BROWN, L. D., ZHAO, W., NELSON, K. D. ErAL. 1996. Bright spots, structure and magmatism in southern Tibet from INDEPTH seismic reflection profiling. Science, 274, 1688-1690. BRUN, J.-P., BURG, J.-P. & CHEN, G. M. 1985. Strain trajectories above the Main Central Thrust (Himalaya) in southern Tibet. Nature, 313, 388-390.

BUCK, W. R. 1988. Flexural rotation of normal faults. Tectonics, 7, 959-974. BURBANK, D. W., BLYTHE, A. E., PUTKONEN, J. K., PRATT-SITAULA, B. A., GABET, E. J., OSKIN, M. E., BARROS, A. P. & OJHA, T. P. 2003. Decoupiing of erosion and precipitation in the Himalaya. Nature, 426, 652-655. BURCHFIEL, B. C. & ROYDEN, L. H. 1985. North-south extension within the convergent Himalayan region. Geology, 13, 679-682. BURCHFIEL, B. C., CHEN, Z., HODGES, K. V., LIU, Y., ROYDEN, L. H., DUNG, C. & Xu, J. 1992. The South Tibetan Detachment System, Himalayan Orogen: Extension Contemporaneous With and Parallel to Shortening in a Collisional Mountain Belt. Geological Society of America, Special Papers, 269. BURCHFIEL, B. C., CHEN, Z., LIu, Y. & ROYDEN,L. H. 1995. Tectonics of the Longmen Shan and adjacent regions, central China. International Geology Review, 37, 661-735. BURG, J. P., BRUNEL, M., GAPAIS, D., CHEN, G. M. LIu, G. H. 1984a. Deformation of leucogranites of the crystalline Main Central Sheet in southern Tibet (China). Journal of Structural Geology, 6, 535 -542. BURG, J. P., GUIRAUD, M., CHEN, G. M. ~z LI, G. C. 1984b. Himalayan metamorphism and deformations in the North Himalayan belt, southern Tibet, China. Earth and Planetary Science Letters, 69, 391-400. BURG, J.-P., DAVY, P. & MARTINOD,J. 1994. Shortening and analogue models of the continental lithosphere: New hypothesis for the formation of the Tibetan plateau. Tectonics, 13, 475-483. CABY, R., PI~CHER,A. & LEFORT, P. 1983. Le grande chevauchement central himalayen: nouvelles donnres sur le m&amorphisme inverse ~ la base de la Dalle du Tibet. Revue de Gdographie physiquean et Gdologie dynamique 24, 89-100. CHEMENDA, A. I., MATTAUER, M. & BOKUN, A. N. 1996. Continental subduction and a mechanism for exhumation of high-pressure metamorphic rocks: new modelling and field data from Oman. Earth and Planetary Science Letters, 143, 173-182. CHEN, L., BOOKER, J. L., JONES, A. G., Wu, N., UNSWORTH, M. J., WEI, W. & TAN, H. 1996. Electrically conductive crust in southern Tibet from INDEPTH magnetotelluric surveying. Science, 274, 1694 - 1696. CHEN, Z., LIU, Y., HODGES, K. V., BURCHFIEL,B. C., ROYDEN, L. H. & DUNG, C. 1990. The Kangmar Dome - a metamorphic core complex in southern Xizang (Tibet). Science, 250, 1552-1556. CHEN, Z., BURCHFIEL, B. C., LIU, Y. ET AL. 2000. Global positioning system measurements from eastern Tibet and their implications for India/ Eurasia intercontinental deformation. Journal of Geophysical Research, B, Solid Earth and Planets, 105, 16,215-16,227. CHUNG, S.-L., LO, C.-H., LEE, T.-Y., ZHANG, Y., XIE, Y., LI, X., WANG, K.-L. & WANG, P.-L. 1998. Diachronous uplift of the Tibetan plateau starting 40 Myr ago. Nature, 394, 769-773.

CHANNEL FLOW-EXTRUSION HYPOTHESIS CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. CLARK, M. K., ROYDEN, L. H., WHIPPLE, K. X. ETAL. 2004. Surface uplift, tectonics, and erosion of eastern Tibet from large-scale drainage patterns. Tectonics, 23. DOI: 10.1029/2002TC001402. CLARK, M. K., HOUSE, M. A., ROYDEN, L. H., WHIPPLE, K. X., BURCHFIEL, B. C., ZHANG, X. & TANG, W. 2005. Late Cenozoic uplift of southeastem Tibet. Geology, 33, 525-528. CLIFT, P. D., CAMPBELL,I. H., PRINGLE,M. S., CARTER, A., ZHANG, X., HODGES, K. V., KHAN, A. A. & ALLEN, C. M. 2004. Thermochronology of the modem Indus River bedload: New insight into the controls on the marine stratigraphic record. Tectonics, 23. DOI: 10.1029/2003TC001559. COLEMAN, M. E. 1996. Orogen-parallel and orogenperpendicular extension in the central Nepalese Himalayas. Geological Society of America Bulletin, 108, 1594-1607. COLEMAN, M. E. 1998. U-Pb constraints on OligoceneMiocene deformation and anatexis within the central Himalaya, Marsyandi valley, Nepal. American Journal of Science, 298, 553-571. COLEMAN, M. & HODGES, K. 1995. Evidence for Tibetan Plateau uplift before 14 Myr ago from a new minimum age for east-west extension. Nature, 374, 49-52. COPELAND, P. & HARRISON, T. M. 1990. Episodic rapid uplift in the Himalaya revealed by 4~ analysis of detrital K-feldspar and muscovite, Bengal Fan. Geology, 18, 354-357. COTTE, N., PEDERSEN,M., CAMPILLO,J., MARS, J. C., NI, J. F., KIND, R., SANDVOL, E. & ZHAO, W. J. 1999. Determination of the crustal structure in southern Tibet by dispersion and amplitude analysis of Rayleigh waves. Geophysical Journal International, 138, 809-819. CULSHAW,N. G., JAMIESON,R. A., KETCHUM,J. W. F., WODICKA, N., CORRIGAN, D. & REYNOLDS, P. H. 1997. Transect across the northwestern Grenville orogen, Georgian Bay, Ontario: Polystage convergence and extension in the lower orogenic crust. Tectonics, 16, 966-982. CURRIE, B. S., ROWLEY, D. B. & TABOR, N. J. 2005. Middle Miocene paleoaltimetry of southern Tibet: Implications for the role of mantle thickening and delamination in the Himalayan orogen. Geology, 33, 181-184. DE SIGOYER, J., CHAVAGNAC,V., BLICHERT-TOFT, J., VILLA, I. M., LUAIS, B., GUILLOT, S., COSCA, M. & MASCLE, G. 2000. Dating the Indian continental subduction and collisional thickening in the Northwest Himalaya; multichronology of the Tso Morari eclogites. Geology, 28, 487-490. DE SIGOYER, J., GUILLOT, S. & DICK, P. 2004. Exhumation of the ultrahigh-pressure Tso Morari unit in eastern Ladakh (NW Himalaya): A case study. Tectonics, 23. DOI: 10.1029/2002TC 001492. DECELLES, P. G., ROBINSON, D. M., QUADE, J., OJHA, T. P., GARZIONE, C. N., COPELAND, P. & UPRETI, B. N. 2001. Stratigraphy, structure, and tectonic

85

evolution of the Himalayan fold-thrust belt in western Nepal. Tectonics, 20, 487-509. DEWEY, J. F. & BURKE, K. C. A. 1973. Tibetan, Variscan and Precambrian reactivation: Products of continental collision. Journal of Geology, 81, 683 -692. DI~ZES, P., VANNAY, J. C., STECK, A., BussY, F. & COSCA, M. 1999. Synorogenic extension: quantitative constraints on the age and throw of the Zanskar shear zone (NW Himalaya). Geological Society of America Bulletin, 111,364-374. EDWARDS, M. A. & HARRISON, T. M. 1997. When did the roof collapse? Late Miocene north-south extension in the high Himalaya revealed by ThPb monazite dating of the Khula Kangfi granite. Geology, 25, 543-546. EDWARDS, M. A., KIDD, W. S. F., LI, J., YUE, Y. & CLARK, M. 1996. Multi-stage development of the southern Tibetan detachment system near Khula Kangri. New data from Gonto La. Tectonophysics, 260, 1-19. EKSTROM, G. A. 1987. A broad band method of earthquake analysis. PhD thesis, Harvard University. ENGLAND, P. C. & HOUSEMAN,G. A. 1988. The mechanics of the Tibetan Plateau. In: SHACKLETON, R. M., DEWEY, J. F. & WINDLEY, B. F. (eds)

Tectonic Evolution of the Himalayas and Tibet. Philosophical Transactions of the Royal Society of London, Series A, 326, 302-319. FAN, G.-W. & LAY, T. 2003. Strong Lg attenuation in the Tibetan Plateau. Bulletin of the Seismological Society of America, 93, 2264-2272. FIELDING, E. J., ISACKS, B. L., BARAZANGI, M. & DUNCAN, C. 1994. How flat is Tibet? Geology, 22, 163-167. FRANCE-LANORD, C., DERRY, L. & MICHARD, A. 1993. Evolution of the Himalaya since Miocene time: Isotopic and sedimentologic evidence from the Bengal Fan. In: TRELOAR, P. J. & SEARLE, M. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 605-621. FRANCHETEAU, J., JAUPART, C., XIAN, J. S., WEN-HUA, K., DE-LU, L., JIA-CHA, B., HUN-PIN, W. & HSIA-YEU, D. 1984. High heat flow in southern Tibet. Nature, 307, 32-36. GAILLARD, F., SCAILLET, B. & PICHAVANT,M. 2004. Evidence for present-day leucogranite pluton growth in Tibet. Geology, 32, 801-804. GARZIONE, C. N., DETTMAN, D. L., QUADE, J., DECELLES, P. G. & BUTLER, R. F. 2000a. High times on the Tibetan Plateau: Paleoelevation of the Thakkhola Graben, Nepal. Geology, 28, 339-342. GARZIONE, C. N., QUADE, J., DECELLES, P. G. & ENGLISH, N. B. 2000b. Predicting paleoelevation of Tibet and the Himalaya from 6180 vs. altitude gradients of meteoric waters across the Nepal Himalaya. Earth and Planetary Science Letters, 183, 215-229. GARZIONE, C. N., DECELLES, P. G., HODKINSON, D. G., OJHA, T. & UPRETI, B. N. 2003. Eastwest extension and Miocene environmental change in the southern Tibetan Plateau: Thakkhola

86

K.V. HODGES graben, central Nepal. Geological Society of America Bulletin, 115, 3-20.

GERBAULT, M. & WILLINGSHOFER, E. 2004. Lower crustal indentation or horizontal ductile flow during continental collision? Tectonophysics, 387, 169-187. GERBAULT, M., MARTINOD, J. & HERAIL, G. 2005. Possibly orogeny-parallel lower crustal flow and thickening in the central Andes. Tectonophysics, 399, 59-72. GILOTTI, J. A. & MCCLELLAND, W. C. 2005. Leucogranites and the time of extension in the East Greenland Caledonides. Journal of Geology, 113, 399-417. GRASEMANN, B., FRITZ, H. & VANNAY, J.-C. 1999. Quantitative kinematic flow analysis from the Main Central Thrust Zone (NW-Himalaya, India): implications for a decelerating strain path and the extrusion of orogenic wedges. Journal of Structural Geology, 21, 837-853. GRATTON, J. 1989. Crustal shortening, root spreading, isostasy, and the growth of orgenic belts: A dimensional analysis. Journal of Geophysical Research, 94, 15,627-15,634. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KONDIG, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. GRUJIC, D., HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel; insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. GUILLOT, S., HODGES, K. V., LE FORT, P. & P~CHER, A. 1994. New constraints on the age of the Manaslu leucogranite: Evidence for episodic tectonic denudation in the central Himalayas. Geology, 22, 559-562. HACKER, B. R., GNOS, E., RATSCHBACHER, L., GROVE, M., MCWILLIAMS, M., SOBOLEV, S. V., WAN, J. & ZHENHAN, W. 2000. Hot and dry deep crustal xenoliths from Tibet. Science, 287, 2463 -2466. HARTZ, E. H., ANDRESEN, A., HODGES, K. V. & MARTIN, M. W. 2001. Syncontractional extension and exhumation of deep crustal rocks in the east Greenland Caledonides. Tectonics, 20, 58-77. HAUCK, M. L., NELSON, K. D., BROWN, L. D., ZHAO, W. & ROSS, A. R. 1998. Crustal structure of the Himalayan Orogen at approximately 90 ~ east longitude fi'om Project INDEPTH deep reflection profiles. Tectonics, 17, 481-500. HEIM, A. & GANSSER, A. 1939. Central Himalaya Geological observations of Swiss expedition, 1936. Memoire, Societe Helvetique Science Naturelle, 73, 1-245. HERREN, E. 1987. Zanskar shear zone: Northeastsouthwest extension within the Higher Himalayas (Ladakh, India). Geology, 15, 409-413. HIGGINS, A. K., ELVEVOLD, S., ESCHER, J. C. ETAL. 2004. The foreland-propagating thrust architecture of the East Greenland Caledonides 72~176 Journal of the Geological Society, London, 161, 1009-1026.

HODGES, K. V. 1998. The thermodynamics of Himalayan orogenesis. In: TRELOAR, P. J. & O'BRIEN, P. (eds) What Drives Metamorphism and Metamorphic Reactions? Geological Society, London, Special Publications, 138, 7-22. HODGES, K. V. 2000. Tectonics of the Himalaya and southern Tibet from two perspectives. Geological Society of America Bulletin, 112, 324-350. HODGES, K. V. & WALKER, J. D. 1992. Extension in the Cretaceous Sevier orogen, North American Cordillera. Geological Society of America Bulletin, 104, 560-569. HODGES, K. V., PARRISH, R., HOUSH, T., Lux, D., BURCHFIEL, B. C., ROYDEN, L. & CHEN, Z. 1992. Simultaneous Miocene extension and shortening in the Himalayan orogen. Science, 258, 14661470. HODGES, K. V., BURCHFIEL, B. C., ROYDEN, L. H., CHEN, Z. & LIU, Y. 1993. The metamorphic signature of contemporaneous extension and shortening in the central Himalayan orogen: Data from the Nyalmn transect, southern Tibet. Journal of Metamorphic Geology, 11, 721-737. HODGES, K. V., PARRISH, R. R. & SEARLE,M. P. 1996. Tectonic evolution of the central Annapurna Range, Nepalese Himalayas. Tectonics, 15, 1264-1291. HODGES, K., BOWRING, S., DAVIDEK, K., HAWKINS, D. & KROL, M. 1998. Evidence for rapid displacement on Himalayan normal faults and the importance of tectonic denudation in the evolution of mountain ranges. Geology, 26, 483-486. HODGES, K. V., HURTADO, J. M. & WHIPPLE, K. X. 2001. Southward extrusion of Tibetan crust and its effect on Himalayan tectonics. Tectonics, 20, 799-809. HODGES, K., WOBUS, C., RUHL, K., SCHILDGEN,T. & WHIPPLE, K. 2004. Quaternary deformation, river steepening, and heavy precipitation at the front of the Higher Himalayan ranges. Earth and Planetary Science Letters, 220, 379-389. HOLLAND, T. & POWELL, R. 2001. Calculation of phase relations involving haplogranitic melts using an internally consistent thermodynamic dataset. Journal of Petrology, 42, 673-683. HUERTA, A. D., ROYDEN, L. H. & HODGES, K. V. 1999. The effects of accretion, erosion and radiogenic heat on the metamorphic evolution of collisional orogens. Journal of Metamorphic Geology, 17, 349-366. HURTADO, J. M., HODGES, K. V. & WHIPPLE, K. X. 2001. Neotectonics of the Thakkhola Graben and implications for Recent activity on the South Tibetan Fault System in the central Nepalese Himalaya. Geological Society of America Bulletin, 113, 222-240. HUYGHE, P., GALY, A., MUGN1ER, J.-L. & FRANCELANORD, C. 2001. Propagation of the thrust system and erosion in the Lesser Himalaya: Geochemical and sedimentological evidence. Geology, 29, 1007-1010. ISACKS, B. L. 1988. Uplift of the central Andean plateau and bending of the Bolivian orocline. Journal of Geophysical Research, 93, 3211 - 3231.

CHANNEL FLOW-EXTRUSION HYPOTHESIS JACKSON, J. A., AUSTRHEIM, H., MCKENZm, D. & PRIESTLEY, K. 2004. Metastability, mechanical strength, and the support of mountain belts. Geology, 32, 625-628. JAMES, D. E. 1971. Plate-tectonic model for the evolution of the central Andes. Geological Society of America Bulletin, 82, 3325-3346. JAMIESON, R. A., BEAUMONT, C., NGUYEN, M. H. & LEE, B. 2002. Interaction of metamorphism, deformation, and exhumation in large convergent orogens. Journal of Metamorphic Geology, 20, 9-24. JAMIESON, R. A., BEAUMONT, C., MEDVEDEV, S. & NGUYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109. DOI: 10.1029/2003 JB002811. JESSUP, M. J., LAW, R. D., SEARLE, M. P. & HUBBARD, M. 2006. Structural evolution and vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/ Nepal: implications for orogen scale flow partitioning. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 379413. JIN, Y., McNuTT, M. K. & Znu, Y. 1994. Evidence from gravity and topography data for folding of Tibet. Nature, 371, 669-674. KENLS, I., URAL J. L., VAN DER ZEE, W., HILGERS, C. & SINTUBIN, M. 2005. Rheology of fine-grained siliciclastic rocks in the middle crust - evidence from structural and numerical analysis. Earth and Planetary Science Letters, 233, 351-360. KIND, R., Ni, J., ZHAO, W. ETAs 1996. Evidence from earthquake data for a partially molten crustal layer in southern Tibet. Science, 274, 1692-1694. KIND, R., YUAN, X., SAUL, J. e r AL. 2002. Seismic images of crust and upper mantle beneath Tibet: Evidence for Eurasian plate subduction. Science, 298, 1219-1221. KING, R. W., SHEN, F., BURCHFIEL, B. C. E T A L . 1997. Geodetic measurement of crustal motion in southwest China. Geology, 25, 179-182. KIRBY, E., WmPPLE, K. X., BURCHFIEL, B. C., WENQING, T., BERGER, G., ZHIMING, S. & ZHILIANG, C. 2000. Neotectonics of the Min Shan, China; implications for mechanisms driving Quaternary deformation along the eastern margin of the Tibetan Plateau. Geological Society of America Bulletin, 112, 375-393. KIRBY, E., REINERS, P. W., KROL, M. A., WHIPPLE, K. X., HODGES, K. V., FARLEY, K. A., TANG, W. Q. & CHEN, Z. L. 2002. Late Cenozoic evolution of the eastern margin of the Tibetan Plateau: Inferences from 40Ar/ 39Ar and (U-Th)/ He thermochronology. Tectonics, 21, 3-22. KLEMPERER, S. L. 2006. Crustal flow in Tibet: geophysical evidence for the physical state of Tibetan lithosphere and inferred patterns of active flow. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and

87

Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 39-70. LAMB, S. & HOKE, L. 1997. Origin of the high plateau in the Central Andes, Bolivia, South America. Tectonics, 16, 623-649. LANGIN, W. R., BROWN, L. D. & SANDVOL, E. 2003. Seismicity of central Tibet from Project INDEPTH III seismic recordings. Bulletin of the Seismological Society of America, 93, 2146-2159. LAVt~, J. & AVOUAC, J. P. 2000. Active folding of fluvial terraces across the Siwalik Hills, Himalayas of central Nepal. Journal of Geophysical Research, 105, 5735-5770. LAW, R. D., SEARLE, M. P. & SIMPSON, R. L. 2004. Strain, deformation temperatures and vorticity of flow at the top of the Greater Himalayan Slab, Everest Massif, Tibet. Journal of the Geological Society, London, 161, 305-320. LE FORT, P. 1975. Himalayas: the collided range. Present knowledge of the continental arc. American Journal of Science, 275-A, 1-44. LEE, J., HACKER, B. R., DINKLAGE, W. S. E T A L . 2000. Evolution of the Kangmar Dome, southern Tibet; structural, petrologic, and thermochronologic constraints. Tectonics, 19, 872-895. LEE, J., HACKER, B. & WANG, Y. 2004. Evolution of North Himalayan gneiss domes: structural and metamorphic studies in the Mabja Dome, southern Tibet. Journal of Structural Geology, 26, 22972316. LEECH, M. L., SINGH, S., JAIN, A. K., KLEMPERER, S. L. & MANICKAVASAGAM, R. M. 2005. The onset of India-Asia continental collision: Early, steep subduction required by the timing of UHP metamorphism in the western Himalaya. Earth and Planetary Science Letters, 234, 83-97. LI, S., UNSWORTH, M. J., BOOKER, J. R., WEI, W., TAN, H. & JONES, A. G. 2003. Partial melt or aqueous fluid in the mid-crust of southern Tibet? Constraints from INDEPTH magnetotelluric data. Geophysical Journal International, 153, 289-304. MCCLELLAND, W. C. & GILOTTI, J. A. 2003. Latestage extensional exhumation of high-pressure granulites in the Greenland Caledonides. Geology, 31, 259-262. MCQUARRIE, N. & CHASE, C. G. 2000. Raising the Colorado Plateau. Geology, 28, 91-94. MAKOVSKY, Y. & KLEMPERER,S. L. 1999. Measuring the seismic properties of Tibetan bright spots: Evidence for free aqueous fluids in the Tibetan middle crust. Journal of Geophysical Research, 104, 10,795-10,825. MAKOVSKY, Y., KLEMPERER, S. L., LIYAN, H. & DEYUAN, L. 1996. Structural elements of the southern Tethyan Himalaya crust from wideangle seismic data. Tectonics, 15, 997-1005. MASEK, J. G., ISACKS, B. L. & FIELDING, E. J. 1994. Rift flank uplift in Tibet: Evidence for a viscous lower crust. Tectonics, 13, 659-667. MEANS, W. D. 1989. Stretching faults. Geology, 17, 893 -896. MECHIE, J., SOBOLEV, S. V., RATSCHBACHER, L. ET AL. 2004. Precise temperature estimation in the

88

K.V. HODGES

Tibetan crust from seismic detection of the ~x-/3 quartz transition. Geology, 32, 601-604. MEISSNER, R. & MOONEY, W. 1998. Weakness of the lower continental crust: a condition for delamination, uplift, and escape. Tectonophysics, 296, 47-60. MOLNAR, P. 2004. Late Cenozoic increase in accumulation rates of terrestrial sediment: How might climate change have affected erosion rates?

Annual Review of Earth and Planetary Sciences, 32, 67-89. MOLNAR, P. 2005. Mio-Pliocene growth of the Tibetan Plateau and evolution of East Asian climate. Palaeontologia Electronica, 8, 1-23. MOLNAR, P., ENGLAND, P. & MARTINOD, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the Indian monsoon. Reviews of Geophysics, 31, 357-396. NAKATA, T. 1989. Active faults of the Himalaya of India and Nepal. In: MALINCONICO, L. L. & LmLIE, R. J. (eds) Tectonics of the Western Himalayas. Geological Society of America, Special Papers 232, 243-264. NELSON, K. D., ZrtAO, W., BROWN, L. D. eTAL. 1996. Partially molten middle crust beneath southern Tibet: Synthesis of Project INDEPTH Results. Science, 274, 1684-1688. NI, J. & BARAZANGI, M. 1984. Seismotectonics of the Himalayan collision zone; Geometry of the underthrusting Indian Plate beneath the Himalaya. Journal of Geophysical Research, 89, 1147-1163. OUIMET, W., WHIPPLE, K., ROYDEN, L. & GRANGER, D. 2005. Long transient response times of rivers in eastern Tibet to regional plateau uplift: The affect of mega-landslides. Geophysical Research Abstracts, 7, EGU05-A-05743. OWENS, T. J. & ZANDT, G. 1997. Implications of crustal property variations for models of Tibetan Plateau evolution. Nature, 387, 37-43. PEARSON, O. N. & DECELLES, P. G. 2005. Structural geology and regional tectonic significance of the Ramgarh thrust, Himalayan fold-thrust belt of Nepal. Tectonics, 24. DOI: 10.1029/2003TC001617. P~CHER, A. 1991. The contact between the higher Himalyan crystallines and the Tibetan sedimentary series: Miocene large-scale dextral shearing. Tectonics, 10, 587-599. POWELL, C. M. & CONAGHAN, P. J. 1973. Plate tectonics and the Himalayas. Earth and Planetary Science Letters, 20, 1 - 12. RANALLI, G. 1997. Rheology of the lithosphere in space and time. In: BURG, J.-P. & FORD, M. (eds) Orogeny Through Time. Geological Society, London, Special Publications, 121, 19-37. ROWLEY, D. B., PIERREHUMBERT, R. T. & CURRIE, B. S. 2001. A new approach to stable isotopebased paleoaltimetry: Implications for paleoaltimetry and paleohypsometry of the High Himalaya since the late Miocene. Earth and Planetary Science Letters, 188, 253-268. ROYDEN, L. 1996. Coupling and decoupling of crust and mantle in convergent orogens: Implications for strain partitioning in the crust. Journal of Geophysical Research, 101(B8), 17,679-17,705.

ROYDEN, L. H., BURCHFIEL, B. C., KING, R. W., WANG, E., CHEN, Z., SHEN, F. & LIu, Y. 1997. Surface deformation and lower crustal flow in eastern Tibet. Science, 276, 788-790. RUHL, K. W. & HODGES, K. V. 2005. The use of detrital mineral cooling ages to evaluate steady-state assumptions in active orogens: an example from the central Nepalese Himalaya. Tectonics, 24. DOI: 10.1029/2004TC001712. RUHL, K. W., BLYTHE, A. E. & HODGES, K. V. 2005. Accelerated Late-Pliocene Himalayan erosion from fission-track and 4~ thermochronology and the possible role of climate change. Geo-

logical Society of America Abstracts Programs, 37, 346.

with

SACHAN, H. K., MUKHERJEE,B. K., OGASAWARA,Y., MARUYAMA, S., ISHIDA, H., MUKO, A. & YOSmOKA, N. 2004. Discovery of coesite from Indus Suture Zone (ISZ), Ladakh, India: Evidence for deep subduction. European Journal of Mineralogy, 16, 235-240. SCAILLET, B., PICHAVANT, M. & ROUX, J. 1995. Experimental crystallization of leucogranite magmas. Journal of Petrology, 36, 663 -705. SCHELLING, D. 1992. The tectonostratigraphy and structure of the eastern Nepal Himalaya. Tectonics, 11, 925-943. SCHELLING, D. & ARITA, K. 1991. Thrust tectonics, crustal shortening, and the structure of the fareastern Nepal Himalaya. Tectonics, 10, 851-862. SCHLUP, M., CARTER, A., COSCA, M. & STECK, A. 2003. Exhumation history of eastern Ladakh revealed by 4~ and fission-track ages: the Indus River - Tso Morari transect, NW Himalaya. Journal of the Geological Society, London, 160, 385 -399. SCHOENBOHM, L. M., WHIPPLE, K. X. & BURCHFIEL, B. C. 2004. Geomorphic constraints on surface uplift, exhumation, and plateau growth in the Red River region, Yunnan Province, China. Geological Society of America Bulletin, 116, 895-909. SEARLE, M. P. 1986. Structural evolution and sequence of thrusting in the High Himalayan, Tibetan-Tethys and Indus suture zones of Zanskar and Ladakh, Western Himalaya. Journal of Structural Geology, 8, 923-936. SEARLE, M. P. 1996. Cooling history, erosion, exhumation, and kinematics of the HimalayaKarakoram-Tibet orogenic belt. In: YIN, A. & HARRISON, T. M. (eds) Tectonics of Asia. Cambridge University Press, Cambridge, 110-137. SEARLE, M. P. 1999. Extensional and compressional faults in the Everest-Lhotse massif, Khumbu Himalaya, Nepal. Journal of the Geological Society, London, 156, 227-240. SEARLE, M. P. & GODIN, L. 2003. The South Tibetan detachment and the Manaslu leucogranite: A structural reinterpretation and restoration of the Annapurna-Manaslu Himalaya, Nepal. Journal of Geology, 111, 505-523. SEARLE, M. P. & REX, A. J. 1989. Thermal model for the Zanskar Himalaya. Journal of Metamorphic Geology, 7, 127-134.

CHANNEL FLOW-EXTRUSION HYPOTHESIS SEARLE, M. P. & SZULC, A. G. 2005. Channel flow and ductile extrusion of the high Himalayan slabthe Kangchenjunga-Darjeeling profile, Sikkim Himalaya. Journal of Asian Earth Sciences, 25, 173-185. SEARLE, M. P., WATERS, D. J., REX, D. C. & WILSON, R. N. 1992. Pressure, temperature, and time constraints on Himalayan metamorphism from eastern Kashmir and western Zanskar. Journal of the Geological Society, London, 149, 753-773. SEARLE, M. P., PARRISH, R. R., HODGES, K. V., HURFORD, A., AYRES, M. W. & WHITEHOUSE, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya; field relations, geochemistry, age, origin, and emplacement. Journal of Geology, 105, 295 -317. SEARLE, i . , SIMPSON, R. L., LAW, R. D., PARRISH, R. R. 8r WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal-South Tibet. Journal of the Geological Society, London, 160, 345-366. SHAPIRO, N. M., RITZWOLLER, M. H., MOLNAR, P. & LEVIN, V. 2004. Thinning and flow of Tibetan crust constrained by seismic anisotropy. Science, 305, 233 -236. SHEN, F., ROYDEN, L. H. & BURCHFIEL, B. C. 2001. Large-scale crustal deformation of the Tibetan Plateau. Journal of Geophysical Research, B, Solid Earth and Planets, 106, 6793 -6816. SORKHABI, R. B., STUMP, E., FOLAND, K. A. & JAIN, A. K. 1996. Fission-track and 4~ evidence for episodic denudation of the Gangotri granites in the Garhwal Higher Himalaya, India. Tectonophysics, 260, 187-199. SPICER, R. A., HARRIS, N. B. W., WIDDOWSON, M., HERMAN, A. B., Guo, S., VALDES, P. J., WOLFE, J. A. & KELLEY, S. P. 2003. Constant elevation of southern Tibet over the past 15 million years. Nature, 421, 622-624. STRACHAN, R. A. & MARTIN, M. W. 2001. Evidence for contemporaneous yet contrasting styles of granite magmatism during extensional collapse of the northeast Greenland Caledonides. Tectonics, 20, 458-473. THEIDE, R. C., BOOKHAGEN, B., ARROWSMITH, J. R., SOBEL, E. R. & STRECKER, M. R. 2004. Climatic control on rapid exhumation along the southern Himalayan front. Earth and Planetary Science Letters, 222, 791-806. TURNER, S., HAWKESWORTH,C., LIU, J., ROGERS, N., KELLEY, S. & VAN CALSTEREN,P. 1993. Timing of Tibetan uplift constrained by analysis of volcanic rocks. Nature, 364, 50-54. VAN NGOC, P., BOYER, D., THERME, P., YUANG,X. C., LI, L. & J1N, G. Y. 1986. Partial melting zones in the crust in southern Tibet from magnetotelluric results. Nature, 311, 310-314. VANNAY, J.-C. & GRASEMANN, B. 2001. Himalayan inverted metamorphism and syn-convergence extension as a consequence of a general shear extrusion. Geological Magazine, 138, 253-276. VANNAY, J.-C. & HODGES, K. V. 1996. Tectonometamorphic evolution of the Himalayan

89

metamorphic core between Annapurna and Dhaulagiri, central Nepal. Journal of Metamorphic Geology, 14, 635-656. VANNAY, J.-C., GRASEMANN, B., RAHN, M., FRANK, W., CARTER, A., BAUDRAZ, V. & COSCA, M. 2004. Miocene to Holocene exhumation of metamorphic crustal wedges in the NW Himalaya: Evidence for tectonic extrusion coupled to fluvial erosion. Tectonics, 23. DOI: 10.1029/ 2002TC001429. WANG, E. & BURCHFIEL,B. C. 2000. Late Cenozoic to Holocene deformation in southwestern Sichuan and adjacent Yunnan, China, and its role in formation of the southeastern part of the Tibetan Plateau. Geological Society of America Bulletin, 112, 413-423. WATTS, D. R., HARRIS, N. B. W. & GROUP, T. N. G. S. W. 2005. Mapping granite and gneiss in domes along the North Himalayan antiform with ASTER SWIR band ratios. Geological Society of America Bulletin, 117, 879-886. WDOWINSKI, S. ~ AXEN, G. J. 1992. Isostatic rebound due to tectonic denudation: A viscous flow model of a layered lithosphere. Tectonics, 11, 303-315. WEI, W., UNSWORTH, M., JONES, A. G. ET AL. 2001. Detection of widespread fluids in the Tibetan crust by magnetotelluric studies. Science, 292, 716-718. WERNICKE, B. P. 1985. Uniform-sense normal simple shear of the continental lithosphere. Canadian Journal of Earth Sciences, 22, 108-125. WERNICKE, B. & AXEN, G. J. 1988. On the role of isostacy in the evolution of normal fault systems. Geology, 16, 848- 851. WESTAWAY, R. 1995. Crustal volume balance during the India-Eurasia collision and altitude of the Tibetan plateau: A working hypothesis. Journal of Geophysical Research, 100, 15,173-15,192. WHITE, R. W., POWELL, R. & HOLLAND,T. J. B. 2001. Calculation of partial melting equilibria in the system Na20-CaO-KzO-FeO-MgO-AlzO3SiO2-H20 (NCKFMASH). Journal of Metamorphic Geology, 19, 139-153. WHITE, A. P., HODGES, K. V., MARTIN, M. W. & ANDRESEN, A. 2002a. Geologic constraints on middle-crustal behavior during broadly synorogenic extension in the central East Greenland Caledonides. International Journal of Earth Sciences, 91, 187-208. WHITE, N. M., PRINGLE, M., GARZANTI, E., BICKLE, M., NAJMAN, Y., CHAPMAN, H. & FRIEND, P. 2002b. Constraints on the exhumation and erosion of the High Himalayan Slab, NW India, from foreland basin deposits. Earth and Planetary Science Letters, 195, 29-44. WILLIAMS, H., TURNER, S., KELLEY, S. & HARRIS, N. 2001. Age and composition of dikes in southern Tibet: New constraints on the timing of east-west extension and its relationship to postcollisional volcanism. Geology, 29, 339-342. WOBUS, C. W., HODGES, K. V. & WHIPPLE, K. X. 2003. Has focused denudation sustained active thrusting at the Himalayan topographic front? Geology, 31, 861-864.

90

K.V. HODGES

WOBUS, C., HEIMSATH,A., WHIPPLE,K. & HODGES, K. 2005. Active out-of-sequence thrust faulting in the central Nepalese Himalaya. Nature, 434, 1008-1010. Wu, C., NELSON, K. D., WORTMAN, G., SAMSON, S. D., YUE, Y., LI, J., KIDD, W. S. F. & EDWARDS, M. A. 1998. Yadong cross structure and South Tibetan detachment in the east central Himalaya (89~176 Tectonics, 17, 28-45. YUAN, X., NI, J., KIND, R., SANDVOL, E. 8z MECHIE, J. 1997. Lithospheric and upper mantle structure of southern Tibet from a seismological passive source experiment. Journal of Geophysical Research, 102, 27,491-27,500.

ZHANG, P., SHEN, Z., WANG, M. ETAL. 2004. Continuous deformation of the Tibetan Plateau from global positioning system data. Geology, 32, 809-812. ZHAO, W., NELSON, K. D. & Project INDEPTH Team. 1993. Deep seismic reflection evidence for continental underthrusting beneath southern Tibet. Nature, 366, 557-559. ZHAO, W. L. & MORGAN, W. J. 1985. Uplift of the Tibetan Plateau. Tectonics, 4, 359-369. ZHAO, W. L. & MORGAN, W. J. P. 1987. Injection of Indian crust into Tibetan lower crust: a twodimensional finite element model study. Tectonics, 6, 489-504.

Crustal flow modes in large hot orogens C. B E A U M O N T 1, M. H. N G U Y E N 1'2, R. A. J A M I E S O N 2 & S. E L L I S 3

l Oceanography Department, Dalhousie University, Halifax, Nova Scotia, Canada B 3 H 4J1 (e-mail: chris, beaumont @dal.ca) 2Department o f Earth Sciences, Dalhousie University, Halifax, Nova Scotia, Canada B 3 H 3J5 3Institute f o r Geological and Nuclear Sciences, L o w e r Hutt, N e w Zealand Abstract: Crustal-scale channel flow numerical models support recent interpretations of

Himalayan-Tibetan tectonics proposing that gravitationally driven channel flows of low-viscosity, melt-weakened, middle crust can explain both outward growth of the Tibetan Plateau and ductile extrusion of the Greater Himalayan Sequence. We broaden the numerical model investigation to explore three flow modes: homogeneous channel flow (involving laterally homogeneous crust); heterogeneous channel flow (involving laterally heterogeneous lower crust that is expelled and incorporated into the mid-crustal channel flow); and the hot fold nappes style of flow (in which mid-/lower crust is forcibly expelled outward over a lower crustal indentor to create fold nappes that are inserted into the mid-crust). The three flow modes are members of a continuum in which the homogeneous mode is driven by gravitational forces but requires very weak channel material. The hot fold nappe mode is driven tectonically by, for example, collision with a strong crustal indentor and can occur in crust that is subcritical for homogeneousflows. The heterogeneous mode combines tectonic and gravitationally driven flows. Preliminary results also demonstrate the existence and behaviour of mid-crustal channels during advancing and retreating dynamical mantle lithosphere subduction. An orogen temperature-magnitude (T-M) diagram is proposed and the positions of orogens in T-M space that may exhibit the flow modes are described, together with the characteristic positions of a range of other orogen types.

During the last decade we have developed and used a range of finite-element numerical models to gain insight into collisional orogenesis. These types of models include two-dimensional doubly (bi-), vergent (Willett et al. 1993; Beaumont et al. 1994; Beaumont & Quinlan 1994), three-dimensional doubly (bi-)vergent (Braun & Beaumont 1995), vise (Ellis et al. 1998), accretionary wedge (Beaumont et al. 1999), Pyrenean (Beaumont et al. 2000) and Alpine styles (Beaumont et al. 1996a; Ellis et al. 1999; Pfiffner et al. 2000). Both mechanical and thermo-mechanically coupled (Jamieson et al. 1998) techniques have been applied to small cold and large hot orogens (Jamieson et al. 2002). The applications to small orogens include the Pyrenees, Alps, Southern Alps of New Zealand (Beaumont et al. 1992, 1996b; Waschbusch et al. 1998), and to examples studied by the Canadian Lithoprobe programme (Ellis & Beaumont 1999). Applications to large hot orogens include the Himalayas and Tibet (Willett & Beaumont 1994; Beaumont et al. 2004; Jamieson et al. 2004b). Each type of orogen occupies a characteristic position in an orogenic temperature-magnitude (T-M) diagram (Fig. 1). This concept is inspired by the astrophysical Hertzsprung-Russell (H-R) star diagram, in which luminosity (or absolute

visual magnitude) is plotted against spectral type (or surface temperature) for star populations (Fig. la, top right) (Hertzsprung 1905; Russell 1914). The H-R diagram concisely describes stellar conditions and provides insight into the range of stellar evolution. The T-M diagram (Fig. 1) is intended to express relationships between the magnitude of the orogen, measured by excess crustal or lithospheric thickness of the orogen relative to that of undeformed standard continental lithosphere, and the excess heat content or temperature of the orogen relative to the same undeformed lithosphere with standard heat production. The T-M diagram provides a first-order classification of orogen types (e.g. Dwarfs, Giants, Fig. la) and offers insight into the underlying tectonic processes. In addition, the evolution of orogens can be represented by evolutionary paths in the T-M diagram. For example, an accretionary wedge may evolve into a cordilleran orogen and thence to a large continent-continent collisional orogen. An 'orogenic main sequence' (MS, Fig. l) extends from bottom left toward the top right. Orogens on the main sequence have excess conductive steady-state temperatures complementing their excess crustal thicknesses. The main sequence is non-linear for two reasons. Curvature of the lower

Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 91-145. 0305-8719/06l$15.00 9 The Geological Society of London 2006.

From: LAW,R. D., SEARLE,M. P. & GODIN,L. (eds)

92

C. BEAUMONT ET AL.

Fig. 1. Orogen temperature-magnitude (T-M) diagram. (a) Classification of orogen types and comparison of T-M diagram with the HertzsprungRussell (H-R) diagram for stars. (b) Suggested classification of particular orogens. (e) Classification of types of mechanical and thermal-mechanical models (see text) that have been used to model different orogen types according to position in T-M space (colour figures and animations are available at http:// ge~176176176176 extras.html).

part occurs because conductive steady-state temperatures contain a quadratic or higher order term owing to radioactive heat production. Steady-state average temperatures in thickened crust/lithosphere increase disproportionately faster than the corresponding thickening of the orogen and, therefore, the MS plots below the M = ceT line, where c~ is a measure of the ratio between magnitude and temperature in the standard state. The convex-up curvature for large orogens expresses the trend to a gravitational limit on the maximum thickness of hot, weak crust and lithosphere. Here the T-M diagram is presented in a one-dimensional form, but even this conceptual view can take account of some three-dimensional aspects. For example, orogens that grow during orthogonal collision accumulate mass more rapidly than equivalent orogens where motion is primarily transcurrent, and the former will evolve more rapidly in T-M space. Transcurrent orogens may never evolve out of the small cold part of T-M space. The primary division in a T-M diagram is between small cold and large hot orogens (Fig. lb). Small cold orogens such as accretionary wedges, Southern Alps of New Zealand, Pyrenees, and Alps plot in the lower left part of diagram (Fig. lb); they lack the levels of crustal thickening and associated thermal relaxation necessary to achieve high temperatures. This may be because they are young (Proto-Main Sequence, Fig. la), strongly denuded (Denudation Dwarfs, Fig. la), dominantly transcurrent, or have low levels of radioactive heating (Accretionary Coldies, Fig. la). It follows that their minimum crustal viscosities are too high for large-scale fluid-like flows in the middle and/or lower crust. At the other end of the orogenic main sequence, the large hot orogens (Giants and Super Giants, Fig. 1a) are both massive and hot, leading to weak viscous regions in the crust that may contain in situ partial melts and may undergo gravitationally driven channel flows (Bird 1991; Westaway 1995; Royden 1996; Royden et al. 1997; Beaumont et al. 2001; Shen et al. 2001). Such flows of 'meltweakened' crust can explain both eastward growth of the Tibetan Plateau, as the channel tunnels outward (Clark & Royden 2000, and references therein), and ductile extrusion of the Greater Himalayan Sequence (Grujic et al. 1996, 2002; Beaumont et al. 2001, 2004; Jamieson et al. 2004b). We regard gravitationally driven channel flow as an end-member requiring a combination of sufficiently low viscosities, thick channels and large differences in mean elevation between the orogen and its foreland to allow the available differential pressures to drive efficient flow (Bird 1991; Clark & Royden 2000). If gravitationally driven crustal flow is an end-member that exists only in

CRUSTAL FLOW MODES Giant and Super Giant orogens like Tibet (Fig. lb & a), do other flow modes occur when conditions are subcritical for gravitational forcing? If so, what drives these flows? How do these more general flow regimes relate to the evolutionary paths outlined in Figure 1? We address these questions in this paper. The mechanics of and types of models used to investigate small cold orogens (Fig. lc) have been described in the earlier papers cited above. Here we expand on the types of flow that can occur within large hot orogens. We use numerical models to investigate three modes of crustal flow. Mode 1 is homogeneous channel flow; Mode 2 is heterogeneous channel flow, which incorporates lower crustal blocks within the channel; Mode 3, hot fold nappes, measures the response of the model orogen to the insertion of progressively stronger blocks of lower crust. We interpret these as members of a continuum of gravitationally and tectonically driven flow modes, and relate the results to the corresponding deformation predicted for orogenic crust in the large hot orogen region of the T-M diagram. We also provide preliminary results from upper-mantle-scale models that address the fate of lithospheric mantle during continent-continent collision, and show that channel flows also exist within this model style.

Numerical calculation of crustal- and upper-mantle-scale flows The numerical modelling methodology is outlined in this section. An explanation for the choice of the model parameter values and the sensitivity of the results to this choice is included in the Appendix. There are two types of models. Crustal-scale (CS) models were described in Beaumont et al. (2004) but an explanation is included here for completeness. Upper-mantle-scale (UMS) models are discussed later but their primary properties are described here. In both CS and UMS we model the development of large hot orogens using a twodimensional (2D) finite element code that assumes plane-strain conditions in a vertical cross-section through the orogen. The codes compute thermal and mechanical evolution subject to velocity boundary conditions applied at the sides and base of the CS model region, and applied at the sides of the UMS model region. Thermal-mechanical coupling occurs through the thermal activation of viscous power-law creep in the model materials and through the redistribution of radioactive crust by material flow. The CS model properties are similar to those described by Beaumont et al. (2004) and Jamieson et al. (2004b). This model, which is 2000 km wide, has two regions: the crust (Fig. 2a & b), in which the

93

velocity and deformation are calculated dynamically; and the mantle, where the velocity is prescribed kinematically (Fig. 2b). In the 2000 • 600 km UMS model, the velocity and deformation for the whole model domain are calculated dynamically subject to far-field lateral velocity boundary conditions. The associated temperature field is calculated for the whole model domain. Model parameters and values for CS and USM models are given in Tables 1 and 2 respectively.

CS model velocity boundary conditions and reference f r a m e s In the CS models described in this paper, both the pro- and retro-mantle lithospheres (Fig. 2) converge at a uniform velocity, Vp = -VR, and detach and subduct beneath the stationary S-point (Willett et al. 1993). This frame of reference and associated symmetric convergence were chosen to give results that are least dependent on the motion of the lithospheric plates with respect to the sublithospheric mantle. The subducted mantle lithosphere descends into the mantle at constant dip with constant kinematically specified velocity (Fig. 2). The CS models can be interpreted in other reference frames (Beaumont et al. 2004; fig. 9) by adding or subtracting a fixed velocity to all of the boundary velocities and the velocity of the S-point. For example, Jamieson et al. (2006) investigate the case where the promantle lithosphere converges at 2Ve, the S-point advances at Vs = Ve, and the retro-mantle lithosphere is stationary, Vn = 0. This reference frame is considered most appropriate for HimalayanTibetan models (Beaumont et al. 2001, 2004; Jamieson et al. 2004b). The change in reference frame does not change the model results, only the way in which they are viewed.

U M S model velocity boundary conditions and reference f r a m e s The UMS model (Fig. 16) is designed to correspond approximately to the collision of India with Asia. The boundary condition has pro-lithosphere, equivalent to India, converging from the left boundary at a uniform velocity of Ve = 5 cm a-1 against a stationary retro-lithosphere, VR ----0 cm a-1, at the fight boundary corresponding to Asia. In contrast to the CS models, Vs is not specified but is determined by the dynamical evolution of the model. In the sublithospheric mantle region, the sides and base have free slip boundary conditions. A small uniform symmetric outward leakage flux of material is specified through the side boundaries to balance the flux of pro-lithosphere into the

94

C. BEAUMONT ETAL.

Fig. 2. Initial crustal model conditions. Only the central part of the 2000 km long model is shown. (a) Passive Lagrangian marker grid and mechanical layers; '0' model surface suture position above subduction point, S. (b) Initial thermal structure, radioactive layers, A1 and A2 and conductive steady-state isotherms, and general velocity vectors, showing convergence with Vp = 1.5, VR = - 1.5 and Vs = 0 cm a ~and implied double subduction of the mantle lithospheres beneath S. (c) Relationship between initial mechanical and thermal layers and summary of parameters (see also Table 1); effect of reduction in viscosity for quartz-rich upper and mid-crust from flow law value at 700~ to 1019 Pa s at 750~ (melt-weakening); effective viscosity used in model is shown by solid line.

CRUSTAL FLOW MODES

95

Table 1. Parameters used in models (see also Fig. 2) Parameter

(a) Mechanical parameters Pcrust Pmantle

D

0 ~b~#( 0 - 1 0 kin)

Meaning Crustal density Mantle density Flexural rigidity (isostasy model) Crustal thickness Lower crustal thickness Subduction dip angle Effective internal angle of friction

~beff ( 1 0 - 3 5 kin)

C P

J~ "0effv= B ' 0 2 ) (1-')/2n exp[Q/nRTx]

Cohesion Solid pressure Second invariant of the deviatoric stress tensor General equation for effective viscosity

.!

12 R

Tx B*, n, Q as below WQ ( 0 - 1 0 km)

WQ z 5 1 0 - 2 5 km or

Value(s) 2700 kg m -3 3300 kg m -3 1022 Nm 35 km see below 20 ~ 5~ 15 ~ 10 MPa Pa Pa 2

Second invariant of strain rate tensor Gas constant Absolute temperature

s-2 8.314 J m o l - ~ K -~ K

Wet Black Hills quartzite flow law (after Gleason & Tullis 1995)

n = 4.0 B* = 2.92 • 106 Pa S1/4 Q = 223 kJ mol-1 B* = B*(WQ) • 5 (etc.)

Modified wet Black Hills quartzite flow law

1 0 - 2 0 km (see below)

DMD

Dry Maryland diabase flow law (after Mackwell et al. 1998)

n = 4.7 B* = 1.91 x 105 Pa S1/4"7 Q = 485 kJ mol-1

DMD/f (see below)

Scaled dry Maryland diabase flow law Linear reduction in effective viscosity over T range 7 0 0 - 7 5 0 ~ for W Q only Length of Eulerian model domain

B* = B*(DMD)/f

'melt-weakening'

'0700 = flow law value "0750 = 1019 Pa s 2000 km

(b) Crustal-scale models basal velocity boundary conditions

v~ vR Vs

(c) Thermal parameters Ce K K

qm

qs A 1 ( 0 - 20 km) A2 ( 2 0 - 3 5 km)

Pro-side (convergence) velocity Retro-side velocity S-point velocity (subduction advance) Heat capacity Thermal conductivity Thermal diffusivity (K = K/pCp, where pCp = 2 • 106) Surface temperature Temperature at lithosphere/ asthenosphere boundary Basal mantle heat flux Initial surface heat flux Upper-crustal heat production Lower-crustal heat production

(d) Crustal-scale models surface denudation slope • f(t) • g(x) Denudation rate slope

f(t) g(x) f(t) g(x)

1.5 c m a - l - 1.5 c m a - 1

0 c m a -1 750 m 2 K-~s 2 2.00 W m - l K -1 1.0 • 10 -6 m 2 s -1 0~ 1350~ 20 m W m -2 71.25 m W m -2 2.0 txW m -3 0.75 txW m -3 ma

-1

Local surface slope measured from finite element mesh Time function (specifies how denudation rate varies with time when g(x) and slope = 1) Spatial function (specifies how denudation rate varies with position x)

g(x) -- 0 -- arid g(x) = 1 -- wet

0.0399 m a-1 1.0 1.0 --+ 0.0 (varies linearly) 0.0

t > 0 (i.e. constant) 0 < x < 500 km 500 < x < 550 km x >_ 550 km

constant

(Continued)

C. BEAUMONT ET AL.

96 Table 1. Continued Parameter

Meaning

Value(s)

(e) Specific model parameters - crustal-scale models LHO-1 Lower crust (25-35 km)

B*(DMD/5) 15 ~

LHO-2 Lower crust (25-35 km) Altemating 250 km long blocks

B*(DMD) B*(DMD/IO)

LHO-3 Lower crust (25-35 km) 250 km long blocks arranged symmetrically with respect to S. Blocks have properties B*(DMD), B*(DMD/4), B*(DMD/8), B*(DMD/12), B*(DMD/16), B*(DMD/20). Order is from external to internal part of model.

model; there is no material flux through the base of the model.

CS and UMS mechanical models The mechanical models used to calculate the CS and UMS velocity fields and deformation (Fullsack 1995) use an Arbitrary Lagrangian Eulerian (ALE) methodology in which flows with free upper surfaces and large deformation are calculated on a Eulerian finite-element grid that stretches in the vertical direction to conform to the material domain. A Lagrangian grid, which is advected with the model velocity field, is used to update the mechanical and thermal material property distributions on the Eulerian grid as their position changes. Flow is driven by the basal and lateral velocity boundary conditions described above. The finite-element model uses a viscous-plastic rheology. The plastic (frictional or brittle) deformation is modelled with a pressure-dependent Drucker-Prager yield criterion. Yielding occurs when (j;)l/2 = p sin 4,eft + C cos beg

(1)

where J~ is the second invariant of the deviatoric stress, P is the dynamical pressure (mean stress), C is the cohesion, and the internal angle of friction, 4,e#-, is defined to include the effects of pore fluid pressures through the relation P sin 4,e# = (P - PU) sin 4'

(2)

For dry frictional sliding conditions (approximating Byeflee's law), 4, = 30 ~ when the pore fluid pressure, Pf = 0. For hydrostatic fluid pressures and typical crustal densities 4,e~-is approximately 15 ~ and for overpressured pore fluid conditions we use 4,eft = 5 ~ (see Appendix). The incompressible plastic flow becomes equivalent to a viscous material (Fullsack 1995; Willett

1999) such that rleeff=(J;)'/2/2('Is '/2, where 02) 1/2 is the second invariant of the deviatoric strain rate. Setting the viscosity to r/~ in regions that are on frictional-plastic yield satisfies the yield condition and allows the velocity field to be determined from the finite-element solution for viscous creeping flows. The overall non-linear solution is determined iteratively using ~ = r/e~ for regions of plastic flow, and r / = r/verr, as defined below, for regions of viscous flow. The flow is viscous when the flow stress is less than the plastic yield stress for the local ambient conditions. Under these circumstances the powerlaw creep effective viscosity is v = B.(~,2)~,-.,~/2. exp[Q/nRTg] ~efr

(3)

The values of B*, n, and Q (Table 1) are based on laboratory experiments with A values converted to B* assuming cylindrical creep tests. The theology of the upper and middle crust is based on the 'wet Black Hills quartzite' (WQ) flow law (Gleason & Tullis 1995). In the CS model experiments (Table 1) we use the flow law with B* = B*(WQ) in the uppermost crust (initially 0 - 1 0 k m ) . In the mid-crust (initially 10-25 km) B* is scaled by a factor of 5 (B* = B*(WQ x 5)), as explained in the Appendix. In the UMS models the upper and middle crust have B * = B*(WQ • 5) (Table 2). The rheology of the lower crust (initially 2 5 - 3 5 km) is based on the 'dry Maryland diabase' (DMD) flow law (Mackwell et al. 1998) (Table 1), which is also scaled to achieve a range of effective lower-crustal strengths (see Appendix). The reference CS rheological structure in model LHO-I represents a laterally uniform three-layer crust. The upper layer (B*(WQ)) has weak frictional-plastic properties, 4,eft = 5 ~ The middle layer ( B * ( W Q x 5)) has standard hydrostatic

CRUSTAL FLOW MODES

97

Table 2. Parameters used in upper-mantle-scale (UMS) models (where different from those of the

crustal scale (CS) models) (see also Fig. 16 LHO-LS1 and LHO-LS2) Parameter

Meaning

Value(s)

(a) Geometry

(b) Mechanical parameters Puc Pk, pe Pml

Pure

~bee(O- 28 km)

~beff(28-34 km)

~be#(34-600 km)

Cuc WQ x 5 ( 0 - 2 8 km) DMD/IO ( 2 8 - 3 4 km) WO x 10 ( 3 4 - 1 0 0 km) WO (100 - 600 km) WO

Model domain Eulerian mesh

2000 x 600 km 101 x 201

Upper crust density at 194~ * Lower crust density at 457~ * Nominal lower crust density when transformed to eclogite facies Mantle lithosphere density at 937~ * in model LHO-LS 1 in model LHO-LS2 Uniform sublithospheric upper mantle density Effective internal angle of friction (strain softens linearly over range 0.5 --~ 1.5 of second invariant of strain) Effective internal angle of friction (strain softens linearly over range 0.5 ~ 1.5 of second invariant of strain) Effective internal angle of friction (strain softens linearly over range 0.5 --~ 1.5 of second invariant of strain) Cohesion Viscous flow laws (see Table 1)

2800 kg m -3 2950 kg m - 3 3100 kg m -3

Scaled olivine flow law Olivine flow law Wet Aheim dunite (olivine) flow law (after Chopra & Paterson 1984)

Minimum effective viscosity in sublithospheric mantle

3300 kg m - 3 3310 kg m -3 3260 kg m - 3 15 ~ --+ 2 ~

15~

~

15 ~ --+ 2 ~

10 MPa

B* = B* = B* = B* =

B*(WQ x 5) B* (DMD/10) B*(WO x 10) B*(WO)

n ----4.48 B* = 7.75 x 104 Pa s 1/44a Q = 498 kJ mol-1 V*=0 1019 Pa s

(c) Upper-mantle-scale models velocity boundary conditions Velocity boundary conditions Vp ( 0 - 1 0 0 km) VR ( 0 - 1 0 0 kin)

5 c m a -1 Ocm a - 1 Small flux through side boundaries (see text) Other boundaries, free slip; upper surface, free surface

(d) Thermal properties K K

Kum

Ts T~ T~ qm

Melt weakening (see Table 1 and Fig. 2) Thermal conductivity Thermal diffusivity Thermal diffusivity of sublithospheric upper mantle (adiabatic temperature gradient) Surface temperature Temperature at lithosphere/asthenosphere boundary Initial temperature at model base Basal mantle heat flux

2.00 W m - l K - 1 8.0 x 10 - 7 m 2s -1 3.2 x 10 -5 m 2 S - 1

0~ 1350~ 1493~ 20 m W m - 2

(Continued)

C. BEAUMONT ETAL.

98 Table 2. Continued Parameter qs av A1 (0-20 km) A2 (20-34 km)

Meaning Initial surface heat flux Volume coefficient of thermal expansion Upper crustal heat production Lower crustal heat production

Value(s) 70.5 mW m 2 3 • 10-5/C -1 2.0 IxW m -3 0.75 IxW m -3

*Initial averagetemperature of this model layer.

frictional-plastic properties, 4)eft= 15 ~. This is underlain by lower crust with ~beg----15 ~ and B * = B*(DMD/5). This layering is designed to approximate the continental margin crust commonly involved in collisional orogenesis, with a refractory, intermediate granulite lower crust overlain by middle crust comprising fertile quartzo-feldspathic low-grade metasedimentary and granitic rocks and an upper crust dominated by quartz-rich sedimentary rocks with high pore fluid pressures. In UMS models, the crustal rheology is similar to the CS models (Tables 1 and 2) except that there is no separate weak upper-crustal layer. The initial thicknesses of the layers are also slightly different because the Eulerian finite-element resolution is lower, comprising 17 as opposed to 40 crustal elements. The rheology of the mantle in the UMS models is based on the 'wet Aheim dunite' (olivine) flow law (WO) (Chopra & Paterson 1984) which is similar to that from Karato & Wu (1993) for wet olivine. This flow law is used for the sublithospheric mantle which is considered to be water-saturated ('wet'). The lithospheric mantle in the UMS experiments is assumed to be more refractory and waterpoor. The value of B* is therefore scaled to B*(WO x 10) to represent mantle lithosphere that is stronger owing to lower water fugacity (Appendix). This scaled flow law predicts effective viscosities that are intermediate between the 'wet' and 'dry' olivine-controlled flow laws of Chopra & Paterson (1984) and corresponding results in Karato & Wu (1993). The effect of the activation volume is not included in the calculation of the power-law creep flow laws (Appendix). There is no strain dependence of the material properties in the CS models described here or in our other papers on large hot orogens (Beaumont et al. 2001, 2004; Jamieson et al. 2004b, 2006). However, strain-softening is included in the UMS models, in the same parametric manner as described by Huismans & Beaumont (2003), by reducing the value of ~be#-linearly from 15 ~ to 2 ~ as the second invariant of the deviatoric strain increases from 0.5 to 1.5 (Table 2, Appendix). Strain-softening occurs in all plastic materials but there is no strain-softening during viscous flow.

Melt-weakening The most important additional property in both CS and UMS models is an extra increment of viscous weakening in the upper and middle crustal materials (those based on the WQ flow law) such that the effective viscosity decreases linearly with temperature from the dynamically determined powerlaw creep value at T = 700~ to 10~9pas at T > 750~ (Fig. 2). This weakening approximates the reduction in bulk viscosity caused by a small amount of in situ partial melt, estimated to be c. 7% at the melt connectivity transition (Rosenberg & Handy 2005). This weakening does not correspond to, and is not designed to represent, the additional decrease in effective viscosity that occurs at much larger partial melt fractions at the solid to liquid transition. The models therefore cannot be interpreted in terms of magma accumulation, transport or emplacement. The 'meltweakening' used in the present models amounts to approximately a factor of 10 decrease in effective viscosity, probably a conservative estimate for melt-weakening by a small percentage of in situ melt. The lower crust in the models does not melt-weaken because it is interpreted to be refractory intermediate granulite not prone to dehydration melting at the temperatures achieved in the models. Model materials can therefore deform according to two mechanisms - plastic or viscous flow - and in the latter case the viscosity may be further reduced by melt-weakening. In all instances, the material deforms according to the mechanism that produces the lowest level of the second invariant of the deviatoric stress for the prevailing conditions; that is, the weakest of the available flow regimes is chosen.

Density structure and isostatic compensation In CS models the crust has a uniform density (Table 1); no account is taken of density changes owing to variations in thermal expansion, melting or metamorphism. This approach is adopted so that buoyancy forces act equally on all materials and none of the flow results from differential buoyancy forces caused by density variations. The

CRUSTAL FLOW MODES changing crustal thickness is isostatically compensated by elastic flexure of a beam embedded in the model at the base of the crust. The flexural rigidity, D = 1022 Pa s, is sufficiently low that broad regions of uniform-thickness crust beneath plateau regions in the model are effectively locally compensated. Only at the transition from plateau to undeformed crust is the effect of flexure apparent. Model topography depends on the choice of crustal and mantle densities (e.g. fig. 4 of Beaumont et al. 2004). The UMS models are more dynamical than their CS equivalents and, consequently, they are more sensitive to their density structure. The upper and middle crust has a uniform density, and the lower crust has a higher density which increases over the P-T range corresponding to the granuliteeclogite transition (Table 2). With the exception of the sublithospheric mantle, all materials have a uniform volume coefficient of thermal expansion (Table 2). The sublithospheric mantle has a constant density and is, like the other materials, incompressible. The UMS models are 'isostatically' compensated at the scale of the model by their internal density structure and the associated gravitationally driven component of the flow.

Thermal model The thermal evolution is calculated by solving the heat balance equation

pCpOT/Ot + v_VT = KVZT + A

(4)

on a Eulerian finite-element mesh, where p is density, Cp is specific heat, T is temperature, t is time, v is the advection velocity of the material, K is thermal conductivity, and A is radioactive heat production per unit volume. In CS models the Eulerian finite element mesh is the same as that for the mechanical model in the crust and continues into the underlying mantle as shown in Figure 2b. The advection velocities are calculated dynamically in the crust and are prescribed kinematically in the mantle. In UMS models the heat balance equation is solved for the whole model domain using dynamically calculated velocities. In both CS and UMS models the values of K, p (thermal density) and Cp are uniform throughout the model lithosphere, resulting in uniform thermal diffusivity, K (Tables 1 and 2). The upper crust ( 0 - 2 0 km) has a uniform radioactive heat production, A1 = 2.0 IxW m -3, and the lower crust (20-35 km CS, 2 0 - 3 4 km UMS) has lower heat production, A2 = 0.75 fxW m -3 (Jamieson et al. 2002) (Tables 1, 2 and Appendix). For each model run, the initial steady-state temperature field is calculated at the scale of the model,

99

with a surface temperature of 0~ and no heat flux through horizontal side boundaries. The basal heat flux, qm = 20 mW m -z, is applied at the base of UMS models, and at the lithosphere-asthenosphere boundary, defined to coincide with the 1350~ isotherm, in CS models. For these conditions and thermal conductivity K = 2 . 0 0 W m -1 ~ -~, the initial surface heat flux qs = 71.25 m W m -z, and the Moho temperature in CS models is 704~ These values are slightly lower in UMS models (Table 2) because the crust is 34 km rather than 35 km thick. The effect of a precursor phase of oceanic subduction, included in some of our models (e.g. Vanderhaeghe et al. 2003), is not included because it has little effect on the evolving crustal temperatures and peak metamorphic conditions at the longer timescales considered here (Jamieson et al. 2002).

Surface processes In CS models, the surface processes model specifies the local erosion rate as b(t,x) = slope x f(t) x g(x), where slope is the local surface slope determined from the Eulerian finite-element mesh, f(t) is a time function, and g(x) is a 'climate' function (Fig. 2). To a first approximation g(x) is a measure of the spatial variation of aridity (0 = dry, 1 = wet) across the model. In the CS models described here f(t) is constant, but it varies in the HT-series models (Appendix; Jamieson et al. 2004b, 2006). There are no surface processes in the UMS models.

Numerical parameters For the CS models, the initial dimensions of the Lagrangian crustal grid are 5000 x 3 5 k m (501 • 41 nodes); each element is initially 10 km wide and 0.875 km deep. The Eulerian mechanical grid has 201 • 41 nodes (2000 • 35 km; crust only) and the thermal grid (crust and mantle) has 201 x 68 nodes (2000 • 96 km on undeformed pro-side, Fig. 2a). In the diagrams, deformation is displayed using a passive marker grid in which initial vertical markers are spaced at 40 km and horizontal markers at 5 km, with heavy vertical lines initially at 200 km intervals. Model times are quoted either in My (millions of years after start of model) or Ma (millions of years before end of model). The length of the timesteps, (At, is 3000 years in the CS models and 1000 years in the UMS models.

Crustal-scale model results In this section we describe results from three CS models. Model LHO-1 illustrates Mode 1,

100

C. BEAUMONT ET AL.

o;..~

o

,....

.-.2 r

9 c~ i

r

9

rm

o

~N

CRUSTAL FLOW MODES

101

..a

102

C. B E A U M O N T E T A L .

,,_.,

2 II

9

9

o o

w~

,-'5 9

.~ ZZ

~D

L'~ ii

CRUSTAL FLOW MODES

103

C ~

104

C. B E A U M O N T ET AL.

9

II

?

e,l r

V~ ?

,.....,

~8 9

,o. ~

t~

-.,~

~5_~

. ,.,.~

CRUSTAL FLOW MODES

105

C

106

C. BEAUMONT E T AL.

o~

o

,._J

. ,,,,~

,.....,

" 7 'e q~ 9

CRUSTAL FLOW MODES

107

.a

108

C. BEAUMONT E T A L .

~s

a~ T~, but not the lower crust which is assumed to be at granulite facies and not subject to melting. This model approximates radioactive self-heating, partial melting, weakening of crustal material, and a consequently increasing region of actively deforming low-viscosity crust, with a critical temperature T ~ = 750~ chosen to represent the effect of muscovite dehydration melting (Johannes & Holtz 1996). The associated thermal model (Appendix), which calculates the one-dimensional vertical thermal evolution, and the self-similar widening of the plateau (Fig. 4), is used to determine the horizontal distribution of temperature T(x, z, t) = T(z, t - x/Va). To make the results of models CI-2 and CI-3 comparable, we chose the thermal parameters of the model (Table 1) so that the mid-channel temperature also reaches the critical value after 20 million years of transition-zone thickening. The results (Fig. 6) look similar to those of model CI-2 (Fig. 4), but in this case the channel becomes thicker as the plateau crust self-heats and approaches a steady thermal state. The thermal dependence of the channel geometry provides an independent self-consistent measure for thickness of the channel, whereas in models CI-1 and CI-2 the choice of channel thickness is arbitrary. Comparison of the model topographies (Fig. 7) shows that a similar geometry to that of CI-2 is obtained for CI-3 with a smaller decrease in viscosity because the channel thickness increases from its initial value, h0, in CI-3, but remains constant in CI-2 (compare Figs 4 and 6).

Thermal and rheological consistency The previous section presented model CI-3, which combines mechanical and thermal models. These are separable into two parts. The next step is to analyse the relationship between these two parts for internal consistency, and consistency with natural systems. We first discuss the correlation between melt weakening, its consequences for the boundary between the transition zone and plateau, and then examine the effect of temperature variations along the channel in the transition zone. Lastly, we ask whether the viscosity of channel material implied by the model results is consistent with laboratory measurements of quartz-controlled rheologies.

154

S. MEDVEDEV & C. BEAUMONT

Table 1. Parameters of the reference model Variable

Geometry nc nco

i- T*. Channel thickness is initially ho. It then increases to include mid- and upper crust where T > T*. Other parameters are the same as in Figure 4. Note: (1) decrease in viscosity under transitional crust leads to plateau-like geometry; (2) self-similar widening of the plateau is characterized by the plateau advancing at velocity Va and by kinematic waveform of evolution of upper surface $3 = S3(x + Vat); (3) control parameter for the model is h 3//x (demonstrated by the effective viscosity, dotted line on (e), but here h and/z are determined independently by the thermal state; (4) channel flux varies spatially: under the plateau q ~ qb = const and in the transition zone flux decreases.

thickening timescales that are very long ( > 4 0 million years), both mechanisms predict significant heating. At intermediate timescales, which are c o m m o n in natural orogens, CI is seen to be a m o r e efficient m e c h a n i s m for crustal heating than PS. The longer-term differences are a c o n s e q u e n c e o f the relative contributions f r o m radioactivity. Pure shear thickens the crust uniforlrdy p r o d u c i n g a

than the timescale for mechanical thickening. Comparison o f the increase o f the temperature at the base o f the crust versus the timescale for crustal thickening (Fig. 8) shows that fast mechanical thickening (timescale < 10 million years) results in only moderate increases in temperature so that both mechanisms, the channel injection (CI) and pure shear (PS), (see Appendix) are ineffective. For

c,_3decrease,n-

," ~-

",

(transition zone/plateau)=25 \ ' ~ , CI-2: decrease in ~t: "~._k (transition zone/plateau)=lO0

. 0

.

.

.

, 50

.

.

.

.

, 100

.

.

.

.

, 150

.

.

. X,

.

i 200

.

.

.

.

, 250

.

.

.

.

,

300

]

350

km

Fig. 7. Comparison of the final geometries of channel injection models CI-3 (dashed line, see details on Fig. 6) and CI-2 (solid line, see details on Fig. 4). Similar geometry is obtained, but for a different viscosity decrease under the plateau owing to thicker channel in the model CI-3. See text for details.

156

S. MEDVEDEV & C. BEAUMONT 300 250

oo

~

200 ~ , , 100

~

-

i

/

I P i

5135

1~0

1~5 20 25 30 Timescale of thickening, Myr

35

40

Fig. 8. Comparison of temperature rise after thickening the crust by 30 km at a constant rate for either the channel injection mode (solid line) or pure shear mode of thickening (dashed line). The thermal state depends on the time during which thickening occurs, but CI heats crust faster than PS. Note that although the initial distribution of radioactive heat sources are equal for two models, the PS model results in a thicker, highly radioactive crustal layer (up to 44 km at the edge of plateau) than model CI-3 (which thickens the upper, high-radioactivity crust up to 22 km). See details of the thermal model in Appendix. Dotted vertical line indicates duration of thickening for the reference model. 44 km thick high-radioactivity layer (this layer is only 22 km thick in model CI), but the crust selfheats slowly, redistributing heat conductively. Channel injection heating is more efficient for thickening timescales less than 40 million years because CI injects hot material from lower in the crust and channel flow advects heat from hot crust beneath the plateau (Appendix). The third consistency criterion concerns the timing of the thermal evolution in the transition zone and requires that the base of the crust be close to the melting temperature at the time the transition zone transforms into a plateau. This transformation is interpreted to be a consequence of the decrease in viscosity owing to melt weakening.

Figure 9a shows the temperature predicted at the depth of the middle of the channel for model CI-3 after 20 million years of crustal thickening as a function of the radioactivity of the upper crust, A1, and thermal capacity, Cp (other parameter values are as given Table 1). High values of A 1 and low values of Cp produce temperatures compatible with the onset of melting. Equivalent results for the PS model (Fig. 9b) are similar but the crust is colder for the same parameter values. Rather than analysing for consistent models we reject inconsistent ones on the basis that melting most likely occurs in the temperature range 7 0 0 750~ (Johannes & Holtz 1996). Models that do not achieve this temperature range in 20 million years, and models that achieve these temperatures too quickly, were rejected leaving those with parameters in the range shown by the two shaded bands (Fig. 9c) for the CI-3 and PS models. Similar plots can be made for other combinations of parameters and duration of thickening. Particular cases can also be analysed for consistency. For example, parameter values used by Beaumont et el. (2001, 2004) for HimalyanTibetan models are inconsistent with thickening by CI-3 but consistent with PS (Fig. 9c). This result also agrees with their numerical model results which showed PS thickening in the transition zone and channel flows only penetrating this region when there is significant erosion. The parameters used by Babeyko et el. (2002) to model the Altiplano are incompatible with both modes of thickening on the 20 million years timescale and would require a minimum 35 million years interval of thickening for CI to operate and 45 million years for PS. If the Altiplano crust is expanding outward and thickening the transition zone it must have an additional source of heat not considered in our CI and PS models. This conclusion was also reached be Babeyko et el. (2002) who applied a high mantle heat flux, 60 ~W m -2, in order to achieve the necessary conditions.

Fig. 9. Contour plots of final temperature in middle of the transitional channel for channel injection (a) and pure shear (b) models. (e) Areas of thermal parameters of the crust (radioactivity of the upper crust, A1 (initial thickness 20 km) and heat capacity, Cp) that can result in a consistency between thermal and mechanical models (grey). Other parameters are the same as in reference model (Table 1). Circles illustrate thermal parameters of reference model (RM), and thermal parameters used to model Tibetan Plateau (Beaumont et el. 2004) and Andes Altiplano (Babeyko et el. 2002).

GROWTH OF PLATEAUS BY CHANNEL INJECTION

Criteria for rheological consistency of the channel below the transition zone

The results show that the extrapolated effective viscosity estimates are approximately 102tPas with a factor of 3 uncertainty at 400~ and 1019 Pa s at 750~ but with a much larger factor of uncertainty of approximately 10. If we eliminate the two lowest viscosity estimates based on the reliability of the laboratory measurements as discussed by Gleason & Tullis (1995), we predict that for a temperature range of 400-750~ calculated for the channel beneath the transition zone, associated viscosities should correspondingly vary from 1021Pas to 1019pas. Figure l la shows the plateau-edge geometry for this choice of viscosity variation compared with that predicted by the CI-3 model with constant viscosity in the transitional channel and an abrupt decrease in the viscosity at the plateau edge. Despite the admittedly large uncertainties, it can be seen that the geometry based on extrapolated laboratory data, and not including melt weakening, fails to predict the abrupt plateauedge geometry. In addition, this model predicts that topography increases steeply at the toe, much more steeply than in the CI-3 model (Fig. 1 la). This steepness is because channel flow is very inefficient at viscosities higher than 102~ Pa s, which suggests that normal crustal thickening at the foreland end of the transition zone is not rheologically consistent

In this section we investigate under what circumstances channel injection is consistent with laboratory measurements of the rheological properties of quartz-rich rocks. Natural plateaus usually exhibit an abrupt boundary between the transition zone and the plateau, which we have interpreted to reflect a significant step decrease in channel viscosity related to the onset of partial melting (Fig. 6). An alternative explanation is that the viscosity variation of quartzrich crust with temperature and strain rate is sufficient to explain the abrupt plateau-edge transition without melt weakening. The large uncertainty in the effective viscosity of quartz-rich rocks is seen from Figure 10, which shows laboratory data (table 3 in Gleason & Tullis 1995) plotted as effective viscosity versus temperature. The plot differs from the usual ones of this type in that the assumed strain rate is not constant but varies from 10-15s-1 to 10 -13 s -1, to reflect the anticipated variation in strain rate between the foreland end of the channel, where temperatures may be as low as 300-400~ and the plateau end of the transition zone, where temperatures may reach 700-800~ and the strain rates are larger.

300

. . . . . . . .

,

. . . . . . . .

,

157

. . . . . . . .

,

.

.

~ ~ - - ~- - ~-- - "~ . . . . - -

,

400

oo

500

/." 600

-

70O

.

/

/ /

/

/ / .... 800 1018

," //

//

o" z.'" /

,.

/

/" " 2 : . 11., .... 1019

"'"'"

./ "~ .... , 1 020

:_

/

increase~

//

..'"

.: .."

/

/"

....'"

/

.~

...-

/

/.;7 .Zg" //

. .."

.;f "'/"

.//"

,

/

/

/

from 10-!5s-!

/

to

1 0 " 1 3 S "1

i i --

/ I

. / .......

, 1021 viscosity, Pa s

........

, 1022

....... 1 023

Fig. 10. Dependence of effective viscosity of quartzite on temperature and strain rate. Here we approximate thermal and strain regime inside the channel in the transitional zone. Foreland end of channel is characterized by low temperature (300-400~ and low strain rate (< 1015 s-l). Plateau end of transitional channel has high temperature (700-800~ and high strain rate (up to 2 • 1013 s -1 for the reference model). Plot shows variations of effective viscosity of more than two orders of magnitude over evolution of the transitional channel. Lines show effective viscosity versus temperature, for the assumed 10-13 to 10-15 s-1 range of strain rates, compared from several sets of experimental laboratory creep data on the deformation of quartz-rich rocks. Data compiled by Gleason & Tullis (1995) and shown in their table 3.

158

S. MEDVEDEV & C. BEAUMONT

(a) Topography 6

5 E4 r4 3 2 1

m

"" " ~ ~

m

Exponential decrease of channel viscosity

/~~L-~

m

m

m

CI-3 0

100

5O (b) Sichuan Basin margin

~

m

150

200

250

m

5 ...... ....

3 4

:_ ,_,r

.,,~./k~

-

~ ,

2 1

~ , ,

( ,

,

%

5

C1-1:#10 = 3-102~ Pa.s = 102, Pa.s)

^

CI-3: PlO = 3:1027 P a . s . ~ . ~ (!.t15 = 1021 Pa.s) '--*~",,."~.._

_

9. . . . . . " y ~ , , ~ - _ . ~

50 1oo (r Southeast margin (south of Sichuan Basin) 6

,IA

E4 N'3 2 1

. . = a t , IL

500

1000

-

.._.

150

200

250

1~

CI-3: glO = 2.10 ,7 Pa.s . ~ (~t15 = 6.7.1017 Pa.s) 0

CI-3:#10 = 1.2.1020 Pa.s (~15 = 4 "102~ Pa.s)

~

....

1500 x. km

2000

~g~.~,.~. 2500

3000

Fig. 11. Comparison of model CI-3 results with experimental and natural examples. (a) Result of model CI3 (constant viscosity along channel and melt weakening at the edge of plateau, solid line) differs significantly from model with an exponential decrease of channel viscosity during evolution of the orogen (dashed line). The latter model (see text) is based on laboratory results presented in Figure 10. (b) Comparison of model results with topographic profile across Sichuan Basin margin of Tibetan Plateau (topographic data modified from Clark & Royden 2000, fig. 4c). Results for CI-3 (/z = 1.2 x 1020 Pa s for channel thickness 10 km, solid line) and CI-1 (/.L= 3 x 1020 for h = 10 km, dashed line; see also fig. 4c in Clark & Royden 2000) almost coincide in the transitional zone (between x = 100 and 200 kin) and match well the natural topography, but differ by a factor of c. 3, the value of the estimated viscosity. Model CI-1 does not account for the development of the plateau. For comparison, the dash-dot line shows result for model CI-3 using the value of the viscosity of the channel material in the transition zone estimated by CI-1. (c) Comparison of models results with topographic profile across the SE margin of the Tibetan Plateau (topographic data modified from Clark & Royden 2000, fig. 4a). The result obtained by model CI-3 (solid line) compares well with the natural example, but the predicted viscosity of the channel material is unreasonably low unless the lower crust was exceptionally hot at the time of channel injection. Viscous thickening (VT, dashed line) produces a more reasonable estimate of the average viscosity of the entire orogenic crust.

with the channel injection model, and that an alternative mechanism needs to be considered. A similar question of rheological consistency arises in regard to the estimates by Clark & Royden (2000). Whereas estimated viscosities for the channel material under the margins with short transitional zones (e.g. Sichuan Basin margin of Tibetan Plateau, Fig. 1 lb) m a y be consistent with the laboratory data, the estimate of channel viscosities of 10 e8 Pa s for a 2000 k m wide region of the SE Tibetan Plateau margin south of the Sichuan

Basin (and even lower viscosity predicted by the more complete CI-3 model, Fig. 1 l c) is questionable. Such low viscosities require temperatures of approximately 800~ even for the lower effective viscosity extrapolations of the laboratory data (Fig. 10). The results imply that the channel injection explanation for crustal thickening only applies if the crust in this region was already anomalously hot prior to channel injection, which raises the related question of whether it was already anomalously thick.

GROWTH OF PLATEAUS BY CHANNEL INJECTION

159

which is similar to Equations 1 and 2, except the corresponding flux

Behaviour of the upper crust We complete the analysis with a preliminary description of other mechanisms, viscous thickening (VT) and plastic translation (PT) (Fig. lc and d), which may also contribute to deformation and crustal thickening of the transition zone but differ from channel injection.

qv --

pgh 30S 3 3ix

Ox

(8)

results from integration of the velocity profile, which is

Viscous thickening model The CI models assume that crust above the channel is much stronger than the channel material and therefore does not deform during injection. However, if this is not correct the deformation can instead be evenly distributed in the transitional crust. Simple models of this type treat the deformation as viscous, are based on the simple shear thin-sheet approximation (e.g. Buck & Sokoutis 1994; Royden 1996; Medvedev 2002), and assume that the upper surface of the crust is deformable and stress-free. The remaining boundary conditions are the same as those for the CI model. The VT model is of this type (Fig. 12) and the mass balance gives 0S3 ~x 0~- + 9 = 0

(7)

Vx(z) = - pg(z - Si-)(h2 + S + - z) 0S3 2ix Ox

(9)

for a uniformly viscous crust and assuming lithostatic pressure. The resulting system of equations is similar to those for the CI model (Equations 1-3), except for the denominator 3, instead of 12, in the expressions for the fluxes. This difference occurs because the upper surface in the VT model is free to deform. Note also that the VT equations are non-linear (h2 is a linear function of $3) and cannot be solved analytically. The model result (Fig. 12) is shown for the viscosity 1.5 x 1022 Pa s which produces the same transition-zone geometry as models CI-2 and CI-3 with channel material viscosity of 102o Pa s. The difference of two orders of magnitude in viscosity required to produce the same topography for the transition zone demonstrates the sensitivity of the behaviour to the mode of deformation. This result implies that for natural examples the type of deformation must be known, in addition to the geometry, before the viscosity can be accurately estimated. An example is the very wide (c. 2000 km) transition zone SE of the Sichuan basin for which Clark & Royden (2000) estimated a viscosity of 1078 Pa s based on the CI model (model CI-3 predicts an even lower value for viscosity; Fig. l l c ) . The corresponding estimate from the VT model is 5 • 1019 Pa s, which may be more realistic if contraction at the surface cannot be ruled out. S t a b i l i t y o f the u p p e r c r u s t

Fig. 12. Results for viscous thickening model. Special here has /z = 1.5 x 1022 Pa s. Flux case considered ' of material into the system, qb = 10-6 m3 s -~ per m along-strike (as for Fig. 3) results in plateau margin elevated to 5.5 km in 20 million years. The geometry of the transitional zone is similar to one by CI-3 (Fig. 7), but rheological parameters of these two models differ significantly.

The last mode of deformation in the transition zone considers the plastic (brittle) failure and translation (PT) of the crust above the channel injection zone. Here we estimate the stability of the upper crust, which allows us to distinguish when the CI and PT modes will operate. Lehner (2000) considered the stability of frictional overburden coupled to a viscous substratum and used the thin-sheet approach for the channel flow (Equation 1) and the small angle approximation for the overburden. Gemmer et al. (2004) presented a simplified version of the theory for systems with constant-thickness channels and they also derive an expression for the velocity when the crust is unstable. They showed that their results

160

S. MEDVEDEV & C. BEAUMONT

are in good agreement with full 2D numerical calculations. Consider the forces acting on a block of the crust above the channel in the transition zone (Fig. 13, area surrounded by bold line). The horizontal balance of forces is:

Fp - Fo - Fnx -I- Fr = 0

(10)

where Fp and Fo are the normal forces acting on the block from the edge of plateau and from the foreland respectively, F~- is the traction on the base of the block produced by the Poiseuille flow in the transitional channel, and F,,x is the horizontal component of the buoyancy force. If the forces in Equation 10 cannot be balanced, the block becomes unstable and starts to move. The forces can be estimated in the following ways. The wedge-shaped block, thickness hw(x), has two parts, an upper uniform frictional-plastic layer, thickness hb, and a lower approximately triangularshaped viscous region above the channel and below the plastic layer. Therefore S (Xp)

Fp

I

Js+(xp)6rxx dz

Fpb + Fpv

(11)

where Fpb and Fpv are the integrals of the horizontal stress within the plastic and viscous parts, respectively, at x = Xp. Integration of the stress under the small-angle assumption for the geometry of the block, and assuming that the principal stresses are horizontal and vertical and the pressure is equal to the lithostatic pressure, gives (Gemmer et aL 2004): Feb > Fp = pghZk/2 and Fpv = pg(h2p - h2)/2; k = (1 - sin~b)/(1 + s i n & ) is the coefficient of passive earth pressure (Lehner 2000), where r is

the internal angle of friction of the overburden, and hwp = hw(xp) is the thickness of the system at the edge of plateau. F~p is the limiting value of Fpb when the system is on the verge of failure, or has failed. Under the equivalent assumptions, Fo < F~ = pghZk-1/2. The horizontal component of the normal stress at the base of the system can be calculated following Medvedev & Podladchikov (1999) by estimating the buoyancy by integrating the pressure along the base of the block

F,x =

>

ost (x) d pghw(x) ~ z

o

p2g jx p Oh2 d pZg (h 2 -2~2 o~-Xz=2~" w~-h~) (12) Here we use the constraint that the boundary S + is parallel to the base of the crust (Fig. 13), and the Airy isostatic condition

os+ p Ohw Ox --Pro OX

where Pm is the density of the lithospheric mantle (Table 1). The traction force from the viscous flow in the channel is determined by integrating the basal drag caused by the channel injection along the length the block (Gemmer et al. 2004): pgh Fr = -~-(S3(xp) - S3(xo))

_ pCPgh-- (hwp - ht,)

(14)

where 9 is the isostatic amplification factor (Table 1). Substitution of the force estimates into Equation 10 gives the stability condition for the system:

h~(k-k-')+~(h2p-h2)+~h(hwp-hh) 1400 km south of the suture. Material tracking indicates little or no mixing of diverse crustal elements in the exhumed region of the model 'GHS', which is derived from originally contiguous materials that are transported together in the top of the channel flow zone. These results are compatible with provenance data indicating a clear distinction between GHS and LHS protoliths, with the GHS originating from a more distal position (relative to cratonic India) than the LHS. In model H T l l l , domes formed between the suture and the orogenic front are cored by 'Indian' middle crust similar to the 'GHS', consistent with data from the north Himalayan gneiss domes. Material tracking shows that plutons generated south of the suture should have 'Indian' crustal signatures, also compatible with observations. Model 'GHS' pressure-temperature-time (P-T-t) paths pass through the dehydration melting field between 30 and 15 Ma, consistent with observed leucogranite ages. Finally, exposure of midcrustal 'GHS' and 'LHS' material at the model erosion front is consistent with the observed appearance of sedimentary detritus in the Lesser Himalaya. We conclude that channel flow model results are compatible with provenance data from the Himalaya and southern Tibet.

Numerical models for channel flow in the H i m a l a y a n - T i b e t a n system (e.g. B e a u m o n t et al. 2001) predict that low-viscosity middle crust has flowed outward from beneath the Tibetan Plateau, and has been e x h u m e d between the Main Central Thrust zone (MCT) and South Tibetan Detachment system (STD) in response to focused denudation at the erosion front. Model results are compatible with a number of first-order tectonic and metamorphic features of the orogen (Beaumont et al. 2001, 2004; Jamieson et al. 2004). However, previous work did not address in detail the distribution of material from different sources within the model orogen. Material tracking in the models can be used to predict the provenance of rocks now exposed in the Greater Himalayan (GHS) and Lesser Himalayan (LHS) sequences and in the north Himalayan gneiss (NHG) domes, as well as the likely sources of Miocene leucogranites throughout the region. Conversely, provenance data from the orogen provide an important test of the models and additional constraints on material flow patterns within the model orogen. Here we present results from two channel flow models (HTI and HT111) that display contrasting tectonic

styles in the model region corresponding to the Himalaya and southern Tibet and that are reasonably consistent with a range of first-order observations from the orogen. The purpose of this paper is to compare the distribution of crustal material at and below the model surfaces with provenance data from metamorphic, plutonic and sedimentary rocks in the orogen. In order to facilitate comparison between model results and data from the H i m a l a y a n - T i b e t a n orogen, the following nomenclature has been adopted in this paper. Model times are given in Ma, meaning 'millions of years before the end of the model', equivalent to 'millions of years before present' in nature. Model features appear in quotation marks; corresponding features of the real H i m a l a y a n - T i b e t a n system do not. The pro-side of the model is referred to as 'India' (south) and the retro-side of the model is referred to as 'Asia' (north). The upper crust above the model orogen is referred to as 'Tethyan' material. The protolith boundary between incoming pro-crust and the outflowing channel is referred to as the M C T (analogous to the Main Central Thrust zone), and the boundary between the extruded channel

From: LAW, R. D., SEARLE,M. P. & GOD/N,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 165-182. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

166

R.A. JAMIESON ETAL.

and the overlying upper crust is referred to as the STD (analogous to the South Tibetan Detachment system; Jamieson et al. 2004). The initial boundary between pro-side and retro-side crust, interpreted to be equivalent to the Indus-Yarlung Tsangpo suture zone, is referred to simply as the suture. The terms distal and proximal refer to relative distance from the pro-ward end of the model ('cratonic India'); pro-side distal material originates closer to the model suture than proximal material.

Model design Models HT1 and H T l l l (Fig. 1, Table 1) were selected from a series of related models (e.g. Beaumont et al. 2001, 2004) that exhibit features relevant to the Himalayan-Tibetan system. The design of model HT1 and its thermal-mechanical evolution were discussed in detail by Beaumont et al. (2004) and Jamieson et al. (2004); additional information is provided in Beaumont et al. (2006, appendix). The numerical formulation and calculation procedure have been described elsewhere (Fullsack 1995; Beaumont et al. 2004, 2006). Model H T l l l is identical to HTI except that the upper crust contains an embedded weak layer; as detailed below, the weaker upper crust has a significant effect on model evolution. Both models were run for 54 million years, with the convergence velocity, Vp = 5 cm a -l, partitioned between S point (subduction) advance at Vs = 2.5 cm a-J and subduction at Vp - V s = 2 . 5 c m a -1 (Fig. 1; Beaumont et al. 2004). The lower, middle and upper crustal layers (Fig. 1, Table 1) are laterally homogeneous. In model HT1, the upper crustal layer (0-10 kin) has a wet quartzite ( B * ( W Q ) ) rheology (Gleason & Tullis 1995; details in Table 1) with an effective internal angle of friction, ~beff= 5 ~ In model HT111, a B * ( W Q ) layer with ~b4j= 2 ~ has been embedded in the upper crust between 4.5 and 7.0kin (Fig. 1). This layer can be thought of as weak sedimentary rock with high pore fluid pressure. In both models, the rheology of the middle crustal layer ( 1 0 - 2 5 k m ) is equivalent to B * ( W Q x 5) with ~bey= 15 ~ (Table 1, Fig. 1), and the lower crustal rheology (25-35 km) is dry Maryland diabase ( B * ( D M D ) ; Mackwell et al. 1998), also with ~be~= 15~ The middle and lower crustal layers can be thought of as quartzo-feldspathic (granitoid and/or metasedimentary) material and intermediate granulite, respectively (Beaumont et al. 2006, appendix). The strong lower crust is subducted and does not influence pro-side model evolution, although it extends as an undeformed 10 km layer along the base of the orogen to the S point. Both model orogens therefore consist mainly of

thickened middle and upper crust, and their thermal evolution is controlled by crustal heat production (A1 = 2.0 oLW m -3, 0 - 2 0 kin; A2 = 0.75 ~ W m -3, 20-35 km). Material flow in the models is tracked using contrasting grey shades for mid-crustal blocks (Fig. 2), which are initially 200 km wide and have identical material properties. In both models, the effective viscosity is reduced linearly from its flow law value at 700~ to 1019 Pa s at T > 750~ (Fig. 1" Beaumont et al. 2001). For convenience, we refer to this as 'melt weakening', because the effect is probably comparable to that produced by the presence of a small amount (_ 3 km) pluton suites like the > 3 0 km long Manaslu granite in central Nepal (e.g. P~cher et al. 1991) or the > 50 km long Khula-Kangri Monlakarchung Passalum suite in western Bhutan (Edwards et al. 1999), where the HHC width, in addition, is quite thin ( < 1 0 k m ) . Such plutonism corresponds to a volume increase of several tens of per cent. In pure-shear-dominated flow, such a volume change will have a significant effect on extrusion of the slab (compare stretch in x in Fig. 3b). However, this can only happen if thinning of the slab is very

DILATANCY EFFECTS ON EXTRUSION strong. Meanwhile, if flow is simple-shear-dominated, then the slab can maintain its thickness with low sensitivity to volume change. However, the effective shear strain is strongly influenced by that volume change (Fig. 3, lower). If there has been much higher local volume increase in the central Nepal and western Bhutan areas, this will force the slab to locally deform with a higher contribution of pure shear (given that in the Himalaya the underthrusting of India is more or less constant along the whole orogen, i.e. the effective shear strain is similar). An along-orogen switching of pure-versus simple-shear-dominated flow can be checked, for example, with the extensive suites of pre- and syntectonic dykes (abundant around the granites and throughout the upper HHC) using Talbot's method of estimating the rotational and stretching behaviour relative to the strain ellipsoid (Talbot 1970). Kinematic v o r t i c i ~ values Kinematic vorticity analyses in the HHC have dominantly given results for Wk of between 0.7 and 1. Some of these studies have also made a rough estimate of vorticity integrated across the whole slab (Vannay & Grasemann 2001), while others focused on thin portions of otherwise thick shear zones within the slab (Grasemann et al. 1999; Law et al. 2004; Carosi et al. 2006; Jessup et al. 2006). Under such simple-shear-dominated flows the thinning and extrusion component is quite insensitive to volume change. However, the effective shear component shows a strong dependence upon volume change. If strain is largely decoupled from the STDS, this could explain why the shear strain of the STDS varies dramatically along the strike of the orogen. Stretching and thinning o f the slab The stretching component of the slab is difficult to decipher along the MCT. Here, the total or minimum displacement that can partly be reconstructed by either cross-section balancing or by simple inspection of geological maps (tectonic windows, ldippen etc.) is represented by the underthrusting component of India beneath the HHC (the 'slab') and by the stretching component of the slab due to general shear and possible volume addition. Along the STDS, however, just the stretching component of the HHC is recorded, showing an increasing displacement towards the foreland and topographic surface of the Earth (Vannay & Grasemann 2001). Therefore, the significant along-strike variations in the extensional displacement observed along the STDS (e.g. very little in NW India, large in central Nepal) could be a

195

function of the present-day structural level exposed by erosion, which is consistent with the fact that the STDS is not a continuous structure along the strike of the Himalayan orogen. Therefore estimates of the local displacement on the STDS, including the exposed metamorphic grade of the hanging wall and the footwall, can be used to estimate the stretching component of the slab. The thinning of the slab during extrusion is probably the most difficult parameter to decipher from field data alone, because the original thickness of the HHC is unknown. However, it is important to note that even the present-day thickness of the HHC can vary from a few up to 40 km in some sections (e.g. Jain & Manickavasagam 1993). Although there may locally be more than one main component to the MCT system, and the actual HHC or principally extruded slab may not always be clearly identified, there are certainly significant along-strike changes of the slab width. This can be interpreted as variation in either the initial thickness of the HHC, or in the magnitude of the thinning component (we recall that the stretching component of the slab was quite variable). If the finite strain can be estimated, our simple kinematic relationship between yp and AAe (Fig. 3) can be used to derive realistic values for thinning of the slab.

Conclusions Our deformation tensor derived from a redefined kinematic dilatancy term for non-isochoric distortion in the velocity gradient tensor has allowed us to examine kinematic implications of volume change in an extruding orogenic mid-crustal slab. For comparable flow types with a fixed stretchingrate factor, especially at low strains, stretches of material lines parallel and perpendicular to the slab are very sensitive to area change in pureshear-dominated flow, while effective shear strain is very sensitive to area change in simple-sheardominated flow. For non-isochoric distortion where either the thinning of the slab is constrainable or the slab is of a fixed width (channel flow), the orientation of the non-stretching eigenvector, if known from field data, would allow constraint of a possible volume change range if finite strain is known, or vice versa. For geologically reasonable stretching-rate factors, pure-shear-dominated extrusion with constant slab thickness can only achieve naturally observed extrusion amounts with massive (doubling, tripling) volume increase and so is unlikely to play a significant role in collisional orogenesis. Volume increase in the slab appears to fluctuate along the Himalayan orogen and should be associated with heterogeneities in the

196

B. GRASEMANN ETAL.

corresponding local slab stretch values and effective shear strain. V o l u m e change is a key, hitherto rarely considered c o m p o n e n t of extrusional flow regimes. M.E. is supported through a EU-Marie Curie Research Fellowship (contract no. HPMF-CT2002-01703). G.W. is supported through the Austrian FWF (project no. P-15668). We thank reviewers D. De Paor and A. Bobyarchick for excellent reviews, editor R. Law for detailed editorial and review duties, and all the members of the Structural Processes Group, Vienna, especially M. Ebner, for discussions and manuscript readthrough. M.E.'s interests in Himalayan granite generation were first inspired by K. Douglas Nelson; we dedicate this work to his memory.

References BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2002. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAM1ESON, R. A., NGUYEN, M. H. & MEDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109(B06406). DOI" 10.1029/2003JB002809. BOBYARCHICK, A. R. 1986. The eigenvalues of steady flow in Mohr space. Tectonophysics, 122, 35-51. BURCHFIEL, B. C. & ROYDEN, L. H. 1985. North south extension within the convergent Himalayan region. Geology, 13, 679-682. BURG, J. P. 1983. Tectogdnkse comparde de deux segments de chafne de collision: le Sud du Tibet (suture du Tsangpo); la chafne hercynienne en Europe (suture du Massif-Central). PhD thesis, Universit6 Montpellier, France. CAROS1, R., MONTOMOLI, C., RUBATTO, D. & VISONA, D. 2006. Normal-sense shear zones in the core of the Higher Himalayan Crystallines (Bhutan Himalaya): evidence for extrusion? In: LAW, R. D., SEARLE, M. P. & GOD1N, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications 268, 425-444. CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. DANIEL, C. G., HOLL1STER, L. S., PARRISH, R. R. & GRUJIC, D. 2003. Exhumation of the Main Central Thrust from Lower Crustal Depths, Eastern Bhutan Himalaya. Journal of Metamorphic Geology, 21, 317-334. DE PAOR, D. G. & MEANS, W. D. 1984. Mohr circles of the First and Second Kind and their use to represent tensor operations. Journal of Structural Geology, 6, 693 -701. DEBON, F., ZIMMERMANN, J. L., LIU, G. J. & CH, X. 1985. Time relationship between

magmatism, tectonics and metamorphism in three plutonic belts in Southern Tibet: new K-Ar data. Geologische Rundschau, 74, 229-236. DEBON, F., LE FORT, P., SHEPPARD, S. M. F. & SONET, J. 1986. The four plutonic belts of the Transhimalaya-Himalaya; a chemical, mineralogical, isotopic, and chronological synthesis along a Tibet-Nepal section. Journal of Petrology, 27, 219-250. DENIEL, C., VIDAL, P., FERNANDEZ,A., LE FORT, P. & PENCAT, J. J. 1987. Isotopic study of the Manaslu granite (Himalaya, Nepal): interferences on the age and source of Himalayan leucogranites. Contributions to Mineralogy and Petrology, 96, 78-92. EDWARDS, M. A. & HARRISON, T. M. 1997. When did the roof collapse? Late Miocene N-S extension in the High Himalaya revealed by Th-PB monazite dating of the Khula Kangri granite. Geology, 25, 543 -546. EDWARDS, M. A., KIDD, W. S. F., LI, J., YUE, Y. & CLARK, M. 1996. Multi-stage development of the southern Tibet detachment system near Khula Kangri. New data from Gonto La. Tectonophysics,

260, 1-19. EDWARDS, M. A., PI~CHER, m., KIDD, W. S. F., BURCHFIEL, B. C. & ROYDEN, L. H. 1999. Southern Tibet detachment at Khula Kangri, Eastern Himalaya: a large-area, shallow detachment stretching into Bhutan? The Journal of Geology, 107, 623-631. ENGLAND, P. C. • HOLLAND, T. J. B. 1979. Archimedes and the Tauern eclogites: the role of buoyancy in the preservation of exotic eclogite blocks. Earth and Planetary Science Letters, 44, 287-294. FOSSEN, H. & TIKOEF, B. 1993. The deformation matrix for simultaneous simple shearing, pure shearing and volume change, and its application to transpression-transtension tectonics. Journal of Structural Geology, 15, 413-422. FOSSEN, H. & TIKOFE, B. 1997. Forward modeling of non-steady-state deformations and the 'minimum strain path'. Journal of Structural Geology, 19, 987-996. GANSSER, A. 1964. Geology of the Himalayas. Interscience Publishers, London. GANSSER, A. 1983. Geology of the Bhutan Himalayas. Birkhauser Verlag, Basel. GRASEMANN, B. & VANNAY, J.-C. 1999. Flow controlled inverted metamorphism in shear zones. Journal of Structural Geology, 21, 743-750. GRASEMANN, B., FRITZ, H. & VANNAY, J. C. 1999. Quantitative kinematic flow analysis from the Main Central Thrust Zone (NW-Himalaya, India): implications for a decelerating strain path and the extrusion of orogenic wedges. Journal of Structural Geology, 21, 837-853. GRASEMANN, B., STI3"WE, K. & VANNAY, J.-C. 2003. Sense and non-sense of shear in flanking structures. Journal of Structural Geology, 25, 19-34. GRUJ1C, D. 2006. Channel flow and continental tectonics: an overview. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision

DILATANCY EFFECTS ON EXTRUSION

Zones. Geological Society, London, Special Publication, 268, 25-38. GUILLOT, S. & Le FORT, P. 1995. Geochemical constraints on the bimodal origin of High Himalayan leucogranites. Lithos, 35, 221-234. GUILLOT, S., LE FORT, P., P~CHER, A., BARMAN,M. R. & APRAHAMIAN,J. 1995a. Contact metamorphism and depth of emplacement of the Manaslu granite (Central Nepal). hnplications for Himalayan orogenesis. Tectonophysics, 241, 99-119. GUILLOT, S., PI~CHER, A. & LE FORT, P. 1995b. Contrrles tectoniques et thermiques de la mise en place des leucogranites himalayens. Comptes Rendus de l'Academie des Sciences, 320 (Ser. IIa), 55-61. HARRIS, N. & MASSEY, J. 1994. Decompression and anatexis of himalayan metapelites. Tectonics, 13, 1537-1546. HARRISON, T. M., GROVE, M., LOVERA, M. & CATLOS, E. J. 1998. A model for the origin of Himalayan anatexis and inverted metamorphism. Journal of Geophysical Research, 103, 27,01727,032. HAUCK, M. L., NELSON, K. D., BROWN, L. D., ZHAO, W. & Ross, A. R. 1998. Crustal structure of the Himalayan orogen at 90 ~ east longitude from Project INDEPTH deep reflection profiles. Tectonics, 17, 481-500. HENRY, P., LE PICHON, X. & GOFFE, B. 1997. Kinematic, thermal and petrological model of the Himalayas: constraints related to metamorphism within the underthrust Indian crust and topographic elevation. Tectonophysics, 273, 31-56. HIPPERTT, J. F. 1998. Breakdown of feldspar, volume gain and lateral mass transfer during mylonitization of granitoid in a low metamorphic grade shear zone. Journal of Structural Geology, 202, 175-193. HODGES, K. V., PARRISH, R. R., HOUSH, T. B., Lux, D. R., BURCHFIEL, B. C., ROYDEN, L. H. & CHEN, Z. 1992. Simultaneous Miocene extension and shortening in the Himalayan orogen. Science, 258, 1466-1470. INGER, S. & HARRIS, N. 1993. Geochemical constraints on leucogranite magmatism in the Langtang Valley, Nepal Himalaya. Journal of Petrology, 34, 345-368. JAIN, A. K. & MANICKAVASAGAM, R. M. 1993. Inverted metamorphism in the intracontinental ductile shear zone during Himalayan collision. Geology, 21, 407-410. JAMIESON, R. A., BEAUMONT, C., MEDVEDEV, S. & N~UYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan Orogen. Journal of Geophysical Research, 109, B06407. DOI: 10.1029/2003JB002811. JESSUP, M. J., LAW, R. D., SEARLE,M. P. & HUBBARD, M. S. 2006. Structural evolution and vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/Nepal: implications for orogen-scale flow partitioning. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion

197

and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 379-413. JIANG, D. 1998. Forward modeling of non-steady-state deformations and the 'minimum strain path': Discussion. Journal of Structural Geology, 20, 975 -977. KLEMPERER, S. 2006. Crustal flow in Tibet: geophysical evidence for the physical state of Tibetan lithosphere and inferred patterns of active flow. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications 268, 39-70. LAW, R., SEARLE, M. P. & SIMPSON, R. L. 2004. Strain, deformation temperatures and vorticity of flow at the top of the Greater Himalayan Slab, Everest Massif, Tibet. Journal of the Geological Society of London, 161, 305-320. LE FORT, P. 1975. Himalayas: The collided range. Present knowledge of the continental arc. American Journal of Science, 275, 1-44. LE FORT, P., CUNEY, M., DENIEL, C., FRANCELANORD, C., SHEPPARD, S. M. F., UPRETI, B. N. & VmAL, P. 1987. Crustal generation of the Himalayan leucogranites. Tectonophysics, 134, 39-57. MAKOVSKY, Y., KLEMPERER, S. L., RATSCHBACHER, L., BROWN, L. D., LI, M., ZHAO, W. ~; MENG, F. 1996. INDEPTH wide-angle reflection observation of P wave to S wave conversion from crustal bright spots in Tibet. Science, 274, 1690-1691. MANCKTELOW, N. S. 1995. Nonlithostatic pressure during sediment subduction and the development and exhumation of high pressure metamorphic rocks. Journal of Geophysical Research, 100, 571-583. MOI-IR, O. 1882. fQber die Darstellung des Spannungszustandes und des Deformationszustandes eines K6rperelementes und fiber die Anwendung derselben in der Festigkeitslehre. Der Zivilingenieur, 28, 113-156. MONTEL, J. M. 1993. A model for monazite/melt equilibrium and application to the generation of granitic magmas. Chemical Geology, 110, 127-146. MUNSON, B. R., YOUNG, D. F. & OKnsm, T. H. 2002. Fundamentals of Fluid Mechanics. John Wiley & Sons, New York. NELSON, K. D., ZHAO, W., BROWN, L. D. eraL. 1996. Partially molten middle crust beneath southern Tibet: synthesis of Project INDEPTH results. Science, 274, 1684-1688. O'HARA, K. D. 1988. Fluid flow and volume loss during mylonitization: an origin for phyllonite in a n overthrusting setting, North Carolina, U.S.A. Tectonophysics, 156, 21-36. PASSCHIER, C. W. 1987. Efficient use of the velocity gradient tensor in flow modelling. Tectonophysics, 136, 159-163. PASSCHIER, C. W. 1988. Analysis of deformation paths in shear zones. Geologische Rundschau, 77, 309-318.

198

B. GRASEMANN ETAL

PASSCHIER, C. W. 1991. The classification of dilatant flow types. Journal of Structural Geology, 13, 101-104. PATIlqO-DOUCE, A. E. & HARRIS, N. B. W. 1998. Experimental constraints on Himalayan anatexis. Journal of Petrology, 39, 689-710. PATIlqO-DOUCE, A. E. & JOHNSTON, A. D. 1991. Phase equilibria and melt productivity in the pelitic system: implications for the origin of peraluminous granitoids and aluminous granulites. Contributions to Mineralogy and Petrology, 107, 202-218. P~CHER, A., BOUCHEZ, J. L. & LE FORT, P. 1991. Miocene dextral sheafing between Himalaya and Tibet. Geology, 19, 683-685. PROVOST, A., BUISSON, C. & MERLE, O. 2004. From progressive to finite deformation and back. Journal of Geophysical Research, 109, B02405. DOI: 10.1029/2001JB001734. RAMBERG, H. 1975. Particle paths, displacement and progressive strain applicable to rocks. Tectonophysics, 28, 1-37. RAMSAY, J. G. & GRAHAM, R. H. 1970. Strain variation in shear belts. Canadian Journal of Earth Sciences, 7, 786-813. RATSCHBACHER, L., FRISCH, W. & L1NZER, H. G. 1991. Lateral extrusion in the Eastern Alps. Part 2: structural analysis. Tectonics, 10, 257-271. RING, U. 1999. Volume loss, fluid flow, and coaxial versus noncoaxial deformation in retrograde, amphibolite facies shear zones, northern Malawi, east-central Africa. Geological Socie~ of America Bulletin, 111, 123-142. ROYDEN, L. H. 1996. Coupling and decoupling of crust and mantle in convergent orogens: Implications for strain partitioning in the crust. Journal of Geophysical Research, 101, 17,679-17,692. ROYDEN, L. H. & BURCHFIEL, B. C. 1987. Thinskinned N-S extension within the convergent Himalayan region: gravitational collapse of a Miocene topographic front. In: COWARD, M. P., DEWEY, J. F. & HANCOCK, P. L. (eds) Continental Geological Society, Extensional Tectonics. London, Special Publications, 28, 611-619. SCAILLET, B. & SEARLE, M. P. 2006 Mechanisms and timescales of felsic magma segregation, ascent and emplacement in the Himalaya. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 293-308. SCAILLET, B., PI~CHER,A., ROCHETTE, P. & CHAMPENO1S, M. 1995. The Gangotri granite (Garhwal Himalaya): laccolithic emplacement in an extending collisional belt. Journal of Geophysical Research, 100, 585-607. SCAILLET, B., HOLTZ, F., P1CHAVANT, M. & SCHMIDT, M. 1996. Viscosity of Himalayan ieucogranites: Implications for mechanisms of granitic magma ascent. Journal of Geophysical Research, 101, 27,691-27,699. SCHNEIDER, D. A., EDWARDS, M. A., KIDD, W. S. F., KHAN, A. M., SEEDER, L. & ZEITLER, P. K. 1999. Tectonics of Nanga Parbat, western Himalaya: Synkinematic plutonism within the doubly

vergent shear zones of a crustal-scale pop-up structure. Geology, 27, 999-1002. SCHWERDTNER, W. M. 1982. Calculation of volume change in ductile band structures. Journal of Structural Geology, 4, 57-62. SEARLE, M. P., NOBLE, S. R., HURFORD, A. J. & REX, D. C. 1999. Age of crustal melting, emplacement and exhumation history of the Shivling leucogranite, Garhwal Himalaya. Geological Magazine, 136, 513-522. SIMPSON, C. & DE PAOR, D. 1993. Strain and kinematic analysis in general shear zones. Journal of Structural Geology, 15, 1-20. SRIVASTAVA, H. B., HUDLESTON, P. & EALEY III., D. 1995. Strain and possible volume loss in a highgrade ductile shear zone. Journal of Structural Geology, 17, 1217-1231. TALBOT, C. J. 1970. The minimum strain ellipsoid using deformed quartz veins. Tectonophysics, 9, 46-76. TIKOFF, B. & FOSSEN, H. 1993. Simultaneous pure and simple shear: the unifying deformation matrix. Tectonophysics, 217, 267-283. VANNAY, J.-C. & GRASEMANN, B. 2001. Himalayan inverted metamorphism and syn-convergence extension as a consequence of a general shear extrusion. Geological Magazine, 138, 253-276. VIELZEUF, D. & HOLLOWAY,J. R. 1988. Experimental determination of the fluid-absent melting relations in the pelitic system. Contributions to Mineralogy and Petrology, 98, 257-276. V~SSERS, R. L. M. 1989. Asymmetric quartz c-axis fabric and flow vorticity: a study using rotated garnets. Journal of Structural Geology, 11, 231-244. WALLIS, S., PLATT, J. P. & KNOTT, S. D. 1993. Recognition of synconvergence extension in accretionary wedges with examples from the Calabrian Arc and the Eastern Alps. American Journal of Science, 293, 463-495. WALL1S, S. R. 1992. Vorticity analysis in a metachert from the Sanhagwa Belt, SW Japan. Journal of Structural Geology, 14, 271-280. WATTS, D., HARRIS, N., GAINES, J. S. ET AL. 2005. Mapping granite and gneiss in domes along the North Himalayan antiform with ASTER SWIR band ratios. Geological Society of America Bulletin, 117, 879-886. WHITTINGTON, A. & TRELOAR, P. J. 2002. Crustal anatexis and its relation to the exhumation of collisional orogenic belts, with particular reference to the Himalaya. Mineralogical Magazine, 66, 53-91. WIESMAYR, G. & GRASEMANN, B. 2002. Eohimalayan structural evolution of the fold and thrust belt in the Tethyan Himalaya (Spiti, N-India). Tectonics, 21, 8, 1-18. WIESMAYR, G., EDWARDS, M. A., MEYER, M., KIDD, W. S. F., LEBER, D., HAEUSLER, H. & WANGDA, D. 2002. Evidence for steady fault-accommodated strain in the High Himalaya; progressive fault rotation of the southern Tibet detachment system in NW Bhutan. In: DE MEER, S., DRURY, M. R., DE DRESSER, J. H. P. & PENNOCK, G. M. (eds)

DILATANCY EFFECTS ON EXTRUSION

Deformation Mechanisms, Rheology and Tectonics; Current Status and Future Perspectives, Geological Society, London, Special Publications, 200, 371-386. Wu, C., NELSON, K. D., YuE, Y., LI, J., KIDD, W. S. F. & Et)wagI)S, M. 1995. Yadong cross structure and South Tibetan detachment, southern Yadong-Gulu Rift, Tibet. Geological Society of America, Abstracts with programs, 27, 334-335. Wu, C., NELSON, K. D., WORTMAN, G. E;r AL. 1998. Yadong cross structure and South Tibetan detachment

199

in the east central Himalaya (89 degrees - 90 degrees E). Tectonics, 17, 28-45. XYPOLIAS, P. & KOUKOUVELAS,I. K. 2001. Kinematic vorticity and strain rate patterns associated with ductile extrusion in the Chelmos Shear Zone (External Hellenides, Greece). Tectonophysics, 338, 59-77. ZHAO, W., NELSON, K. D., CHE, J. EZrAt.. 1993. Deep seismic reflection evidence for continental underthrusting beneath southern Tibet. Nature, 366, 557-559.

Ductile extrusion in continental collision zones: ambiguities in the definition of channel flow and its identification in ancient orogens R. R. J O N E S 1'2, R. E. H O L D S W O R T H 3, M. H A N D 4 & B. G O S C O M B E 4'5

1Geospatial Research Ltd, Department of Earth Sciences, University of Durham, DH1 3LE, UK (e-mail: [email protected]) 2e-Science Research Institute, University of Durham, DH1 3LE, UK 3Reactivation Research Group, Department of Earth Sciences, University of Durham, DH1 3LE, UK 4Continental Evolution Research Group, Geology and Geophysics, University of Adelaide, 5005, Australia 5Northern Territory Geological Survey, PO Box 8760, Alice Springs, 0871 NT, Australia Abstract: Field characteristics of crustal extrusion zones include: high-grade metamorphism flanked by lower-grade rocks; broadly coeval flanking shear zones with opposing senses of shear; early ductile fabrics successively overprinted by semi-brittle and brittle structures; and localization of strain to give a more extensive deformation history within the extrusion zone relative to the flanking regions. Crustal extrusion, involving a combination of pure and simple shear, is a regular consequence of bulk orogenic thickening and contraction during continental collision. Extrusion can occur in response to different tectonic settings, and need not necessarily imply a driving force linked to mid-crustal channel flow. In most situations, field criteria alone are unlikely to be sufficient to determine the driving causes of extrusion. This is illustrated with examples from the Nanga Parbat-Haramosh Massif in the Pakistan Himalaya, and the Wing Pond Shear Zone in Newfoundland.

In this paper we review channel flow and crustal extrusion in relation to field geology. Our aim is to provide a critical evaluation of whether the predictions of channel flow models are testable using field observations, and also to consider the relevance of channel flow in relation to existing models of crustal extrusion. We illustrate the inherent difficulty and ambiguity of applying the channel flow concept with field examples from both active and ancient orogens (the Nanga Parbat-Haramosh Massif, Pakistan Himalaya, and the Wing Pond Shear Zone in the Appalachian orogen, Newfoundland). General aspects of orogenic thickening and crustal extrusion are presented first, followed by more specific discussion of channel flow and the nature and limitations of possible diagnostic characteristics.

Orogenic thickening and crustal extrusion Field studies in orogenic belts have long shown that deformation arising from continental collision

generally involves a combination of large-scale overthrusting and bulk crustal thickening. This combination is sometimes considered in terms of the relative importance of simple shear and pure shear to the mountain-building process, although these are end-member strain components and most parts of an orogen will experience more general shear (e.g. Sanderson 1982; Platt & Behrmann 1986; Passchier 1986, 1987; subsimple shear of Simpson & De Paor 1993). The variety and complexity of structures typically observed in orogenic belts illustrate that the distribution and localization of strain is highly heterogeneous (e.g. Jones et al. 2005). Overthrusting is associated with zones of high strain concentration and high vorticity that give the orogen marked polarity (i.e. asymmetry), recognizable in outcrop as dominantly forelandverging structures, including shear zones with topto-foreland shear sense. Regions of bulk crustal thickening represent a more widely distributed response to horizontal shortening across the orogenic belt, and show large variations in width and in localization of strain. The boundaries to these

From: LAW, R. D., SEARLE,M. P. & GODIN,L. (eds) ChannelFlow,DuctileExtrusionand Exhumationin Continental Collision Zones. Geological Society, London, Special Publications, 268, 201-219. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

202

R.R. JONES ET AL.

regions can be well-defined shear zones or diffuse transitions with gradually varying style and intensity of strain. Some areas of large-scale crustal thickening can involve broad zones characterized by intense folding and widespread formation of pervasive tectonic fabrics. Crustal shortening can also be localized into narrow regions of more intense deformation (often marked by well-defined high strain boundaries), in which shortening is matched by thickening occurring as concentrated zones of crustal extrusion. Irrespective of whether strain is widely distributed across a broad area, or localized into a more concentrated zone of extrusion, diagnostic characteristics of crustal extrusion include the uplift of high-metamorphic-grade rocks in the core of the zone (e.g. the Kaoko Belt in Namibia; Goscombe et al. 2003a, b, 2005), and the reversal in the sense of vorticity across the region of deformation (e.g. southwestern Hellenides; Xypolias & Doutsos 2000; Xypolias & Koukouvelas 2001; Xypolias & Kokkalas 2006). Numerical modelling of crustal extrusion

Regions of crustal extrusion bounded by coeval, opposing shear zones are a characteristic feature of numerical models of deformation that involve combinations of coaxial and non-coaxial strain (e.g. Sanderson 1982; Sanderson & Marchini 1984). Although strain matrix modelling is naively simplistic in approach, and explores the kinematics of deformation without any thermalmechanical considerations, it is computationally efficient, and has the advantage that it can be used to model first-order relationships between bulk finite strain and boundary conditions. In addition, it also helps to demonstrate the innate complexity of progressive deformation in four dimensions (4D) (i.e. one temporal and three spatial dimensions). Strain matrix modelling has been applied extensively to improve understanding of various non-coaxial, non-plane strains that typify zones of transpressional and transtensional deformation

(e.g. Fig. 1; Tikoff & Fossen 1993; Jones et al. 1997, 2004; Fossen & Tikoff 1998; De Paola et al. 2005). The significance of transpression is that deformation resulting from oblique plate collision cannot be accommodated by simple shear alone, and must involve a significant component of crustal extrusion, so numerical modelling of transpression can provide insight into the interplay between orogenic shortening, overthrusting and crustal thickening. Application of transpressional strain modelling to areas of oblique convergence and divergence has shown close correlation between model predictions and observed aspects of crustal extrusion (e.g. Robin & Cruden 1994; Teyssier et al. 1995; Jones & Tanner 1995; Holdsworth et al. 1998, and references therein). In the case of transpression (Fig. lc, d), upward extrusion of crustal material is caused by shortening across the deformation zone, and is thus equally applicable to plane-strain coaxial shortening during orthogonal collision (Fig. l a, b). The thermal implications of crustal thickening and extrusion have been explored for both orthogonal and oblique shortening by Thompson et al. (1997a, b). Thermal-mechanical models (Ellis et al. 1998), representing an extra level of geological realism compared with simplistic strain modelling, reproduce similar combinations of large-scale overthrusting and broad zones of bulk crustal extrusion resulting from continental collision. In general, strain modelling usually assumes that crustal extrusion occurs by upward migration, deformation and erosion of the Earth's surface, although it is also possible for crustal material to be extruded laterally (i.e. axially, parallel to the orogenic trend; Dias & Ribeiro 1994; Jones et al. 1997; Dewey et al. 1998). Note that extrusion in this sense refers to the lateral escape of crustal blocks (e.g. Eastern Anatolia; Dewey et al. 1986). Lateral extrusion is also used in a slightly different sense to describe outward ductile flow at lower crustal levels, driven by lateral pressure gradients (e.g. Bird 1991; Clark & Royden 2000).

Fig. 1. First-order kinematics of tectonically driven crustal extrusion caused by boundary displacement: (a) vertical extrusion, plane strain; (b) inclined extrusion, plane strain; (c) vertical extrusion, non-coaxial non-plane strain (transpression); (d) inclined extrusion, non-coaxial non-plane strain (inclined transpression).

DUCTILE EXTRUSION AND CHANNEL FLOW E v i d e n c e f o r c r u s t a l e x t r u s i o n in c e n t r a l and eastern Himalayas

Recognition of north-dipping normal faulting over large distances along-strike in Tibet (Burg et al. 1984), led Burchfiel & Royden (1985) to suggest that the Greater Himalayan Sequence (GHS) has been extruded relative to Tibet and the Lesser Himalaya. Subsequent work has recently been extensively discussed and summarized by Searle et al. (2003), so only an overview of field-based evidence for crustal extrusion is given here. The possible relationship between crustal extrusion of the GHS and mid-crustal channel flow (together with evidence from regional geophysics) is discussed later. The GHS is basally bounded by the Main Central Thrust (MCT), and separated from the overlying Tibetan crust by the South Tibetan Detachment (STD) system (Gansser 1964; Burg et al. 1984). Structural descriptions of normal displacement on the STD are given by Burchfiel et al. (1992), Edwards et al. (1996), Grujic et al. (1996, 2002) and Law et al. (2004). Evidence for contemporaneity of movement on the MCT and STD between c. 22 and 17 Ma is documented by Hodges et al. (1992, 1993) and Walker et al. (1999). Detailed kinematic analyses from the MCT (Grasemann et al. 1999) and STD (Law et al. 2004; see also Carosi et al. 2006; Jessup et al. 2006) suggest that

203

finite strain in the GHS is a subsimple shear (i.e. a combination of pure shear flattening and simple shear). Higher metamorphic grades are recorded from central parts of the GHS, with inverted pressure-temperature profiles across the MCT (e.g. Hodges et al. 1988, 1993; Searle & Rex 1989; Vannay & Grasemann 2001). Crustal-melt Miocene leucogranites were emplaced between the MCT and STD (e.g. Searle et al. 1993; Searle 1999; Edwards et al. 1996; Wu et al. 1998). Notwithstanding ongoing debate regarding some of the field evidence (e.g. Hubbard & Harrison 1989; Harrison et al. 1999), the combined stratigraphical, structural, metamorphic and geochronological data provide a strong indication that the Higher Himalayan slab has undergone southward-directed extrusion. The existence of this body of fieldbased evidence is independent of geodynamic interpretations that are based on regional geophysics, or that link crustal extrusion to mid-crustal channel flow beneath Tibet. Summary: crustal extrusion

In this paper we use crustal extrusion as a general term to signify the commonly recognized thickening/ extrusion of crust in response to bulk contraction. In this sense, the term implies no specific geodynamic process or driving force. The bulk geometry and kinematics of crustal extrusion can be modelled in

Fig. 2. Crustal extrusion in relation to channel flow: (a) schematic (scale-independent)depiction of typical criteria that characterize regions of crustal extrusion: such criteria are consistent with crustal-scale channel flow, but may also be the result of other driving mechanisms; (b) variation in the degree of strain localization at the margins of crustal extrusion zones: upper figures are schematic cross-sections; lower graphs show generalized variation in strain magnitude for a typical transect across the zone; (c) schematic depiction of causative processes and metamorphic and structural characteristics of channel flow predicted from thermal-mechanical modelling of Beaumont et al. (2001, 2004) and Jarn]~0n et al. (2002, 2004).

204

R.R. JONES ET AL.

4D using strain matrix modelling. Diagnostic field characteristics of crustal extrusion include: metamorphic inversion at the lower boundary of a highgrade zone flanked by lower-grade rocks; highgrade regional metamorphism, possibly with partial melting in the core of the zone; broadly coeval shear zone margins with opposing senses of shear; early ductile fabrics progressively overprinted by later brittle structures; and localization of strain, with some phases of deformation only seen in the zone and not in the flanking rocks (Fig. 2a).

Channel flow D u c t i l e f l o w in the l o w e r crust

Laboratory studies of rock strength (e.g. Goetze & Evans 1979; Brace & Kohlstedt 1980), as well as deep seismic data (e.g. Chen & Molnar 1983; Kuznir & Matthews 1988), have underpinned suggestions that layers of low viscosity in the lower crest can undergo ductile flow (e.g. Kuznir & Matthews 1988; Bird 1989, 1991; Block & Royden 1990; Molnar 1992; Bott 1999; see also Maggi et al. 2000, and Jackson 2002). These ductile zones may be weak enough to allow a degree of mechanical decoupling between upper and lower crust (e.g. Oldow et al. 1990; Royden 1996; Ellis et al. 1998; Teyssier et al. 2002; Tikoff et al. 2002; Grocott et al. 2004; Klepeis et al. 2004). Possible representatives of lower-crustal flow in the rock record include large tracts of granulite and migmatite terranes, and regions of large-scale ductile sheath folding. Numerical modelling of ductile flow in the lower crust has shown that the flow may be driven by gravitational potential in regions of elevated topography, and could provide an effective mechanism for the compensation of local gravity anomalies (Bott 1999), and for large-scale variations in Moho topography to be gradually smoothed over geological time (Kuznir & Matthews 1988; Bird 1991). This approach to modelling has utilized standard fluid dynamics equations (e.g. Turcotte & Schubert 1982, 2002), generally assuming laminar flow within a channel geometry. Similar flow conditions have also been used to model exhumation of material along subduction channels by reverse flow (England & Holland 1979; Mancktelow 1995). Common to all the above models is that the location of channel flow is coincident with lowstrength regions of the lower crust above the assumed rheology transition at the crust-mantle interface.

applicable to mid- and upper-crustal levels, and used channel flow terminology in a conceptual model to explain field observations from Bhutan. This extended ideas developed by Nelson et al. (1995, 1996; see also Hauck et al. 1998, and Alsdorf et al. 1998), who presented geophysical data from the INDEPTH project, and postulated that mid-crustal flow is occurring under the Tibetan Plateau, with the southerly continuation of this mid-crustal low-viscosity layer being extruded as the GHS. An assumption of this model is that partial melting related to crustal thickening can radically alter the standard strength profile of the crust, allowing weaker zones of reduced viscosity to develop at mid-crustal levels that would generally be too strong for extrusive flow to occur. This concept has influenced thermal-mechanical models of the Himalaya, which have incorporated low-strength mid-crustal theologies in order to induce channel flow to develop. Thermal-mechanical

models

and channel flow

Recent advances in thermal-mechanical finite element modelling (e.g. Ellis et al. 1998; Jamieson et al. 2002) represent a new level of sophistication in the analysis of progressive evolution of orogenic belts. Thermal-mechanical models allow testable predictions to be made in relation to geodynamics, tectonics, geomorphology, and gross structural and metamorphic patterns. Application of this approach to lithospheric-scale modelling has been applied to the Himalaya-Tibet orogen (e.g. Beaumont et al. 2001, 2004, 2006; Jamieson et al. 2004, 2006). In these cases, the boundary conditions and input parameters are carefully chosen to be geologically realistic and to match actual Himalayan values as closely as possible. Conformity between the results of modelling and many first-order geological and geophysical characteristics of the Himalaya-Tibet belt suggests that the models are well calibrated to this specific orogen. The importance of this is that proper calibration is essential prior to using the modelling to test the sensitivity of different input parameters in relation to resultant tectonism. In other words, it should not be considered a weakness of the modelling method that a model is tuned to match a specific orogenic setting. On the contrary, this is the basis for a better understanding of which factors are likely to have most influence on dynamic orogenic processes such as orogenic thickening, crustal extrusion and channel flow. Implications from thermal-mechanical

C h a n n e l f l o w in the mid- a n d u p p e r crust

models

Grujic et al. (1996) subsequently suggested that comparable fluid dynamics principles might be

Iterations of the thermal-mechanical modelling that test the sensitivity of various input parameters

DUCTILE EXTRUSION AND CHANNEL FLOW (Beaumont et al. 2004; Jamieson et al. 2004) emphasize that channel flow is unlikely to have developed unless certain boundary conditions were favourable. The aim of the following characterization of channel flow is to separate key attributes of the models into two main types, broadly corresponding to 'cause and effect': firstly, transient features relating to prerequisite tectonic and geodynamic boundary conditions shown to be important in inducing channel flow ('cause'); and secondly, attributes relating to the resultant channel ('effect'), evidence for which may remain in the rock record over significant periods of geological time, long after orogenesis has abated (Fig. 2c).

Modelling: requisite conditions for channel flow to occur Large, hot, long-lived orogen.

Modelling suggests that in order for a mid-crustal low-viscosity channel to develop, the crust must be suitably radiogenic, and large-scale crustal thickening must last long enough for partial melting to occur. An ancient orogen that was small, short-lived or had radiogenically depleted crust would be unlikely to have been able to support significant gravity-driven channel flow.

towards the Earth's surface. Thus, high denudation rates are not a prerequisite for the channel to develop at depth (e.g. fig. 6a of Beaumont et al. 2004), but may be instrumental in promoting the exhumation of an active channel. In most ancient orogens, direct evidence of elevated rates of denudation will have disappeared, although indirect evidence may be present (e.g. from pressuretemperature-time (P-T-t) data, or from the sedimentary record in adjacent deposits; e.g. Galy et al. 1996).

Field characteristics of resultant channel flow inferred from modelling Pressure, temperature and relative displacement paths in thermal-mechanical models allow inferences to be made about the resultant metamorphic and structural characteristics that should be visible in outcrop if channel flow had developed. Here we evaluate the degree to which specific characteristics will be useful as diagnostic attributes in ancient orogens. We consider features of the midcrustal channel, as well as upper-crustal extrusion zones where the channel reaches the Earth's surface.

Subhorizontal mid-crustal channel.

Extensive elevated plateau region.

One of the main conclusions of modelling is that mid-crustal channel flow is driven by a large horizontal gradient in lithostatic pressure relating to a wide upland plateau region (corresponding to Tibet in the application of the model to the Himalayas), flanked by an erosional/topographic front. Indeed, the thermalmechanical models for the Himalaya-Tibet orogen show that there appears to be such a strong dependency on gravitational loading that channel flow is unable to develop unless a large plateau is present. In an ancient orogen, lack of evidence of a sufficiently large source of gravitational potential (such as a plateau) will therefore make it significantly more difficult to demonstrate conclusively that gravity-driven channel flow has occurred. In many old orogens it may no longer be possible to demonstrate the presence of an ancient plateau region, particularly in orogenic belts that have been heavily reworked (e.g. Fenno-Scandian, Grenvillian, Grampian, Acadian, Variscan and many others), or those in which large lateral motions along strike-slip faults have dismembered much of the original collisional architecture (e.g. Western Cordillera, Scandian, Taconic, etc.).

Evidence of high denudation rates.

Modelling has shown that a high rate of denudation at the margins of the elevated topographic front can have a major effect in causing the channel to propagate ('tunnel') upwards from mid-crustal levels,

205

Post-orogenic denudation and exhumation of mid-crustal rocks may expose evidence of the fossilized channel. This should contain regions of partial melting, migmatization and elevated metamorphic grade, flanked above and below by rocks that have experienced lower temperature metamorphism. Therefore, for positive recognition of possible channel flow, a full section showing both upper and lower channel margins should ideally be exposed. Modelling implies that structural evidence should be useful in recognizing whether channel flow may have occurred. One of the most important diagnostic characteristics is to demonstrate that shear zones were active simultaneously at both top and base of the channel, and for the zones to have demonstrably opposing shear sense ('toptowards-orogenic-front' for the basal shear zone, 'top-towards-orogenic-core' for the upper shear zone). Furthermore, the age of shearing should be broadly similar to the timing of migmatization in the channel. Difficulties may arise if the channel margins have been reworked or reactivated by subsequent tectonic episodes, as this may remove much of the evidence for earlier shear during the original channel flow. Model predictions suggest that recognition of high strains might provide useful diagnostic information. In an idealized channel, intensity of shear strain should be greatest at the margins of the channel, and lower in the channel centre, though in reality strain is likely to be more heterogeneously

206

R.R. JONES ET AL.

distributed. Consideration of particle path trajectories in thermal-mechanical models (e.g. figs 4 & 6 of Jamieson et al. 2004) show that for welldeveloped channel flow, shear strains (3/) of 102 or more may be common along the channel margins, with y in excess of 103 also possible. Fault rocks that have experienced such high levels of strain are generally recognizable by their ultra-mylonitic fabric, even if the exact magnitude of shear strain is difficult to quantify precisely. Within the channel there should be evidence of subsimple shear strains (i.e. a non-coaxial shear with shortening across the deformation zone; examples are given by Grasemann et al. 1999; Xypolias & Doutsos 2000; Vannay & Grasemann 2001; Xypolias & Koukouvelas 2001; Law et al. 2004). In such cases, the strain magnitude in the channel margins should be lower towards the core of the orogen where the channel starts, with an increase in strain outwards towards the orogenic front (cf. 'cream-cake tectonics' of Ramsay & Huber 1987, pp. 610-613). Clearly this diagnostic attribute is only testable if a large extent of the mid-crustal channel is exposed across the orogen. Strongly localized deformation due to high shear strain within the channel is likely to give rise to narrow terranes (domains) characterized by complex structural relationships, in which there is repeated overprinting of successive deformational fabrics. This may give the appearance of several phases of deformation (fabric crenulation, refolded folds, sheath fold development, fold and fabric transposition), though these should span a relatively short time period. Because these deformational 'phases' are caused by spatial and temporal heterogeneity along the channel, and relate to localized perturbations within the flow, the rocks outside the channel should be markedly less deformed, and some (if not most) of the 'phases' should be restricted to the channel, and will be largely absent beyond the channel margins. Zone crust.

of

crustal

extrusion

in

the

lower-grade flanks, with the high P-T migmatitegrade channel enveloped by markedly lower pressure and temperature greenschist- or amphibolite-grade rocks typical of upper-crustal orogenic levels. P-T-t conditions should indicate that rapid exhumation of the channel occurred synchronously with the timing of shear along the channel margins. Coeval shear zones at the top and base of the channel should once again show opposing shear sense, with thrust displacement along the lower channel margin and normal-sense displacement at the upper margin. Modelling emphasizes that in order for the channel to have propagated towards the surface, both margins must have experienced very large strain magnitudes (material in the core of the channel must be transported large distances during displacement from mid- to upper crust). For flow in the channel to remain ductile at high crustal levels, high strain rates are needed, so that there is insufficient time for the channel to cool to temperatures at which increase in viscosity would impede flow. Thus in ancient orogens, typical characteristics of upper-crustal regions of the channel should include the presence of high strain mylonites or ultramylonites that are progressively overprinted by successive down-temperature and pressure phases of semi-ductile, semi-brittle and then brittle deformation, as material in the channel is exhumed and extruded towards the Earth's surface. Younger brittle structures should generally show the same sense of shear as the older ductile fabrics that they overprint. In summary, thermal-mechanical modelling allows specific predictions to be made about the field characteristics that can be expected in zones of channel flow. However, many of these characteristics are non-unique to channel flow: they are common to general zones of bulk crustal extrusion, and are widely recognized in areas in which channel flow is not believed to have occurred. The significance of this in relation to active and ancient orogens is discussed in the following sections.

upper

Modelling implies that the exposure of a fossilized channel that has 'tunnelled' from midcrustal to upper-crustal levels at the orogenic front, will share some diagnostic attributes with a corresponding channel at mid-crustal levels. As at greater crustal depths, the migmatitic channel should be flanked above and below by rocks of lower metamorphic grade, to give characteristic inverted and right-way-up metamorphic isograds at the base and top of the channel respectively (see section 10.1 of Jamieson et al. (2004) for discussion). This metamorphic inversion is a particularly important diagnostic feature in the upper crust, as there is likely to be greater contrast in metamorphic grade between the channel and its

Discussion Application of the channel flow concept to explain Himalayan-Tibetan orogenesis has clearly been a positive catalyst in promoting vigorous scientific discussion. Much of the current debate focuses on the reliability and validity of geophysical data, and its relevance in relation to field observations of Himalayan geology. In the following discussion we focus on other inherent difficulties of the channel flow concept, particularly regarding recognition of channel flow in ancient orogens, and whether or not the concept is useful to field-based analysis of continental collision zones.

DUCTILE EXTRUSION AND CHANNEL FLOW

Ambiguity of channel flow semantics Earlier sections of this paper illustrate that common usage of the term 'channel flow' can convey a range of slightly different meanings, each of which can carry subtle implications depending upon context. For example, with regard to fluid dynamics, channel flow is simply a generic term that encompasses a number of more specific types of quantitative flow equation. In contrast, when used with reference to outcrop observations (e.g. Grujic et al. 1996), channel flow is a conceptual term that is consistent with crustal extrusion (implying inverted metamorphism, high shear strains, and contemporaneous thrusting and normal faulting). In this context, for some workers the term 'channel flow' may also tacitly imply that continuation of the channel extends from the surface down to mid-crustal levels. Finally, within the context of the thermal-mechanical models of Beaumont et al. (2001, 2004), channel flow is a dynamic, gravity-driven process that requires high heat flow and large gravitational potential for significant amounts of flow to occur, as well as high rates of localized exhumation for the channel to reach the surface. In relation to Himalayan geology, these semantic discrepancies are generally unimportant, since most proponents of channel flow generally include discussion that encompasses all the different contexts outlined above. However, in older orogenic belts, the distinction may be much more important, since in general there will be a very different balance of available data upon which interpretation can be based.

Nature of evidence in ancient orogens The study of modern mountain belts is essential in order to improve our understanding of processes that produced ancient orogens. Conversely, ancient orogens can also increase our understanding of modern-day orogenic processes, because exhumation of ancient mountains allows us to study large tracts of deep levels of the orogen that are not widely exposed in active mountain regions today. In modern orogens such as the HimalayanTibetan system, many different types of data are available to help constrain tectonic interpretation and to guide the choice of suitable boundary conditions and geodynamic input parameters to use in modelling. Some of the most useful types of data are transient in nature (e.g. earthquake focal mechanisms, heat flow, bright spots on deep seismic reflection data, geodetic measurements showing relative displacement recorded by arrays of GPS stations). Other criteria can be longer-lived (e.g. topography, and sea-floor magnetic anomalies that

207

record relative plate motion). However, in the majority of older orogens, these types of transient short-lived geodynamic data provide little or no constraint. Most ancient terranes do not have well-preserved boundary conditions, and it can be extremely difficult to infer the nature of the farfield driving forces that controlled deformation (e.g. Dewey et al. 1986). Furthermore, factors suggested by thermal-mechanical modelling to be prerequisite for channel flow to occur - such as high gravitational potential, heat flow and exhumation rates - are also transient in nature, and will generally be of much less use in constraining geodynamic interpretations of ancient orogens. Thus, many of the types of data that are currently driving debate about channel flow in the Himalayas will not be as useful when testing the applicability of the channel flow concept to ancient orogens. In such cases, interpretation will inevitably place greater emphasis on field studies and analysis of samples collected at outcrop, and must rely heavily on a combination of critical types of metamorphic, structural and geochronological evidence. In our opinion, in order to test different geodynamic models as objectively as possible, it is important to maintain a descriptive framework for making field observations that avoids recording data in terms of a genetic description linked to a specific interpretation of causative processes.

Non-uniqueness of field characteristics inferred from modelling Predicted field characteristics of channel flow inferred from thermal-mechanical modelling show close similarity with actual field observations from the Himalayan-Tibetan system (see above); However, while field data do provide strong evidence that crustal extrusion has taken place in the High Himalaya, it is much more contentious to demonstrate incontrovertibly that extrusion at the surface is linked to zones of channel flow beneath Tibet far to the north. Even in modern orogens where transient geodetic and geophysical data can be acquired, a fundamental challenge facing geoscientists is to test the range of possible geodynamic interpretations that are consistent with the available data. At the current level of understanding, the characteristics predicted by thermal-mechanical modelling, as well as actual field observations and analyses from the Himalayas, are non-unique in terms of causative geodynamic driving forces. In particular, regions of crustal extrusion can result from horizontal forces relating to convergent plate motion, and need not imply processes driven predominantly by gravitational potential. This is illustrated below with an example from the Nanga

208

R.R. JONES E T A L .

Parbat-Haramosh Massif of northern Pakistan. In older orogenic systems these challenges are exacerbated, as shown later with an example from the Wing Pond Shear Zone in the Appalachian orogen of Newfoundland.

Crustal extrusion in the Nanga Parbat-Haramosh Massif The Nanga Parbat-Haramosh Massif (NPHM) lies at the NW end of the Himalayan arc, and is bordered to the west by the Hindu Kush, to the north by the Karakorum range, and to the SE by the High Himalayas (Fig. 3a). As part of the long-lived, hot Himalayan system, with the vast Tibetan plateau to the north, the massif shares many of the geodynamic attributes that characterize regions further to the east. The NPHM is marked by extreme topography, with over 7000 m of relief between the summit of Nanga Parbat (8125 m) and the base of the Indus valley. Rapid denudation is facilitated by the enormous sediment-bearing capacity of the Indus, which may have allowed an effective positive feedback mechanism to develop, where increased uplift rates are driven by rapid erosion (Zeitler e t al. 2001 ; Koons e t al. 2002).

The massif is a large asymmetric antiformal dome (Fig. 3b), forming a crustal-scale syntaxis (Fig. 3c) that marks a major change in the overall trend of the orogenic belt (Tahirkheli & Jan 1979; Coward 1983, 1985). Formation of the syntaxis has reworked the earlier structure of the Main Mantle Thrust (Butler & Prior 1988a, b), which represents the tectonic plate boundary between India and Asia. The age of the syntaxis post-dates peak metamorphism, and probably initiated within the last 10 million years (Treloar e t al. 2000; Zeitler et al. 2001; Butler et al. 2002). The core of the massif consists of migmatitic crystalline basement rocks of Indian plate affinity that have been extruded through structurally overlying units of the Kohistan-Ladakh Arc, which had accreted to the SE margin of the Asian plate prior to collision with India from c. 55 Ma onwards (Treloar et al. 1989; Treloar & Coward 1991). Active crustal extrusion is ongoing, and is associated with extremely rapid uplift (Zeitler 1985) and high denudation rates (Bishop & Shroder 2000), suggesting exhumation rates of up to 7 mm a ~ (Zeitler 1985; Zeitler e t al. 2001). Seismic data suggest that the base of the upper-crustal seismogenic zone is significantly elevated beneath the syntaxis (Meltzer & Christensen 2001; Meltzer et al. 2001). The NPHM is one of the most intensely studied regions of Himalayan geology, and there

Fig. 3. Nanga Parbat-Haramosh Massif (NPHM), northern Pakistan: (a) satellite image of Himalaya, Tibet and the Indian subcontinent (courtesy of NASA/Visible Earth); (b) cross-section showing asymmetric antiformal nature of the massif and amount of material removed by erosion (after Butler et al. 2002); (c) map of the massif (collated from Butler et al. 1989, 2000; Searle & Khan 1996; PEcher & Le Fort 1999; Argles 2000; Edwards et al. 2000).

DUCTILE EXTRUSION AND CHANNEL FLOW is a wealth of published data that integrate a wide range of stratigraphic, metamorphic, structural, isotopic, geochemical and geophysical evidence. Detailed descriptions of the geology of the massif are given by Khan et al. (2000) and references therein, as well as Zeitler et al. (2001) and Butler et al. (2002). Geochronological, metamorphic and structural data provide strong evidence that ductile extrusion of mid-crustal material is taking place in the core of the massif. Fission track data and muscovite, biotite and hornblende cooling ages show that the central parts of the NPHM have experienced very rapid exhumation relative to surrounding regions: 4~ mica cooling ages suggest rates of between 3 - 4 mm a -~ (Whittington 1996) and 7 mm a -1 (Zeitler 1985; Zeitler et al. 2001). A range of P / T isotope data suggest that up to 25 km of crust may have been eroded from the central parts of the massif within the last 10 million years (Butler et al. 1997; Whittington et al. 1999). Uplift and exhumation are associated with partial melting and emplacement of small granitoid bodies. These yield zircon ages as recent as 1 Ma (Zeitler et al. 1993), and generally young towards the core of the massif (summarized in Zeitler et al. 2001). Structural studies around the margins of the massif have revealed a protracted kinematic history in which earlier ductile fabrics are generally overprinted by later semi-ductile to brittle structures. The age of deformation is locally well constrained, either by direct dating of structural fabrics, or by dating of cordierite seams and anatectic leucogranites that are intimately associated with shearing (e.g. Zeitler & Chamberlain 1991; Zeitler et al. 1993; Butler et al. 1997; Schneider et al. 1999; Whittington et al. 1999). Currently active thrusting along the western margin of the NPHM is emplacing Nanga Parbat gneisses northwestwards over unconsolidated fiver gravels and rocks of the Kohistan arc (Butler & Prior 1988a; Butler 2000). In their hanging wall, these brittle structures carry semi-brittle and ductile fabrics that also have the same top-to-NW kinematics, including 1 - 2 km of mylonitic gneiss with welldeveloped S-C fabrics (Butler 2000). Detailed fabric analysis comparing variation in the ellipticity and preferred orientation of augen, shows that the finite strain in the gneiss is a subsimple shear, with a strong component of up-dip elongation (Butler et al. 2002). Boudinaged calc-silicate layers suggest elongations of 30%. There is also evidence of a component of right-lateral strikeslip along the western margin (Butler et al. 1989; Butler 2000), showing that the bulk deformation of the region is transpressional. The sequence of ductile to brittle structures now visible at the

209

surface is likely to be very representative of comparable structures that are currently forming at different depths in the massif, and which will be progressively exhumed as crustal extrusion continues. The eastern margin of the NPHM is not as widely studied or well understood as the western margin. There are local areas of top-to-NW thrusting (Argles 2000), while south of Nanga Parbat, topto-south thrusting has been documented (Butler et al. 2000). However, most parts of the eastern margin appear to be marked by ductile fabrics of the Main Mantle Thrust (MMT), subsequently steepened during fonrlation of the syntaxis (Butler et al. 1992; Wheeler et al. 1995), without significant overprinting of recent brittle or semi-brittle deformation (Argles 2000). Although further fieldwork is clearly needed to provide additional constraints on the kinematics of the eastern margin, available evidence suggests that later (post-MMT) strain is more distributed than in the west, and that topto-SE deformation was accommodated largely by upwarping of the eastern antiformal limb during formation of the syntaxis. Summary

There is a wealth of field data showing key characteristics of crustal extrusion in the NPHM, including: high-grade units in the core of the massif, with inverted metamorphism across the lower margin; evidence of partial melting, marked by younger ages in the centre of the zone; active deformation synchronous with crystallization of anatectic melts; ductile fabrics progressively overprinted by semi-ductile and brittle structures; and subsimple shear strains, with a large component of up-dip stretching. Recognizing that the prerequisite geodynamic conditions for channel flow are present (large, hot orogen, elevated plateau, extreme denudation), Beaumont et al. (2004) proposed that the NPHM is another potential example of active channel flow. However, it remains extremely difficult to establish the main causative driving force(s) of extrusion in the massif, and several alternative models for its development have been proposed. These include the NPHM as a fault-related antiform at the lateral tip to major Himalayan thrusts (Coward et al. 1988); a zone of NW shortening in response to orogen-parallel extension around the Himalayan arc (Seeber & P~cher 1998); upwelling of crustal gneiss domes (P~cher & Le Fort 1999; Koons et al. 2002); crustal-scale buckling (Burg & Podladchikov 2000); and north-south constriction/transpression above a thrust tip (Butler et al. 2000). Although the far-field boundary displacements between India and Asia are well established and reasonably straightforward (Treloar & Coward 1991), the

210

R.R. JONES E T A L .

regional-scale tectonics of the syntaxis are highly complicated. Deformation in the massif is nonplane strain, is spatially very heterogeneous, and has changed significantly with time, even over short periods, so understanding the geodynamics of the region requires a 4D analysis of the system. Although vertical forces associated with the gravitational loading by Tibet may conceivably have direct influence on crustal melting and extrusion beneath the NPHM, it remains at least as plausible that the main driving force of extrusion is the far-feld tectonic force of the northward motion of India, and the resultant complex interactions that are occurring at the corner of the Indian indentor between the Himalaya, Karakorum and Hindu Kush.

Crustal extrusion in the Wing Pond Shear Zone, Newfoundland The Gander Zone in the Appalachian orogenic belt of northeastern Newfoundland (Fig. 4a) represents part of a Gondwanan-derived continental fragment thought to have been accreted to Laurentia during the closure of the Lower Palaeozoic Iapetus Ocean (e.g. Williams et al. 1988; Soper et al. 1992). The Iapetus suture in Newfoundland lies within the Dunnage Zone, the eastern part of which was obducted onto the Gander Zone during the Arenig with the boundary marked by an allochthonous unit of melanges (Gander River Complex, Fig. 4a) emplaced initially on eastward-directed thrusts. The present-day eastern margin of the NE Gander Zone is the Dover Fault Zone, a major reactivated brittle-ductile structure (Blackwood & Kennedy 1975; Holdsworth 1994). The Avalon Zone to the east comprises a characteristic Gondwanan Neoproterozoic tectonostratigraphy and Palaeozoic cover (O'Brien et al. 1983). The metasedimentary rocks forming most of the Gander Zone in NE Newfoundland (the 'Gander Group') are a thick sequence of variably deformed and metamorphosed clastic metasediments and gneisses (psammites, pelites) intruded by numerous granites (Fig. 4a; Blackwood 1977; Hanmer 1981; O'Neill 1991; Holdsworth 1994). The deformational and metamorphic character of the Gander Zone is highly heterogeneous. Two regional 'flat belts' are recognized, characterized mainly by greenschist-facies (chlorite-white mica) assemblages which are deformed by generally eastwardverging recumbent folds locally termed 'F2' (e.g. Fig. 5a; Kennedy & McGonigal 1972) thought to have formed during overthrusting of the Dunnage Zone in the Arenig (c. 480 Ma). These are post-dated by two 10-25km wide, high-strain transpressional steep belts (Fig. 4a) of similar

age, both of which are characterized by significantly elevated P-T conditions (D'Lemos et al. 1997; King 1997). The easternmost of these, here termed the Hare Bay Gneiss Shear Zone (HBGSZ), is characterized by rnigmatization and syntectonic granite emplacement during sinistral shear thought to be related to initial docking of the Gander and Avalon terranes during the Silurian (e.g. Holdsworth 1994). Peak metamorphism, derived from mineral reactions, cordierite geobarometry and hornblende-plagioclase geothermometry in the HBGSZ was at high temperature-low pressure conditions (c. 740~ at 4.5 kbar; D'Lemos et al. 1997). Further to the west, the Wing Pond Shear Zone (WPSZ) (O'Neill 1991) is a structurally more complex region of focused metamorphism and high strain up to 10 km across (Fig. 4b); its age relative to the HBGSZ is uncertain, although existing geological constraints suggest that they are broadly contemporaneous. The WPSZ is well exposed in roadside exposures along the Trans Canada Highway north and east of the later Gander Lake granite (Fig. 4b). Moving eastwards from Benton, early recumbent folds and associated greenschist-facies fabrics of the western Gander Zone are progressively overprinted and transposed by upright 'F3' folds and an associated steeply dipping, N E - S W trending '$3' fabric (Fig. 5b; O'Neill 1991). $3 is a prograde foliation defined by increasing proportions of biotite and is associated with low-P prograde metamorphism resulting in formation of cordierite- and andalusite-bearing assemblages as the intensity of $3 increases. In places andalusite and cordierite are overprinted by $3 (Fig. 6a, b), whereas in other instances porphyroblast growth occurred late- or post-S3 (Fig. 6c). For the general range of bulk compositions in the Gander Group, the growth of cordierite- and/or andalusite-bearing assemblages implies prograde metamorphism at _

~

~0

E=

"~.~_

'-."

L (l,1 .t.,.,

~0( HAl

~

7

I

;/, i

IIII

II II

~'~ o~-

~

I1[1111 II

III

II

" "~o\

L

L

ue,o,AoPJO

"--Z

LL

_ ~O'~" ~

.~.,

IIIIIIIII I[[LLLL]I[:IL[LI[LLLL[LIZELI

o

~

~

--03

-ui~unl,~

L L

LLILLL U [ [ L [ L [ [ [ I

III

IIIIIIIIIIIIIIIIIIIIIIII1[111111

I

~

~

[I

~

I[ i

~ I ~~I o

~ ~I

"~

~

- =

~

o

= =

=

o,,

eouenbeSue/~elew!i~i ~oludJoweleLuJeleeJID

.---..

"~I

-~

_o

,o.o~l

(::~

oo,~

N

-'- "-"

1%/

a., ~:

m~

""

~. o ~ro,_o,o,~

""

,

~r-9 1 o~ 7 O6) m

~

II ..o,o,o,o,o,~

/J

,s oS~N

\

dl9

I

~,~

m

~-, : o - - o

"~

~, 9

"~

u,,,,~o/

o

-

,0~ o8~N

O O eJ0O.l . u . u o. . o !~,>

--.I

.c:~ -~-

J

!

eouenbesfuelueuJlpes, ue/~qle.I,

._. _

o

.. .c:

"--I--

~

-"m.-.

"-- -':)

"-"

oS~N

S

olsse,Jj. -snoJe,,uoqJeo

O 0 O,~j~

=,J


-, r r 0 0 0 r

9

,.ii .,..~

0

Cq 0 0 "0

0 C',,I

04

0

r

,iq

,.-..,

Z; 9

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE

273

Fig. 3. (a) Two S W - N E geological cross-sections ( A - A ' and B - B ' ) , and one N W - S E cross-section (C-C') in thermochronology and U - P b the Nar valley, showing distribution of results obtained from the 4~ geochronology study. The section lines and legend are indicated on Figure 2. Modified from Searle & Godin (2003). (b) Schematic illustration of field area divided into three distinct structural levels: the Chame detachment separates the lowest structural level (Level 1) from Level 2; the Phu detachment separates Level 2 from the highest structural level (Level 3).

274

L. GODIN ET AL.

locally placing Tethyan sedimentary rocks in fault contact directly above the Manaslu leucogranite. The Phu detachment therefore must have been active after 19 Ma (Searle & Godin 2003).

Tethyan sedimentary sequence

As indicated on the map (Fig. 2) and cross-sections (Fig. 3a), the distribution of the Tethyan sedimentary sequence (Level 3) in the Nar valley is confined to two regions. The southern region encompasses exposures near Nar village and Pisang Peak, whereas the northern region includes exposures near Phu village and Pangre glacier. These two regions are separated by the Chako antiform, which exposes metamorphic rocks of the Greater Himalayan sequence. The Ordovician Nilgiri Formation, exposed in the Marsyandi valley near Britang and on the southern slopes of Pisang, is the lowest exposed stratigraphic level of the Tethyan sedimentary sequence. It consists of micritic grey massive to metre-thick beds of limestone, with rare poorly preserved brachiopods and crinoids, grading upwards into calcareous shale and siltstone (Bordet et al. 1975). The Nilgiri Formation is overlain by a SiluroDevonian Tentaculite-bearing black shale and siltstone unit (Bordet et al. 1975), exposed on the northern flank of Pisang, and also in the immediate hanging wall of the Phu detachment in the Phu Khola, south of Phu village (Fig. 2). Brachiopodrich Carboniferous-Permian calcareous shale and limestone of the Lake Tilicho and Thini Chu formations cap the Siluro-Devonian shale, and pass upwards into >200 m thick, black to grey Triassic limestone, shale and sandstone of the Thini Formation, visible in the vicinity of Nar and Phu villages (Fig. 2; Colchen et al. 1986). The uppermost stratigraphic unit exposed in the Nar valley is the Upper Triassic-Lower Jurassic Jomsom Formation (or Kioto Limestone; Garzanti 1999), consisting of a c. 500 m thick, grey to dun micritic limestone (Colchen et al. 1986).

Structural constraints Figure 3a presents two c. 35 km long N E - S W cross-sections of the Nar valley (A-A' and B-B'), from the northern slope of Annapurna IV to north of Phu village and the Manaslu leucogranite, respectively. Both sections are linked by a third orthogonal N W - S E cross-section (C-C'). The field area is divided into three distinct structural levels, separated by two segments of the STDS (Fig. 3b): the two lower levels (Levels 1 and 2) are within the Greater Himalayan sequence, whereas the upper level (Level 3) is confined to

the Tethyan sedimentary sequence. Rocks of Levels 1 and 2 are characterized by top-to-thesouth ductile flow structures, locally concentrated into several metre-thick high-strain zones, in contrast with the polyphase-folding style preserved in the overlying Tethyan sedimentary sequence of Level 3. The Chame detachment (Coleman 1996) juxtaposes rocks of similar metamorphic grade (Schneider & Masch 1993), and is interpreted as a ductile top-to-the-north shear zone wholly within the Greater Himalayan sequence, separating Levels 1 and 2 (Searle & Godin 2003). The Phu detachment is a ductile-brittle top-to-the-north normal fault that defines the uppermost, youngest strand of the STDS, separating Level 2 from Level 3. The Phu detachment is folded into a series of open structures (Mutsog synform, Chako antiform; Figs 3a & 4a). Above Nar village on the south limb of the Chako antiform, the fault juxtaposes unmetamorphosed Palaeozoic rocks of the Tethyan sedimentary sequence (Level 3) above biotite marble (Unit IV) and garnet schist (Unit V) of the Greater Himalayan sequence (Fig. 4b). Further north, on the north-dipping limb of the Chako antiform, the detachment juxtaposes unmetamorphosed upper Palaeozoic and lower Mesozoic rocks of the Tethyan sedimentary sequence (Level 3) in its hanging wall against the Manaslu leucogranite (Level 2) in its footwall (Figs 2, 3a & 4c; Searle & Godin 2003). Structural L e v e l 1

The structurally lowest level (Level 1) is composed of hornblende-biotite gneiss, leucogranitic augen gneisses and biotite schist, metamorphosed to amphibolite facies (up to 750~ Schneider & Masch 1993; Vannay & Hodges 1996). These units constitute the footwall of the Chame detachment. Rocks within structural Level 1 are characterized by strongly developed continuous metamorphic foliation and ductile transposition fabrics, locally affected by south-verging folds. Quartz ribbons, mineral aggregate lineations, and a higher concentration of top-to-the-south shear-sense indicators are observed within 1- to 100-m-thick high-strain zones (Fig. 5a, b, c). Top-to-the-south shear-sense indicators include sigma porphyroblasts, C-S fabrics, C' shear bands, and systematic north-dipping subgrain boundaries developed in quartz ribbons within Unit III granitic augen gneiss (Fig. 5b, c). Between high-strain zones, the rocks exhibit anastomosing fabrics, and lack mineral lineations and shear-sense indicators (Gleeson & Godin 2006). The upper boundary of Level 1 is defined by the top-to-thenorth Chame detachment (Coleman 1996; Searle & Godin 2003).

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE

275

Fig. 4. Field photographs of the Phu detachment and immediate footwall rocks. This fault crops out in the Marsyandi valley near Britang, dips west around Pisang peak, dips south over the village of Nar, wraps over the Chhubche massif, then dips to the NE as it wraps around the northwestern margin of the Manaslu leucogranite. (a) Panoramic photograph of Chhubche (5605 m) from Chako. The near-2000-m-high vertical cliff displays, from bottom to top, the hornblende-biotitegneiss (Unit II) with thin interlayer of granitic augen gneiss (Unit III), micaceous marble (Unit IV), and garnet-biotite schist (Unit V) of the Greater Himalayan sequence. Note the pervasive leucogranitic injections within Units II and III. The Phu detachment is projected just above the massif, and is folded by the Chako antiform, which plunges shallowly (7~ to the NW. (b) Phu detachment (Pangla Ri fault) above village of Nar, with Oligoceneage north-verging folds affecting the Tethyan stratigraphy in the immediate hanging wall. (e) NW margin of the Manaslu leucogranite, as seen from the Pangre glacier, east of Phu.

Structural Level 2 The intermediate structural level (Level 2) consists of upper-greenschist to lower-amphibolite facies garnet-bearing schist and gneiss, located in the hanging wall of the Chame detachment. Level 2 structures are characterized by a pervasive metamorphic foliation, overprinted by south-verging folds and related crenulation cleavage (Fig. 6a). Syntectonic garnets within Unit V indicate dominant southward rotation relative to cleavage during development of the metamorphic foliation (Gleeson & Godin 2006), consistent with moderately developed S-C fabrics and strain caps, which display evidence for general non-coaxial (mainly southward) ductile flow (Fig. 6b, c; cf. Passchier & Trouw 1996, p. 178). D2 deformation is characterized by south-verging, asymmetric folds, and development of north-dipping axial planar cleavage ($2) and hinge-parallel mineral lineations. The folds are open to closed, centimetre- to metrescale and overturned to the south with shallowly NW-plunging fold hinges. Level 2 folds exhibit

angular hinge zones and chevron fold shapes suggesting that they formed at higher structural levels than Level 1 folds of the same generation (Gleeson & Godin 2006).

Phu detachment In the Nar valley, Level 2 rocks are overlain by the Level 3 Tethyan sedimentary sequence along the Phu detachment (Searle & Godin 2003). The Phu detachment marks a clear metamorphic boundary between the Greater Himalayan sequence (sensu stricto) below, and the Tethyan sedimentary sequence above (Searle & Godin 2003). The immediate footwall rocks are dominated by intense flattening fabrics, mostly developed in the biotite marble Unit IV (Fig. 7a). In the northern end of the map area, in the Pangre glacier area, the northwestern part of the Manaslu leucogranite is cut by the Phu detachment (Searle & Godin 2003). In this area, the leucogranite contains a moderate solid-state foliation defined by muscovite

276

L. GODIN E T AL.

Fig. 5. Top-to-the-south ductile structures of Level 1. (a) Top-to-the-south hornblende-biotite mylonites within the Greater Himalayan sequence, Chako area. (b) Quartz proto-ribbon texture developed in the Unit III granitic augen gneiss (Sample N00-1 lb). (e) Detail of quartz subgrain boundaries (SB) in proto-ribbons. Their systematic north-dipping orientation is consistent with top-to-the-south ductile flow. and flattened dynamically recrystallized quartz (Fig. 7b). The immediate hanging wall of the Phu detachment is characterized by unmetamorphosed, yet highly deformed, Tethyan sedimentary rocks, which display numerous top-to-the-north brittle-ductile shear-sense indicators, such as sigma porphyroclasts and quarter structures

Fig. 6. Top-to-the-south ductile structures of Level 2. (a) South-verging folds and crenulations developed in the micaceous marble unit (Unit IV). Hammer is 40 cm long. (b) Moderately developed S-C fabric in Unit IV micaceous marble. Scale bar is 2 mm. (c) Opaque backrotated grain (replaced equinoderm?) in Unit V, with asymmetric strain caps and pressure shadows indicating top-to-the-south general non-coaxial ductile flow. Scale bar is 2 mm. (Fig. 7c; cf. Passchier & Trouw 1996). Crosssection constraints (see section A - A ' , Fig. 3a) suggest the Phu detachment may cut the northern end of the Chame detachment, just north of Kyang (Figs 2 & 3a).

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE

277

Structural Level 3

Fig. 7. Structures related to the Phu detachment. (a) Marble mylonite in the immediate footwall of the Phu detachment, between Kyang and Phu. (b) Photomicrograph of preferred alignment of muscovite and dynamically recrystaUized quartz in the Manaslu leucogranite, in the immediate footwall of the Phu detachment, upper Pangre glacier (Sample N02-12). Ms, Muscovite; P1, Plagioclase; Qz, Quartz. (c) Carbonate lithic porphyroclasts displaying sigma-type asymmetry and quarter structures developed in the shortening quadrants of the strain elipse (Passchier & Trouw 1996), compatible with top-to-the-north sense of shear in Triassic calcareous shale, in immediate hanging wall of Phu detachment, above Nar village (Sample N00-38).

The structurally highest rocks of the Nar valley are the Level 3 Palaeozoic to Mesozoic rocks of the Tethyan sedimentary sequence, which lie in tectonic contact with Level 2 rocks. The deformation and metamorphic style of the Tethyan sedimentary sequence is quite distinct from the Greater Himalayan sequence. It is characterized by polyphase folding, and a very low degree of metamorphism, similar to the Tethyan sedimentary sequence in the Kali Gandaki valley, c. 35 km west (e.g. Godin 2003). The Tethyan sedimentary sequence is affected by three fold generations of contrasting vergence. The second-generation structures (D2) are conspicuously defined by large-amplitude north-verging folds, interpreted to account for 150% vertical thickening of the Tethyan sedimentary sequence in Oligocene time, which are overprinted by south-verging kink folds and related crenulation cleavage (D4; Godin 2003). In the southern part of the map area, from Annapuma II to Chhubche, the Tethyan sedimentary sequence displays D2 north-verging fold structures (Fig. 3a). The north side of the AnnapurnaLamjung face is characterized by a large-amplitude north-verging anticline, mostly developed in lower Palaeozoic strata of the Tethyan sedimentary sequence (Fig. 3a). This fold, termed the Annapurna-Nilgiri fold by Border et al. (1975), is part of a larger fold system (e.g. Godin 2003) that is possibly associated with a top-to-the-north (thrust?) fault juxtaposing Ordovician and Silurian rocks (Nilgiri and Sombre formations) above Triassic and Jurassic strata in the valley between Pisang peak and Nar village (Figs 3a & 4b). This possible thrust fault is interpreted to be associated with D2 north-verging deformation, but alternatively could be an early top-to-the north STDSrelated normal fault, subsequently flexed into a south-dipping orientation. Both this fault and the D2 folds are cut by the Miocene Chame-Phu detachment system (D3; Fig. 3a; Searle & Godin 2003). North of Chhubche, on the northern limb of the Chako antiform, the structural style of the Tethyan sedimentary sequence is dominated by southverging to upright folds, and shallow north-dipping top-to-the-south brittle-ductile thrust faults (Fig. 8). These structures are interpreted to be related to the south-verging D 4 structures observed further west in the Kali Gandaki valley (Godin 2003).

Crustal-scale buckling A series of megascopic folds buckle the Greater Himalayan sequence, the overlying Tethyan

278

L. GODIN ET AL.

Fig. 8. North-dipping, top-to-the-south minor thrusts in Triassic limestones, near Phu village. These structures are concentrated on the northern limb of the Chako antiform. Note circled tent for scale. The cliff in the foreground is approximately 15 m high. sedimentary sequence, and all older structures including the Chame and Phu detachments (Fig. 3a; Searle & Godin 2003; Gleeson & Godin 2006). These buckles were previously described as the Mutsog synform and the Chako dome (Bordet et al. 1975; Coleman 1996), and are referred to here as the Mutsog synform and Chako antiform. The megascopic antiformsynform pair control outcrop patterns and foliation orientations in the Nar valley, and also account for the two sharp 90 ~ bends in the surface trace of the STDS (Fig. 1a, b). The non-cylindrical buckles are upright and open with fold axes plunging shallowly NW. The amplitude (c. 4 km) and wavelength (c. 25 km) of the Mutsog synformChako antiform implies crustal-scale folding. These buckles are overprinted by south-verging structures within the Tethyan sedimentary sequence (D4; Godin 2003).

D2

D3

Structural evolution

The following structural evolution sequence is proposed based on field and microstructural observations, supported by regional correlations and previous studies (e.g. Godin 2003; Searle & Godin 2003; Schill et al. 2003; Gleeson & Godin 2006). The deformation phases follow the nomenclature of Godin (2003). The first folding phase (Oa) in the Palaeozoic rocks of the Kali Gandaki valley has not been observed in the Nar valley.

D3_4

North-verging folds in the structural Level 3 Tethyan sedimentary sequence, and burial metamorphism of the Greater Himalayan sequence producing the metamorphic foliation in Levels 2 and 3. This episode is interpreted to occur at c. 35 Ma (Vannay & Hodges 1996; Godin et al. 2001; Godin 2003). Southward extrusion of the Greater Himalayan sequence, producing top-to-thesouth flow structures in Levels 2 and 3, accommodated by the STDS and the MCT. During extrusion, the Greater Himalayan sequence is progressively exhumed, and, at the current exposure level, the detachments evolve from ductile (Chame detachment) to brittle (Phu detachment) behaviour. The Manaslu leucogranite is emplaced in two phases (22 and 19 Ma; Guillot et al. 1994; Harrison et al. 1999), concomitant and slightly prior to motion on the Chame and Phu detachments, respectively. The Phu detachment has to be younger than 19 Ma, since it cuts the youngest intrusive phase of the Manaslu leucogranite (Searle & Godin 2003). Large-amplitude buckling of the Greater Himalayan sequence, the STDS (Chame and Phu detachments), and the overlying Tethyan sedimentary sequence forming the Mutsog and Chako folds.

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE D4

D5

North-dipping, top-to-the-south-verging brittle-ductile thrust structures, and associated folds in the Tethyan sedimentary sequence, localized on the northern limb of the Chako antiform. Localized north-south striking, steep brittle faults with minor offset and spaced brittle cleavage, associated with the onset of east-west extension at c. 14 Ma (Coleman & Hodges 1995).

Timing constraints U - P b geochronology U - P b geochronology was undertaken to: (1) assess the age of the granitic gneiss, and test the field interpretation that this unit is correlative with the Unit III Ordovician augen gneiss mapped further south in the Marsyandi and Kali Gandaki valleys (Coleman 1996; Godin et al. 2001); and (2) constrain the age of leucogranitic injections and deformation of the dykes within Level 1 rocks, as seen on the vertical east face of Chhubche (Fig. 4a). Although four samples were initially chosen for monazite and zircon analyses (see location on Figs 2 & 3a; results are presented in Table 1), only two samples (N-11b, N-22b) yielded reliable zircon results. Monazites were not recovered in any samples, most probably due to the predominance of calcareous rocks in the Nat valley that may have inhibited crystallization of monazite. Most of the analysed zircons contain very significant inheritance, except for fraction Z-1 in sample N-22b (concordant at 20.05 ___0.05 Ma). Most of the inheritance is c. 470 Ma, which suggests the Nar metamorphic rocks are likely related to the uppermost part of the Greater Himalayan sequence (e.g. Godin et al. 2001). Sample N-11b was collected from the granitic augen gneiss unit, at 4250 m elevation above Chako, on the east side of Nar Khola (Fig. 2). The sample was collected less than 1 m above a high-strain zone, defined by biotite schist (Fig. 9a). In thin section, the sample is dominated by quartz + plagioclase + biotite, and contains a strong foliation defined by proto-ribbon textures and preferred alignment of biotite (Fig. 5b, c). The quartz proto-ribbons contain dynamically recrystallized quartz with systematic north-dipping subgrain boundaries, consistent with other top-tothe-south shear-sense indicators visible in outcrop (Fig. 5a). The three zircon fractions analysed contain between three and eleven crystals with varying morphology including equant (Z-3) and

279

flat (Z-2) grains and broken tips (Z-1) from crystals; our intention in selecting these fractions was to reduce or eliminate the influence of inherited cores. Inheritance is least likely in the flat grains and tips, and most likely in equant grains, which have the oldest 2~176 age. Regression of the flat grains and tips yields an upper intercept age of 476.1 __+3.6 Ma with a lower intercept of 29 _+ 34 Ma. We interpret the upper intercept age as the protolith age (Fig. 9b). Sample N-22b is from a boudinaged and folded 1.5-m-thick pegmatitic leucogranite dyke, injected in deformed hornblende-biotite gneisses (Fig. 2), visible on the cliffs north of Kyang (Fig. 9c). In thin section, the sample displays coarse (c. 1 mm) equigranular plagioclase + quartz + tourmaline + muscovite. The sample is not foliated, and does not contain any visible systematic array of subgrain boundaries. Four zircon analyses were produced, with between five and 14 crystals each, one of which (Z-I) produced a perfectly concordant abraded high-precision analysis with a concordia age of 20.05 ___0.06 Ma, which is interpreted as the age of crystallization. All other fractions contained inherited zircons of a wide age range. Analyses of two other samples were attempted: a weakly deformed tourmaline leucogranite dyke (N-6.4), and an undeformed quartz-feldspar porphyry (N-6) late dyke. This would have placed constraints on the minimum age of structures, but unfortunately every single analysis contained appreciable inheritance and no precise age information was obtained, though the data are presented in Table 1. The age of inherited zircons from N-6 appears to be 4 6 9 _ 12 Ma (lower intercept 21 ___ 39 Ma), which is consistent with inheritance from the early Palaeozoic gneisses. The inheritance in sample N-6.4 is much more diverse and includes middle Proterozoic ages. In summary, sample N-11b contains mainly mid-Ordovician zircons with either Pb-loss during Himalayan events, or thin overgrowths of the same age. This confirms that the augen gneiss is equivalent to Unit III exposed elsewhere in the Annapuma range (Hodges et al. 1996; Godin et al. 2001). The deformed leucogranitic dyke (N-22b) yielded a concordant c. 20 Ma zircon analysis, which is interpreted to represent the age of the melt, with some older zircons indicating a Proterozoic source, most probably inherited from the Greater Himalayan sequence into which the dyke was intruding.

4~

thermochronology

Twenty-three mineral separates from 15 samples were analysed by the 40 At/ 39 Ar method to constrain the youngest part of the tectonothermal history,

280

L. G O D I N ETAL.

-.~

9

r...) ~

t,,eq

d

r

~3

eq

e-.

[...

k. "*~ t ~

o ',, ,..--,

I

~ ~'q'~

~

~

',~" ~ w~

t"q o~,_"

,.-,

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE (b)

281

data-point error ellipses are 68% conf

0.076

N-11b 460/~-""

0.074-

0.072-

~.

4 ~ " "

"

~73----

0.070-

0.068-

420J

0,0660.064

Intercepts at & 476.1 ~_ 3.6 [+6.01Ma

.-'\ Zl

MSWD = 0.000 2~

,""

0.062 ....... I 0,49

(d) 0.08

0.06

./

,/'P.,."" ~

I

I

0.51

I

l

0.53

1

0.55

I

I

I ....

0.57

I

0.59

da!a-point error ellipses are 68.3% conf

,,,,,,,,

N-22b Concordia age = 20.05 + 0.03 Ma 1 (1~, decay-const, errs included) / MSWD (of concordance) = 0.31, Probability (of concordance) = 0 . 5 8 4

0.04

2,/| .

200

0,02

.~

/ k" . / loo/Z2-1 l y

" " ""

zl/! /

0o0' 0.0

0,2

0.4

0.6

!

I

Fig. 9. U-Pb geochronology results from the Nar valley. (a) Field photograph of sample N-1 lb, showing the granitic augen gneiss, resting above its high-strain equivalent biotite schist. Hmnmer is 40 cm long. (b) Concordia plot of sample N-1 lb; the protolith age is best interpreted with a discordia line connecting fraction Z1 with Z2, yielding an upper intercept age of 476.1 + 3.6 Ma. (c) Field photograph of sample site of N-22b, showing the folded and stretched nature of the leucogranitic dyke. (d) Concordia plot of sample N-22b; fraction Z1 yielded a concordant age of 20.05 + 0.06 Ma.

and to complement the U - P b data obtained in the Nar valley. Muscovite, biotite and hornblende mineral separates were selected to obtain a range of cooling ages with respect to the different isotopic closure temperatures of the respective minerals. Results (Table 2, Figs 3a & 10) are presented, interpreted and compared to published 40Ar/ 39Ar thermochronologic investigations in the Marsyandi area (e.g. Coleman & Hodges 1995, 1998), and in the Kali Gandaki area (Vannay &Hodges 1996; Godin et al. 2001). The T h - P b ion microprobe ages on monazites from the Manaslu leucogranite define two pulses of magmatism at 22.9 +_ 0.6 Ma and 19.3 + 0.3 Ma (Guillot et al. 1994; Harrison et al. 1999). Peak

metamorphic conditions of the uppermost Greater Himalayan sequence (650~ Gleeson & Godin 2006) are interpreted to have resulted from the Neohimalayan thermal pulse at 24-20 Ma (Coleman 1998; Godin et al. 2001), coeval with emplacement of the Manaslu leucogranite (c. 700-775~ Guillot et aL 1994; Scaillet et al. 1995). The obtained 4~ ages are, for the most part, younger than 20 Ma, and therefore interpreted as cooling ages. The consistently small age differences between the Manaslu leucogranite, the leucogranitic dykes in the Nar (Sample N-22b at 20.05 4-0.05 Ma), and the hornblende, biotite and muscovite ages require rapid cooling from > 700~ to c. 300~ in less than 5 million years (c. 80~ years). For these rapid cooling rates,

L. GODIN ETAL.

282

Table 2. Summary of 39Ar/4~ results from the Nar valley Sample

Rock unit

Mineral*

T-6 N-38 N-40 N-34 L-9 T-21 T-48 T-134 T-105 N-102 L-12 T-141 L-6 T-6 N-11 N-34 N-40 L-9 T-48 N-38 T-134 N-102 T-105 T-147

Bt schist (Unit III) Pelitic schist (Unit V) Augen gneiss (Unit III) Bt-Ms schist (Unit III) Marble (Unit IV) Marble (Unit IV) Pelitic schist (Unit V) Pelitic schist (Unit V) Pelitic schist (Unit V) Pelitic schist (Unit V) Manaslu granite TSS Augen gneiss (Unit III) Bt schist (Unit III) Augen gneiss (Unit III) Bt schist (Unit III) Augen gneiss (Unit III) Marble (Unit IV) Pelitic schist (Unit V) Pelitic schist (Unit V) Pelitic schist (Unit V) Pelitic schist (Unit V) Pelitic schist (Unit V) TSS

Hbl Hbl Ms Ms Ms Ms Ms Ms Ms Ms Ms Ms Bt Bt Bt Bt Bt Bt Bt Bt Bt Bt Bt Ms (WR)

Plateau Total age * (Ma) ___2o- steps 24.9 + 0.6* 16.5 + 0.6 16.9 + 0.2 16.6 + 0.4 16.0 + 1.3 N/A 16.3 + 0.5 16.3 + 0.2 15.5 + 0.5 '~ 14.8 i 0.3 19.3 + 0.2 15.9 + 0.9 16.0 + 0.3 16.6 + 0.3 17.5 __+0.3 16.1 + 0.6 23.1 -I- 0.1 18.1 -I- 0.2 18.2 -4- 0.8 16.9 + 0.6 16.1 _+ 0.2 14.2 _-/- 0.3 14.3 ___0.2 24.5 ___0.7

6 8 7 9 7 N/A 6 13 5 5 13 7 9 8 8 9 8 7 4 7 8 9 9 5

% 39Ar (spectrum)

Isochron age* Isochron (Ma) + 20- MSWD*

91.8 95.3 74.3 97.7 94.1 N/A 71.9 100 89 70.4 100 100 96.3 100 80.5 100 82.1 83.5 43.9 90.5 93.7 100 100 76.2

24.8 16.9 16.9 16.7 16.1 16.3 12.7 16.0 15.8 14.8 19.0 14.7 15.7 16.8 17.9 16.9 23.1 18.3 19.6 16.0 16.4 14.3 14.5 24.8

+ 1.0 + 1.1 + 0.4 + 0.8 + 2.4 + 2.5 + 6.4 + 0.6 + 0.8 + 1.0 + 0.6 + 6.1 + 1.8 + 0.7 + 0.9 + 2.0 ___0.3 __+0.9 _ 1.9 + 1.8 + 0.9 __+0.7 ___0.7 + 2.3

1.15 0.74 0.58 0.62 0.92 0.61 0.65 0.72 1.50 1.40 0.98 1.30 0.90 1.17 0.15 2.00 1.9 0.84 0.35 1.06 0.48 0.70 1.20 1.50

*Bt, biotite; Hbl, hornblende;Ms, muscovite *Ages used in the interpretationare in bold *Mean square weighteddeviation (MSWD) calculatedfollowingRoddick (1987) w temperature steps were lost in the initial run. Age reflects the plateau steps of the first run, combinedwith the secondrun results

we estimate closure temperatures based on published diffusion data as follows: hornblende 535 ___ 50~ (Harrison 1981); muscovite 370 __+50~ (Lister & Baldwin 1996); biotite 335 ___ 50~ (Harrison et al. 1985; Grove & Harrison 1996).

Hornblende 4~

ages

Two hornblende mineral fractions were analysed (T-6 and N-38). Sample T-6 is from the lowest exposed structural level (Level 1), in the core of the Chako antiform, structurally below the Chame detachment (Figs 2 & 3). The sample yielded a fiat release spectrum with a plateau age of 24.9 ___ 0.6 Ma, within error of the inverse isochron age (Table 2; Fig. 10). Sample N-38 was collected above the Chame detachment, in the immediate footwall of the overriding Phu detachment (Fig. 2). The obtained spectrum defines a plateau age of 16.5 ___ 0,6 Ma, consistent with the inverse isochron age (Table 2; Fig. 10). Both samples have well-constrained, near-modern-day ratios of 2 9 6 _ 11 and atmospheric 4~ 288 + 17 defined through regression analysis (modern atmosphere = 295.5), with MSWDs of 1.15 and 0.74, respectively (Table 2).

Muscovite 4~

ages

A total of ten muscovite mineral factions were analysed. They were collected (Fig. 2) from Unit III augen gneiss (N-40) and highly strained equivalent biotite-muscovite schist (N-34), Unit IV (L-9, T-21) and Unit V (T-48, T-134, N-102, T-105). Two additional samples are from the Manaslu leucogranite (L-12), and from the Triassic unit of the Tethyan sedimentary sequence, in the immediate hanging wall of the Phu detachment above Nar village (T- 141). Both samples from Unit III, N-40 (granitic augen gneiss) and N-34 (highly sheared granitic gneiss), yield very good plateau ages of 16.9 +__0.2 Ma and 16.6 _+ 0.4 Ma, respectively (Fig. 10), and well-constrained initial atmospheric 4~ ratios (286 ___ 51 and 293 __+ 23, respectively; Table 2). Since their muscovite cooling ages are within error identical, the strain fabric responsible for the textural difference between samples N-40 and N-34 must therefore be older. Samples from Unit IV (Ms-Bt marble) did not yield robust plateau ages. Sample L-9 yields a poorly constrained plateau age of 16.0 _+ 1.3 Ma, with an initial atmospheric 4~ ratio of

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE

283

Fig. 10. 4~ release spectra diagrams of hornblende, muscovite and biotite samples used for thermochronology. The shaded steps within the release spectra indicate the increments used for calculating plateau ages. See text for explanation and discussion of data.

284

L. GODIN ET AL.

293 _ 22 (Table 2; Fig. 10). In contrast, sample T21 did not yield a plateau age due to excess argon, as indicated by the high initial atmospheric 4~ ratio of 350 ___23 (Table 2). By ignoring the low and high temperature steps, an inverse isochron age of 16.3 _+ 2.5 Ma is obtained, consistent with the age of nearby samples. Both samples from Unit IV were collected on the northern limb of the Chako antiform, within a marble mylonite zone that displays significant top-to-the-south thrust shear-sense indicators. Samples from Unit V (T-48, T-134, N-102, T-105) yielded muscovite cooling ages within error of each other. Sample T-48 has a good plateau age of 16.3 ___0.5 Ma (with 72% of 39mr), although the obtained inverse isochron age is poorly constrained ratio of with a high initial atmospheric 4~ 383 ___ 160 (Table 2; Fig. 10). Samples T-134 and T-105 are from the northern and southem limbs of the Mutsog synform, respectively. Sample T-134 yielded a very robust 16.3 ___0.2 Ma plateau age, in agreement with the inverse isochron age (initial atmospheric 4~ = 307 _ 21). Sample T-105 yielded a similar age, within error, to T-134, of 15.8 • 1.5 Ma; however, some problems were encountered during the initial run, resulting in the loss of the higher-temperature steps. The interpreted age reflects the plateau steps of the first run, combined with the heating steps from the second run, plotted on the inverse isochron diagram (initial atmospheric 4~ = 276 _+ 20; Table 2; Fig. 10). Both plateau and inverse isochron ages are in agreement, supporting this interpretation. Sample N-102 is from an intensely crenulated schist, located in the core of the Mutsog synform (Figs 2 & 3). A reliable plateau age of 14.8 _+ 0.3 Ma was obtained, with a slight increase in age in the higher temperature steps (initial atmospheric 4~ = 295 • 61" Table 2). Sample L-12 is from the northeastern part of the Manaslu leucogranite (Fig. 2), on the south side of the Pangre glacier, in the immediate footwall of the Phu detachment (Searle & Godin 2003). The sample yielded a very robust plateau age of 19.3 _ 0.2 Ma, in agreement with the inverse isochron age (initial atmospheric 4~ = 311 • 32; Table 2; Fig. 10). The obtained age is identical to monazite U - T h ages of the second pulse (Bimtang phase) of Manaslu magmatism at 19.3 • 0.3 Ma (Harrison et al. 1999). The last muscovite sample (T-141b) is from the immediate hanging wall of the Phu detachment, in shales of the Triassic Thini Formation, just above Nar village (Fig. 2). The sample yielded a plateau age of 15.9 • 0.9 Ma, in agreement with the inverse isochron correlation diagram (initial atmospheric 4~ = 288 + 34; Table 2; Fig. 10).

Biotite 4 ~

ages

Eleven biotite samples were analysed. Five of these are from Unit III granitic augen gneiss (L-06, N- 11 and N-40) and high-strain equivalent biotite schist (T-06 and N-34). Sample L-09 is from the marble mylonite zone just north of Kyang. The remaining five samples are from Unit V (T-48, N-38, T-134, N-102 and T-105). Since it is common for Himalayan biotites to contain unresolved excess argon (Roddick et al. 1980; McDougall & Harrison 1988), it was anticipated that the biotite samples would yield anomalously old age spectra compared to muscovite (e.g. Godin et al. 2001). However, although some biotite samples displayed anomalously old characteristics, most biotite ages are similar or younger than muscovite ages, with good plateau and inverse isochron correlation diagrams, supporting the quality of the data. The three biotite samples from Unit III yield plateau ages of 23.1 • 0.1 Ma (N-40), 17.5 • 0.3 Ma (N-11) and 16.0 • 0.3 Ma (L-06). Although all three results agree with their respective inverse isochron correlation diagrams (Table 2; Fig. 10), the age of the biotite in sample N-40 is distinctly older than muscovite in the same sample (16.9 _ 0.2 Ma), which suggests there might be unresolved excess argon. The result from sample N- 11 is also slightly older than muscovites analysed from Unit III (see above), while sample L-06 yields a biotite age in agreement with muscovite results. We therefore cautiously disregard the results of N-40 and N-11. The two samples from the biotite schist yield flat release spectra with plateau ages of 16.6 _ 0.3 Ma (T-06) and 16.9 _+ 0.6 Ma (N-34), in agreement with their inverse isochron ages (initial atmospheric 4~ = 286 _+ 23 and 278 + 32 for T-06 and N-34, respectively; Table 2; Fig. 10). Sample L-09 yields a plateau (83.5% of 39Ar) age of 18.1 + 0.2 Ma, coinciding with the inverse isochron age (initial atmospheric 4~ 287 + 53; Table 2; Fig. 10). However, muscovite analysis from the same sample yields an age 2 million years younger, suggesting this sample might also contain excess argon. Four of the five pelite schist samples from Unit V yield very good plateau ages. The two southern samples yield young robust plateau ages (100% of 39Ar) of 14.2 _+ 0.3 Ma (N-102) and 14.3 + 0.2 Ma (T-105), in agreement with their inverse isochron ages (initial atmospheric 4~ ratio= 292 _ 32 and 284 _+ 33 for N-102 and T-105, respectively; Table 2; Fig. 10). Samples N-38 and T-134 produced older good plateau ages of 16.9 _ 0.6 Ma (90.5% of 39Ar) and 16.1 _ 0.2 Ma (93.7% of 39Ar), respectively. Their initial atmospheric 4~ ratios are well-constrained

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE (310 __+29 for sample N-38, and 277 • 51 for sample T-134; Table 2; Fig. 10). Sample T-48 did not produce a reliable plateau or inverse isochron age (Table 2; Fig. 10). This sample contained significant excess argon, and is therefore discarded in our interpretations. In summary, the Greater Himalayan sequence that crops out in the Nar valley yields hornblende ages ranging from 24.9 to 16.5 Ma, muscovite ages between 16.6 and 14.8 Ma, and biotite ages between 16.9 and 14.2 Ma. However, except for one c. 25 Ma hornblende and two c. 14 Ma biotite ages, all samples yield average ages systematically clustered around 16 Ma. We therefore interpret the 4~ ages to represent fairly homogeneous and rapid cooling of the Nar gneisses at c. 16 Ma.

285

younger than 19 Ma (Searle & Godin 2003), since it cuts the youngest intrusive phase of the Manaslu leucogranite dated at 19.3 + 0.3 Ma (Harrison et al. 1999), and possibly the earlier Chame detachment. Motion on the Phu detachment was followed by large-amplitude buckling of the Greater Himalayan sequence, which folded the South Tibetan detachment system (Chame and Phu detachments), and the overlying Tethyan sedimentary sequence. The Mutsog and Chako folds are surface expressions of this event exposed in the Nar valley. Finally, renewed south-verging deformation created brittle-ductile thrust structures and associated folds in the Tethyan sedimentary sequence. We suggest that the northern limbs of the folded Greater Himalayan sequence provide natural ramps to localize these out-of-sequence thrusts.

Discussion D e f o r m a t i o n f e a t u r e s e x p o s e d in the

I m p l i c a t i o n s f o r cooling rate o f the

N a r valley

Greater Himalayan sequence

The data presented in this paper confirm that the Phu detachment footwall rocks correlate with the Greater Himalayan sequence. Units II and III are very similar to the Greater Himalayan sequence described further south in the Marsyandi and Kali Gandaki valleys, as well as in the Annapurna sanctuary (Coleman 1996; Hodges et al. 1996; Godin et al. 2001). Garnet-biotite thermometry on Unit V rocks indicates temperatures reached 650~ (upper-greenschist to amphibolite-facies metamorphism), confirming that the Chame detachment does not mark a metamorphic discontinuity at the top of the Greater Himalayan sequence (Gleeson & Godin 2006). Rather, it is a shear zone entirely positioned within the Greater Himalayan sequence, as suggested by Searle & Godin (2003). U - P b geochronology of the Unit III granitic augen gneiss reveals a clear Ordovician origin, very similar to previously dated Unit III augen gneiss in the Kali Gandaki valley (c. 480 Ma; Godin et al. 2001). Structurally, the metamorphic rocks display penetrative top-to-thesouth ductile flow structures, locally concentrated in high-strain zones, similar to Greater Himalayan sequence rocks exposed elsewhere in the Himalayan metamorphic core (Hodges et al. 1996; Vannay & Hodges 1996). Although structural Levels 1 and 2 are separated by the top-to-thenorth Chame detachment, both levels are internally dominated by top-to-the-south ductile flow structures. During southward extrusion, the fault system was progressively exhumed, and the topto-the-north detachment system evolved from ductile (Chame detachment) to brittle (Phu detachment) behaviour. The Phu detachment has to be

Figure 1 l a shows three different cooling curves, established by compiling available ages from U - P b geochronology obtained on zircon and monazite from melts, combined with 4~ cooling ages of hornblende, muscovite and biotite. In order to facilitate comparison, only samples collected from the upper levels of the Greater Himalayan sequence (Units II and higher) were considered. The Kali Gandaki data are composed of U - P b monazite ages, and hornblende, muscovite and biotite 4~ cooling ages (Vannay & Hodges 1996; Godin et al. 2001). They indicate a first thermal pulse at c. 35 Ma (Eohimalayan phase), interpreted to represent initial thickening following continental collision (Vannay & Hodges 1996; Godin et al. 2001). After a c. 15 million years isothermal period, the Greater Himalayan sequence underwent a second thermal pulse (Neohimalayan phase), coincident with sillimanite-grade metamorphism (Vannay & Hodges 1996; Godin et al. 2001). The Greater Himalayan sequence then cooled homogeneously, without any major structural disruption, through the muscovite cooling temperature between 15 and 13 Ma. The cooling part of the Kali Gandaki curve represents a cooling rate of c. 37~ years (Fig. 1 la). The Manaslu cooling curve combines T h - P b ages of monazites from the Manaslu leucogranite (Harrison et al. 1999) with muscovite, biotite and alkali feldspar 4~ cooling ages (Copeland et al. 1990). Following emplacement of the youngest phase of the Manaslu leucogranite at c. 19.3 Ma (Harrison et al. 1999), the pluton cooled at a rate of c. 82~ years until c. 13.5 Ma (Copeland et al. 1990). The obtained cooling curve shows a

286

L. GODIN E T AL.

Fig. 11. (a) Temperature-time plot for the uppermost Greater Himalayan sequence (GHS) in the Nat valley. Kali Gandaki data from Vannay & Hodges (1996) and Godin et al. (2001); Manaslu data from Copeland et al. (1990) and Harrison et al. (I 999). The temperature estimates for the Eohimalayan and the Neohimalayan events are taken from Vannay & Hodges (1996). See text for discussion. Ky, kyanite; Sill, sillimanite. (b) Schematic block diagram of present-day geometry of folded upper surface of the South Tibetan detachment system (Phu detachment), in relation to the Manaslu leucogranite.

faster period of exhumation (or unroofing) during the mid-Miocene, compared to the Greater Himalayan sequence in the Kali Gandaki (Fig. 11 a). The Nar cooling curve is obtained from compilation of the 24 hornblende, muscovite and biotite

4~ cooling ages presented in this paper, combined with the zircon age of sample N-22b. The upper part of the Greater Himalayan sequence in the Nar valley must have been at c. 750~ during crystallization of the leucogranitic dyke

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE (L-22b) at 20.05 Ma (Fig. 9d), and cooled through the biotite closure temperature (335 ___50~ at c. 16 Ma (Table 2). This represents a fast cooling rate of c. 80~ years, similar to the calculated cooling rate of the Manaslu leucogranite. However, the rapid cooling, and hence exhumation, of the Greater Himalayan sequence in the Nat valley occurred 2 - 3 million years prior to exhumation of equivalent structural-level rocks now exposed in the Kali Gandaki valley. Figure 1 lb is a three-dimensional block diagram depicting the geometric relationships between the Manaslu leucogranite and the Greater Himalayan sequence of the Nar and Kali Gandaki valleys. The diagram also shows the folded nature of the upper strand of the STDS surface. In the Kali Gandaki valley, the STDS surface dips to the NE, then folds around the Annapurna massif into a series of synforms and antiforms, accounting for the northern step of its surface trace (Fig. 1a, b). The Kali Gandaki cooling data come from the 'frontal' previously well-described part of the Greater Himalayan sequence. The exhumation of the Greater Himalayan sequence in the Kali Gandaki is interpreted to be the result of tectonic denudation from top-to-the-north motion on the STDS, assisted by erosion of the south-facing side of the Annapurna Range. In contrast, the Greater Himalayan sequence in the Nar valley, exposed in the core of the Chako antiform, cooled 2 - 3 million years earlier, and at a faster rate, than the 'frontal' Kali Gandaki Greater Himalayan sequence. North-propagating tectonic denudation assisted by erosion would produce younger cooling ages towards the north. In contrast, the Greater Himalayan sequence in the Nar valley cooled before the frontal Kali Gandaki rocks. We propose that exhumation of the Nar valley rocks resulted from tectonic denudation associated with northward motion on the STDS, and assisted by buckling of the Greater Himalayan sequence shortly after normal-sense motion ceased on the Phu detachment. The Phu detachment cuts the youngest phase of the Manaslu leucogranite, and must therefore be younger than 19.3 Ma (Searle & Godin 2003). This buckling folded the Phu detachment, and probably initiated slightly prior to - or during cooling at c. 16 Ma, which corresponds with the 40 Ar/ 39 Ar ages of most of the hornblende, muscovite and biotite in the Nar valley (Table 2). Buckling causing rapid exhumation is consistent with the fast cooling rates of the Manaslu leucogranite and the Greater Himalayan sequence in the Nar valley. Implications for timing and mechanics of cessation of southward extrusion

The Greater Himalayan sequence extruded southward during the Miocene along two opposite-sense

287

shear-zone boundaries, the structurally lower MCT and structurally higher STDS (Hubbard & Harrison 1989; Searle & Rex 1989; Hodges et al. 1992, 1996). The geometry of this extrusion is commonly envisioned as either a 'channel' (Grujic et al. 2002; Beaumont et al. 2001) or a wedge (Burchfiel & Royden 1985; Hodges et al. 1992; Grujic et al. 1996; Grasemann et al. 1999). In both end-member geometries, the extrusion is possibly rendered feasible by Poiseuille-type flow in which the pressure gradient, induced by variation in crustal thickness at the southern edge of the Tibetan Plateau, produces highest velocities in the centre of the channel and opposite vorticity for the top and bottom of the channel (see review in Godin et al. 2006). Our thermochronology data indicate that folding of the Greater Himalayan sequence quickly followed cessation of motion along the upper extrusion boundary (Phu detachment), and is similar to data from Bhutan (Grujic et al. 2002). In central Nepal, the cessation of southward extrusion occurred between 19 and 16 Ma, prior to growth of the Lesser Himalayan duplex in late Miocene-Pliocene time (Robinson et al. 2003). This does not preclude the possibility that some segments of the MCT and/or STDS may have remained active after this time, as suggested by Harrison et al. (1997) and Hurtado et al. (2001); however, these younger fault motions are not related to synchronous and dynamically linked extrusion along the MCT-STDS couple (Godin et al. 2006). Fold-and-thrust belts can be modelled in terms of a deforming wedge, propagating at a taper angle 05 (Davis et al. 1983). In order to maintain a critical taper angle (05c), changes in 05 may instigate hinterland deformation (when 05 < 05c) or foreland deformation (05> 05c). The buckling of the Greater Himalayan sequence could therefore indicate a northward step in the Himalayan deformation, back-stepping away from the southwardpropagating thrust system. Figure 12 portrays a suggested sequence of early to middle Miocene events that might have generated the geometry of structures observed in the Nar valley. We suggest that the hinterland, out-of-sequence buckling is controlled by a decrease in taper angle of the deforrohlg wedge at the leading edge of the Himalayan deformation front (Fig. 12), a similar scenario to the dynamic compensation model proposed by Hodges et al. (1996). It is suggested that rapid early extrusion during the coupled MCT-STDS phase in the early Miocene occurred at the critical taper angle (05c), where the angle is maintained at equilibrium by southward ductile extrusion of mid-crustal material between the MCT and the lower strand of the STDS (Chame-Deurali-Annapurna detachment), balanced by frontal erosion on the southern slopes of the Himalaya (Fig. 12a).

288

L. GODIN ETAL.

Fig. 12. Schematic evolution of the Greater Himalayan sequence (GHS) in the Nar valley and variation in the taper angle (05) of the deforming Himalayan wedge. (a) Early Miocene southward extrusion of the GHS by coeval motion along the Main Central thrust (MCT) and the ductile South Tibetan detachment system (STDS) fault (Chame detachment) at critical taper angle (4u (b) Decrease of the taper angle at c. 19 Ma, due to brittle activity along the structurally high Phu detachment, and sustained erosion on the southern deforming front of the Himalaya. (c) Increase and return to critical angle of the deforming taper by buckling the hinterland, favouring rapid exhumation of the GHS now exposed in the Nar valley. (d) Continued shortening generates and localizes top-to-the-south thrusts on the north-dipping limbs of the buckle folds. The return to critical angle of the deforming wedge favours foreland propagation, and development of the Main Boundary thrust (MBT) in mid-Miocene.

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE At c. 19 Ma, the angle of the deforming wedge decreased and slip in the STDS migrated structurally higher from the Chame detachment to the Phu detachment, as the Greater Himalayan sequence extruded and cooled (Fig. 12b). The decrease in angle may have resulted from continued activity along the ductile-brittle strand of the STDS (Phu detachment), while MCT motion ceased. Alternatively, the decrease in angle of the deforming wedge may have resulted from a combination of slowing extrusive flow rates and/or an increase in erosion. The 87Sr/86Sr ratio in marine carbonate sediments dramatically increases at c. 18Ma, which is interpreted to represent a sharp increase in weathering denudation rates in the Himalaya (Richter et al. 1992), and may be related to strengthening of the Indian Monsoon (Raymo & Ruddiman 1992; Molnar et al. 1993). If continued activity along the brittle strand of the STDS (Phu detachment) persisted while motion on the MCT ceased, this would imply that the younger normal-sense fault motion resulted in true horizontal extension. This would also imply that motion along the different strands of the STDS may have different causes. Although the lower ductile strand (Chame detachment) may be linked to footwall extrusion of the Greater Himalayan sequence beneath a 'fixed' Tethyan sedimentary sequence, the upper, younger and brittle strand could be linked to 'true' extension, associated with far-field critical taper adjustments. Between 19 and 16 Ma, top-to-the north normal motion along the Phu detachment ceased, and the Greater Himalayan sequence in the Nar valley was buckled. The Chako antiform, combined with subsequent valley incision, forms an apparent structural window in the Greater Himalayan sequence. The 'hinterland' buckling accelerated exhumation of the Manaslu leucogranite and the Greater Himalayan sequence now exposed in the Nar valley, and contributed to an increase of the taper angle (Fig. 12c). Continued shortening between 16 and 14 Ma localized renewed top-to-the-south thrusting on the north-dipping limbs of the folded Greater Himalayan sequence, locally reactivating as thrust faults older top-to-the-north (previously normalsense motion) fault segments, such as the Phu detachment (Fig. 12d). This late thrusting must have occurred after folding of the Greater Himalayan sequence, but before onset of eastwest brittle extension at c. 14 Ma, as documented in the Marsyandi valley (Coleman & Hodges 1995).

Conclusion We suggest that ductile extrusion of the Greater Himalayan sequence occurred at c. 25-19 Ma above the ductile (structurally higher) MCT and

289

below the ductile strand of the STDS (Chame detachment). Peak temperatures were attained within the Greater Himalayan sequence from c. 35 Ma, and temperatures remained high until the end of crustal melting that generated the Manaslu leucogranite and the Chako dyke network at c. 20 Ma. Structural cross-cutting relationships, regional cross-sections and restoration, and metamorphic petrology indicate that southward extrusion of the Greater Himalayan sequence terminated with cessation of normal-sense motion on the brittle upper strand of the STDS, the Phu detachment, at c. 19 Ma. This was followed almost immediately by crustal-scale buckling, resulting in rapid cooling of the Greater Himalayan sequence now exposed in the metamorphic culmination of the Nar valley. The thermochronology of samples from the the Nar valley indicates that cooling occurred very rapidly, 2 - 3 million years earlier than in the frontal part of the extruding Greater Himalayan sequence. Although extrusion in the frontal part of the slab must have locked at the same time as in the Nar valley, the exhumation and cooling of the gneisses was slower, most probably assisted only by erosion, rather than by rapid folding as in the case in the Nar valley samples. This buckling also indicates a step northward in deformation of the Himalayan belt, which may be a response to a lower taper geometry in the foreland due to the erosion that might have initiated and driven extrusion in the first place. Invaluable field assistance was provided by C. Olsen, N. Portelance and P. Tamang and his friendly crew. Funding from a Natural Science and Engineering Research Council of Canada operating grant and graduate scholarship to L.G. and T.G., respectively, and a Natural Environment Research Council grant to M.S. are gratefully acknowledged. Many ideas expressed in this paper benefited from discussions with R. Brown, D. Kellett and K. Larson. The first expedition to the Nar valley in 2000 was made possible by a Natural Science and Engineering Research Council of Canada operating grant to R. Brown, of Carleton University. Reviewers M. Hubbard, S. Guillot and editor R. Law are thanked for constructive comments on an earlier draft of this manuscript. This paper is dedicated to the memory of Pasang Tamang (1960-2005).

References ARITA, K. 1983. Origin of the inverted metamorphism of the lower Himalayas, central Nepal. Tectonophysics, 95, 43-60. BEAUMONT, C., JAMIESON,R. A., NGUYEN,M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742.

290

L. GODIN ET AL.

BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & MEDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109. DOI: 10.1029/ 2003JB002809. BORDET, P., COLCHEN, M., KRUMMENACHER, D., LE FORT, P., MOUTERDE, R. & REMY, M. 1971. Recherches gdologiques dans l'Himalaya du Ndpal: rdgion de la Thakhola. Centre de la Recherche Scientifique, Paris. BORDET, P., COLCHEN, M. & LE FORT, P. 1975. Recherches gdologiques dans l'Himalaya du Ndpal: rigion du Nyi-Shang. Centre de la Recherche Scientifique, Paris. BOUCHEZ, J.-L. & Pt~CHER, A. 1976. Plasticit6 du quartz et sens de cisaillement dans les quartzites du Grand Chevauchement Central Himalayen. Bulletin de la Socidtd Gdologique de France, 18, 1377-1385. BROWN, R. L. & NAZARCHUK,J. H. 1993. Annapurna detachment fault in the Greater Himalaya of central Nepal. In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 461-473. BRUNEL, M. 1986. Ductile thrusting in the Himalayas: shear sense criteria and stretching lineations. Tectonics, 5, 247-265. BURCHFIEL, B. C. & ROYDEN, L. H. 1985. Northsouth extension within the convergent Himalayan region. Geology, 13, 679-682. BURCHFIEL, B. C., CHEN, Z., HODGES, K. V., LIu, Y., ROYDEN, L. H., DENG, C. & Xu, J. 1992. The South Tibetan Detachment System, Himalaya Orogen: Extension contemporaneous with and parallel to shortening in a collisional mountain belt. Geological Society of America, Special Paper, 269. BURG, J.-P. 1983. Carte Giologique du sud du Tibet. Ministry of Geology Peking and Centre National de Recherche Scientifique, Paris. CABY, R., PI~CHER, A. & LE FORT, P. 1983. Le grand chevauchement central himalayen: nouvelles donn~es sur le m&amorphisme inverse a la base de la Dalle du Tibet. Revue de Gdologie Dynamique et de G~ographie Physique, 24(2), 89 - 100. CATLOS, E. J., HARRISON, T. M., MANNING, C. E., GROVE, M., RAI, S. M., HUBBARD, M. S. & UPRETI, B. N. 2002. Records of the evolution of the Himalayan orogen from in situ Th-Pb ion microprobe dating of monazite: eastern Nepal and western Garhwal. Journal of Asian Earth Sciences, 20, 459-479. COLCHEN, M., LE FORT, P. & PgCHER, A. 1981. Geological map of the Annapurnas-Manaslu-Ganesh Himalaya of Nepal. In: GUPTA, H. K. & DELANY, F. M. (eds). Zagros - Hindu Kush - Himalaya: A Geodynamic Evolution. American Geophysical Union, Geodynamics Series 3, Map scale 1:200,000. COLCHEN, M., Le FORT, P. & PI~CHER, A. 1986. Recherches giologiques dans l'Himalaya du Ndpal: , Annapurna - Manaslu - Ganesh Himal. Editions du centre National de la Recherche Scientifique, Paris.

COLEMAN, M. E. 1996. Orogen-parallel and orogenperpendicular extension in the central Nepalese Himalayas. Geological Society of America Bulletin, 108, 1594-1607. COLEMAN, M. E. 1998. U-Pb constraints on Oligocene-Miocene deformation and anatexis within the central Himalaya, Marsyandi valley, Nepal. American Journal of Science, 298, 553-571. COLEMAN, M. E. & HODGES, K. V. 1995. Evidence for Tibetan plateau uplift before 14 Myr ago from a new minimum age for east-west extension. Nature, 374, 49-52. COLEMAN, M. E. & HODGES, K. V. 1998. Contrasting Oligocene and Miocene thermal histories from the hanging wall and footwall of the South Tibetan detachment in the central Himalaya from 40Ar/39Ar thermochronology, Marsyandi valley, central Nepal. Tectonics, 17, 726-740. COPELAND, P., HARRISON, T. M. & LE FORT, P. 1990. Age and cooling history of the Manaslu granite: implication for Himalayan tectonics. Journal of Volcanology and Geothermal Research, 44, 33-50. DAVIS, D., SUPPE, J. & DAHLEN, F. A. 1983. Mechanics of fold-and-thrust belts and accretionary wedges. Journal of Geophysical Research, 88, 1153-1172. FUCHS, G., WIDDER, R. W. & TULADHAR, R. 1988. Contributions to the geology of the Annapurna range (Manang, area, Nepal). Jahrbuch der Geologischen Bundesanstalt, Wien, 131, 593-607. FUCHS, G., REGMI, K. & SCHILL, E. 1999. Note on the geology of the Mar - Manag region in northern Nepal (Himalaya). Abstract volume, 14th Himalaya-Karakorum-Tibet workshop, Kloster Ettal. Terra Nostra, 2, 46-47. GARZANTI, E. 1999. Stratigraphy and sedimentary history of the Nepal Tethys Himalaya passive margin. Journal of Asian Earth Sciences, 17, 805 -827. GARZANTI, E., GORZA, M., MARTELLINI, L. & NICORA, A. 1994. Transition from diagenesis to metamorphism in the Paleozoic to Mesozoic succession of the Dolpo-Manang Synclinorium and Thakkhola graben (Nepal Tethys Himalaya). Eclogae Geologicae Helvetica, 87, 13-632. GLEESON, T. & GODIN, L. 2006. The Chako antiform: a folded segment of the Greater Himalayan sequence, Mar valley, Central Nepal Himalaya. Journal of Asian Earth Sciences, 27, 717-734. GODIN, L. 2001. The Chako dome: an enigmatic structure in the hanging wall of the South Tibetan detachment, Mar valley, central Nepal. Journal of Asian Earth Sciences, 19, 22-23. GOD1N, L. 2003. Structural evolution of the Tethyan sedimentary sequence, central Nepal Himalaya. Journal of Asian Earth Sciences, 22, 307-328. GODIN, L., BROWN, R. L. & HAMMER, S. 1999. High strain zone in the hanging wall of the Annapuma detachment, central Nepal Himalaya. In: MACFARLANE, A. M., SORKHABI, R. & QUADE, J. (eds) Himalaya and Tibet: Mountain Roots to Mountain Tops. Geological Society of America, Special Paper 328, 199-210.

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE GODIN, L., PARRISH,R. R., BROWN,R. L. & HODGES, K. 2001. Crustal thickening leading to exhumation of the metamorphic core of the central Nepal Himalaya: Insight from U-Pb geochronology and 4~ thermochronology. Tectonics, 20, 729-747. GODIN, L., GRUJIC, D., LAW, R. D. & SEARLE, M. P. 2006. Channel flow, ductile extrusion and exhumation in continental collision zones: an introduction. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion, and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 1-23. GRADSTEIN, F. M., VON RAD, U., GIBLING, M. R. Er AL. 1992. The Mesozoic continental margin of central Nepal. Geologisches Jahrbuch, 77. GRASEMANN, B., FRITZ, H. & VANNAY, J.-C. 1999. Quantitative kinematic flow analysis from the Main Central thrust zone (NW-Himalaya, India): implications for a decelerating strain path and the extrusion of orogenic wedges. Journal of Structural Geology, 21, 837-853. GROVE, M. & HARRISON, T. M. 1996.4~ diffusion in Fe-rich biotite. American Mineralogist, 81, 940- 951. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KONDIC, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extension of the Higher Himalayan Crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. GRUJIC, D., HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel: insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. GUILLOT, S., P~CHER, A., ROCHETTE, P. & LE FORT, P. 1993. The emplacement of the Manaslu granite of Central Nepal: field and and magnetic susceptibility constraints. In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 413-428. GUILLOT, S., HODGES, K., LE FORT, P. & PI~CHER, A. 1994. New constraints on the age of the Manaslu leucogranite: evidence for episodic tectonic denudation in the central Himalayas. Geology, 22, 559-562. HARRISON, T. M. 1981. Diffusion of 40Ar in hornblende. Contributions to Mineralogy and Petrology, 78, 324-331. HARRISON, T. M., DUNCAN, I. & McDOUGALL, I. 1985. Diffusion of 4oAr in biotite: Temperature, pressure and compositional effects. Geochimica et Cosmochimica Acta, 49, 2461-2468. HARRISON, T. M., RYERSON, F. J., LE FORT, P., YIN, A., LOVERA, O. M. & CATLOS, E. J. 1997. A late Miocene-Pliocene origin for the Central Himalayan inverted metamorphism. Earth and Planetary Science Letters, 146, E l - E 7 . HARRISON, T. M., GROVE, M., MCKEEGAN, K. D., COATH, C. D., LOVERA, 0. M. & LE FORT, P. 1999. Origin and episodic emplacement of the Manaslu intrusive complex, central Himalaya. Journal of Petrology, 40, 3-19.

291

HODGES, K. V. 2000. Tectonics of the Himalaya and southern Tibet from two perspectives. Geological Society of America Bulletin, 112, 324-350. HODGES, K. V., PARRISH, R. R., HOUSH, T. B., Lux, D. R., BURCHF~L, B. C., ROYDEN, L. H. & CHEN, Z. 1992. Simultaneous Miocene extension and shortening in the Himalayan orogen. Science, 258, 1466-1470. HODGES, K. V., PARRISH, R. R. & SEARLE, M. P. 1996. Tectonic evolution of the central Annapuma Range, Nepalese Himalayas. Tectonics, 15, 1264-1291. HUBBARD, M. S. & HARRISON, T. M. 1989. 4~ age constraints on deformation and metamorphism in the Main Central Thrust Zone and Tibetan Slab, eastern Nepal Himalaya. Tectonics, 8, 865-880. HURTADO Jr, J. M., HODGES, K. V. & WHIPPLE, K. X. 2001. Neotectonics of the Thakkhola graben and implications for recent activity on the South Tibetan fault system in the central Nepal Himalaya. Geological Society of America Bulletin, 113, 222-240. LE FORT, P. 1975. Himalayas: the collided range. Present knowledge of the continental arc. American Journal of Science, 275-A, 1-44. LE FORT, P. 1981. Manaslu leucogranite: a collision signature of the Himalaya - A model for its genesis and emplacement. Journal of Geophysical Research, 86, 545-568. LE FORT, P., P~CHER, A. & UPRETI, B. N. 1986. A section through the Tibetan Slab in central Nepal (Kali Gandaki valley): mineral chemistry and thermobarometry of the Main Central thrust zone. Science de la Terre, M~moire 47, 211-228. LE FORT, P., GUILLOT, S. & P~CHER, A. 1999. Une carte g6ologique de l'Himlung Himal, massif du Manaslu. La Montagne et Alpinisme, 95, 22-27. LISTER, G. S. & BALDWIN, S. L. 1996. Modelling the effect of arbitrary P-T-t histories on argon diffusion in minerals using the MacArgon program for the Apple Macintosh. Tectonophysics, 253, 83-109. McDOUGALL, I. & HARRISON, T. M. 1988. Geochronology and Thermochronology by the 4~ Method. Oxford University Press, New York. MOLNAR, P., ENGLAND, P. & MARTINOD, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the indian monsoon. Reviews in Geophysics, 31, 357-396. MURPHY, M. A. & HARRISON, T. M. 1999. Relationship between leucogranites and the Qomolangma detachment in the Rongbuk Valley, south Tibet. Geology 27, 831-834. PASSCHIER, C. W. & TROUW, R. A. J. 1996. Microtectonics. Springer-Verlag, Berlin. P~CHER, A. 1975. The Main Central Thrust of the Nepal Himalaya and the related metamorphism in the Modi-Khola cross-section (Annapurna range). Himalayan Geology, 5, 115-131. PI~CHER, A. 1991. The contact between the Higher Himalaya crystallines and the Tibetan sedimentary series: Miocene large-scale dextral shearing. Tectonics, 10, 587-598.

292

L. GODIN ETAL

RAYMO, M. E. & RUDDIMAN, W. F. 1992. Tectonic forcing of Late Cenozoic climate. Nature, 359, 117-122. RICHTER, F. M., ROWLEY, D. B. & DE PAOLO, D. J. 1992. Sr isotope evolution of seawater - the role of tectonics. Earth and Planetary Science Letters, 109, 11-23. ROBINSON, D. M., DECELLES, P. G., GARZIONE, C. N., PEARSON, 0. N., HARRISON, T. M., & CATLOS, E. J. 2003. Kinematic model for the Main Central thrust in Nepal. Geology, 31, 359-362. RODDICI,:, J. C. 1987. Generalized numerical error analysis with applications to geochronology and thermodynamics. Geochimica et Cosmochimica Acta, 51, 2129-2135. RODDICK, J. C., CLIFF, R. A. & REX, D. C. 1980. The evolution of excess argon in Alpine biotites - A 4~ analysis. Earth and Planetary Science Letters, 48, 185-208. ROYDEN, L. H. 1996. Coupling and decoupling of crust and mantle in convergent orogens: implications for strain partitioning in the crust. Journal of Geophysical Research, 101, 17679-17705. SCAILLET, B., Pt~CHER, A., ROCHETTE, P. CHAMPENOIS, M. 1995. The Gangotri granite (Garhwal Himalaya): Laccolithic emplacement in an extending collisional belt. Journal of Geophysical Research, 100, 585-607. SCHILL, E., APPEL, E., GODIN, L., CROUZET, C., GAUTAM, P. & REGMI, K. 2003. Record of deformation by AMS and secondary inclinations in the Nar/Phu valley (central Nepal). Tectonophysics, 377, 197-209. SCHNEIDER, C. & MASCH, L. 1993. The metamorphism of the Tibetan Series from the Manang area, Marsyandi valley, central Nepal. In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 357-374. SEARLE, M. P. & GODIN, L. 2003. The South Tibetan detachment system and the Manaslu leucogranite: a

structural reinterpretation and restoration of the Annapurna - Manaslu Himalaya, Nepal. Journal of Geology, 111, 505-524. SEARLE, M. P. & REX, A. J. 1989. Thermal model for the Zanskar Himalaya. Journal of Metamorphic Geology, 7, 127-134. SEARLE, M. P., PARRISH, R. P., HODGES, K. V., HURFORD, A., AYRES, M. W. & WHITEHOUSE, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya: Field relations, geochemistry, age, origin, and emplacement. Journal of Geology, 105, 295-317. SEARLE, M. P., SIMPSON, R. L., LAW, R. D., PARRISH, R. R. & WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal - South Tibet. Journal of the Geological Society, London, 160, 345-366. VANNAY, J.-C. & GRASEMANN, B. 2001. Himalayan inverted metamorphism and syn-convergence extension as a consequence of a general shear extrusion. Geological Magazine, 138, 253-276. VANNAY, J.-C. & HODGES, K. V. 1996. Tectonometamorphic evolution of the Himalayan metamorphic core between the Annapurna and Dhaulagiri, central Nepal. Journal of Metamorphic" Geology, 14, 635-656. VIDAL, P., COCHERIE, A. & LE FORT, P. 1982. Geochemical investigations of the origin of the Manaslu leucogranite (Himalaya, Nepal). Geochimica et Cosmochimica Acta, 46, 2279-2292. WEISMAYR, G. & GRASEMANN, B. 2002. Eohimalayan fold and thrust belt: Implications for the geodynamic evolution of the NW-Himalaya (India). Tectonics, 21, 1058. DOI: 10/1029/2002TC001363. YIN, A. 2003. Reinterpreting the South Tibet Detachment as a passive roof fault in a large scale faultbend fold system during Miocene evolution of the Himalayas. Abstract volume, 18th HimalayaKarakorum- Tibet workshop, Monte Verita, Switzerland, 131.

Mechanisms and timescales of felsic magma segregation, ascent and emplacement in the Himalaya B. S C A I L L E T ~ & M. P. S E A R L E 2

llnstitut des Sciences de la Terre, UMR 6113, CNRS-UO, l a rue de la Fdrollerie, 45071 Orleans, cedex 2, France (e-mail: [email protected]) 2Department o f Earth Sciences, Oxford University, Parks Road, Oxford, OX1 3PR, UK Abstract: We combine field, petrological, geochemical and experimental observations to

evaluate the timescales of compaction-driven and shear-assisted melt extraction and ascent in the Himalaya. The results show that melt migration via compaction and channelling is inescapable and operates on timescales of less than 1 million years and possibly as short as 0.1 million years. Field and petrological data show that such a fast and efficient melt transfer results from a combination of favourable factors, including: (1) low but constant melt viscosity (104.5 Pa s) during extraction and ascent; (2) grain size coarsening of the source rocks in response to prolonged heating prior to melting; and (3) high source fertility and thus high melt fraction, owing to elevated modal amounts of muscovite in leucogranite sources. All three factors dramatically increase source permeability. Calculations show that shear-assisted melt extraction had a time interval recurrence in the range 10 000-100 000 years (10-100 ka), leading to sill thicknesses of 130 m. Yet melts falling at the low end of the viscosity range when coupled to high shear velocities may lead to veins several hundred metres thick. The deepest structural levels (e.g. central Zanskar Range) show that in-situ melts formed where pure shear compaction was greatest and where simple shear was also operative. Magma extracted from migmatite leucosomes was injected along planes of weakness parallel to the ductile shear fabric, probably by some form of hydraulic fracturing crack propagation mechanism. Large High Himalayan leucogranite (HHL) bodies (e.g. c. 5 km thick sills at Manaslu, Makalu and northern Bhutan) may thus represent inflated laccoliths assembled via dykes that tapped a 100-300 m melt layer produced by compaction of the Greater Himalayan Series (GHS). Thermal simulations show that such melt layers may have incubation times of several million years. Although transport time for magmas associated with the HHL is short, the time for assembly may take several million years for the largest HHL, as geochronological data indicate (up to 5 million years for Manaslu, Shisha Pangma). Transport of leucogranite melt from mid-crustal levels towards the surface was concomitant with active low-angle normal faulting along the South Tibetan Detachment (STD) normal fault, a structure that effectively formed the lid to the extrusion of a partially molten layer of mid-crustal rocks (channel flow). Rapid cooling of the granites emplaced at the top of the GHS implies rapid extrusion and lateral flow of GHS rocks beneath the STD during the period c. 20-17 Ma. Weakening of the crust by partial melting is thus likely to be pulsatory in time, and future thermomechanical models should incorporate such aspects to model tectonic evolution of hot orogens.

The role of melt in orogenic processes has received increasing attention in recent years, with adoption of the basic concept that processes of melt formation and continental deformation have a strong feedback relationship (Hollister & Crawford 1986; Hollister 1993; Nelson et al. 1996; Brown & Solar 1998a). Recent field and thermomechanical models of the H i m a l a y a - T i b e t orogen have in particular suggested that melt-assisted middle crustal flow plays a central role in the a c c o m m o d a t i o n of large horizontal and vertical displacements during continental collision (Bird 1991; Clark & Royden 2000; B e a u m o n t et al. 2001, 2004; Grujic et al. 2002). Aside from issues specific to the Himalayan context, the processes of m a g m a extraction, collection and migration in the continental crust have

been the subject of numerous field-based (e.g. Brown & Solar 1998b; Searle 1999a, b; Sawyer 2001) and theoretical (e.g. Clemens & Mawer 1992; Brown et al. 1995; Petford 1995; Vigneresse et al. 1996; Cruden & McCaffrey 2001; Jackson et al. 2003; Bons et al. 2004; Rabinowicz & Vigneresse 2004) studies in recent years, testifying to the complexity of the processes involved and h o w imperfect our current understanding of these processes is. In the present paper we focus on the Himalaya and use its exceptionally well exposed and constrained geological framework to discuss the possible mechanisms of melt extraction, segregation ascent and e m p l a c e m e n t in continental crust. The mechanisms of melt segregation and e m p l a c e m e n t

From: LAW, R. D., SEARLE,M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 293-308. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

294

B. SCAILLET & M. P. SEARLE

in the Himalaya have been addressed in a number of previous works (e.g. Guillot et al. 1993; Scaillet et al. 1995b; Searle 1999a, b), but only on a qualitative basis. On the other hand, timescales of magma segregation and emplacement in the Himalaya have already been considered, but mostly using geochronological or geochemical arguments (e.g. Deniel et al. 1987; Copeland et al. 1990; Ayres et al. 1997; Harris et al. 2000). Here we evaluate the timescales of melt extraction and accumulation in the Himalaya, by combining available experimental constraints on the petrogenesis of the High Himalayan leucogranites (HHL) and theoretical models of magma segregation and migration. This paper is, however, not intended to provide a general review of available models on melt extraction from partially molten crust. Many excellent works on this topic are available (e.g. Wickham 1987; Miller et al. 1988; Bergantz & Dawes 1994; Brown et al. 1995; Petford 1995; Rubin 1995; Vigneresse et al. 1996; Sawyer 2001; Bons et al. 2004). We note only that Jackson et al. (2003) have quantitatively modelled the case of partial melting being triggered by basalt intrusion in the lower crust. This work shows that, in contrast to early views (e.g. Wickham 1987), compaction may be an efficient mechanism for felsic melt segregation from partially molten crust and can occur on timescales as short as 4 ka. Although the boundary conditions of this study do not apply to the Himalaya, the results prompted us to reconsider compaction as a possible mechanism for melt extraction in the Himalaya using the general equations of McKenzie (1985). Similarly, the recent theoretical work of Rabinowicz & Vigneresse (2004) provides a general physical framework for assessing the role of simple shear in segregating melt from a partially molten source, which we apply below to the Himalaya. We do not address the transition from solid to liquid behaviour in partially molten systems (i.e. the concept of rheological critical melt percentage), which has been, and is still, a matter of intense debate (see Mecklenburgh & Rutter (2003) for a recent review). We simply assume that the melt produced in the source migmatite is allowed to move freely once a percolation threshold is attained (at melt fraction higher than 4%; Laporte et al. 1997). We start by briefly reviewing HHL petrogenesis together with their general geometric relationships with their surroundings. We then evaluate the timescale of magma extraction from the partially melted protolith by considering two end-member models, one via compaction and the other via shearing. We speculate that both mechanisms operate along the Himalayan chain, but that their relative importance may vary along-strike and with time. We then consider the timescale of magma cooling.

We end by outlining some general consequences that extraction of magma from its source may have on the development of an orogen, considering in particular the channel flow model as currently proposed for the Himalaya (e.g. Beaumont et al. 2001, 2004; Jamieson et al. 2004).

Geological setting and petrogenesis of HHL The High Himalayan leucogranites are, together with the North Himalayan leucogranites (NHL), one of the rare magmatic products of the Himalaya orogen (Le Fort 1981), with crystallization ages spanning 28-12 Ma (see compilations in Searle et al. 1997, 2003; Harrison et al. 1998; Zhang et al. 2004). Recent U - P b dating of several North Himalayan leucogranites has shown that their ages are not younger than the HHL as previously thought (Harrison et al. 1997) but overlap closely with ages from the HHL (Lee et al. 2000, 2006; Zhang et al. 2004). Figure 1 summarizes the present-day general geometric attributes of the HHL in relation to their host rocks. The HHL are intruded into the upper structural levels of a highgrade metamorphic sequence that includes metapelites, metagreywackes and marbles-calc-silicates, referred to as the Greater Himalayan Sequence (GHS). The structural thickness of the GHS varies from 5 to 25 km along-strike and the HHL are always intruded along the upper structural levels of the GHS and never along the basal levels. The GHS is truncated at its base by the Main Central Thrust (MCT), and along its top, above the HHL, by the South Tibetan Detachment (STD). Slip on the MCT and STD are believed to have been broadly contemporaneous (e.g. Hodges et al. 2001; see also review by Godin et al. 2006). In detail, HHL relationships to host rocks show either a clear intrusive character, with sharp contacts, or more diffuse boundaries with an intermediate zone rich in enclaves of host rocks (Fig. 2). Both types of contact are locally present on the upper and lower surfaces of leucogranite sills (Fig. 2). The deepest structural levels of the GHS contain a structurally thick migmatite zone and migmatite leucosomes can be traced amalgamating into narrow sills and dykes that feed the larger scale sills. Many of the largest HHL, such as those at Shisha Pangma (Searle et al. 1997), Everest-Makalu (Searle 1999a, b; Searle et al. 2003) and Kangchenjunga (Searle & Szulc 2005) have been fed directly from these giant sill complexes. The HHL display a variety of textures ranging from a weak or near-isotropic magmatic fabric, generally found in the thickest intrusions, to a penetrative mylonitic fabric notably observed in sills

MAGMATIC PROCESSES IN THE HIMALAYA

295

Fig. 1. Block diagram showing the different structural positions of Himalayan leucogranites within the Greater Himalaya Series (GHS), lower magmatic source, and feeder dykes. For clarity, the thickness of the zone in between the base of the laccoliths and the compacted melt layer has been increased. Dotted circles illustrate the horizontal extent of zones of compacted layer melt which are tapped by dykes feeding overlying plutons. HHL, High Himalayan Leucogranites; MCT, Main Central Thrust; STD, South Tibetan Detachment; TSS, Tibetan Sedimentary Series.

intruded close to the STD (e.g. Burg et al. 1984; P~cher 1991; Carosi et al. 1998; Murphy & Harrison 1999). Structural studies show that the leucogranite magmatic fabric is characterized by a significant component of pure-shear strain (e.g. Searle et al. 1993; Scaillet et al. 1995b), that can lead to dramatic large-scale boudinage structures as interpreted for the Gangotri lenses in the Garhwal Himalayas (e.g. Searle et al. 1993; Scaillet et al. 1995b). Quantitative strain and vorticity analyses also show that material flow has a strong component of pure shear strain at the top of GHS where HHL are intruded (Law et al. 2004; Jessup et al. 2006). Field observations show that most HHL bodies preserve feeder dykes along their base (Inger & Harris 1993; Searle e t a l . 1993; Scaillet et al. 1995b; Searle 1999a, b). In some areas such dyke swarms can be traced over vertical distances in excess of c. 1 km (Fig. 2), and they presumably extend to the depths where primary stage of melt collection occurred, and tapped these melt-rich zones (Fig. 1). It is not clear, however, that any single dyke has such a vertical extent. Various lines of evidence, including field (e.g. Le Fort 1981; Searle 1999a, b; Searle et al. 2003), petrological, geochemical (e.g. Vidal et al. 1982; France Lanord & Le Fort 1988; P~cher 1989;

Inger & Harris 1993; Searle et al. 1997) and experimental data (Patino Douce & Harris 1998), all indicate that HHL were produced by muscovite dehydration melting of the underlying metapelite sequence. Experimental melting of Himalayan metapelites has shown that HHL-like melts are the by-product of the following master reaction (Patino Douce & Harris 1998): 22 Ms + 7 P1 + 8 Qtz ----+25 Melt + 5 Kfs §

(1)

where stoichiometric coefficients are in wt%. As pointed out by Patino Douce & Harris (1998), the melting reaction 1 implies that the maximum amount of melt that can be produced by muscovite (Ms) breakdown is 1.14 times the mass proportion of Ms in the protolith. There are only a few published modal compositions of Himalayan metapelites (e.g. Patino Douce & Harris 1998; Annen et al. 2006), and they suggest that Ms abundance varies from 10 up to 30 wt%, hence the maximum melt fraction varies between 11 and 31 wt%. The stoichiometry of reaction 1 implies that muscovite abundance will be the limiting factor for melt production, since muscovite is always less abundant

296

B. SCAILLET & M. P. SEARLE

MAGMATIC PROCESSES IN THE HIMALAYA than either plagioclase or quartz in Himalayan metapelites. The temperature range over which reaction 1 proceeds is experimentally constrained to be 7 5 0 820~ in the pressure range 6 - 8 kbar (Patino Douce & Harris 1998). Assuming an average H20 content of metamorphic muscovite of c. 6 wt%, then reaction 1 predicts that the produced melts will have c. 5.3 wt% dissolved H20. These experiments also demonstrate that melts produced during dehydration melting of muscovite have compositions that are both remarkably constant and similar to those of HHL. On the other hand, experimental phase equilibria carried out on HHL samples constrain magma production and emplacement temperature to be in the range 750-800~ with melt water contents of 5 - 7 wt% (Scaillet et al. 1995a); such conditions are identical to those obtained from the forward experimental approach. Close agreement between forward (melting of the source) and inverse (crystallization of HHL) experimental studies strongly suggests that HHL melts were virtually free of entrained restite material during ascent through the crust, and calls upon a very efficient melt-restite segregation process at the source level. The temperature-H20 constraints have enabled direct experimental measurement of HHL viscosity and a viscosity range of 104 to 105 Pa s has been inferred for the early stages of their emplacement in the upper crust (Scaillet et al. 1996). Modelling based upon phase equilibrium constraints has shown that the HHL magma viscosity changes little during cooling, and may even decrease, over more than 80% of the crystallization interval, owing to the near-eutectic character of HHL and the incompatible behaviour of H20 during crystallization, which largely counteracts the opposing effects that a decrease in temperature and increase in crystal content both have on bulk viscosity (Scaillet et al. 1997). Therefore, it appears that HHL viscosity may be taken as remaining nearly constant, at around 104.5 Pa s, during most of their crystallization interval, which

297

simplifies the physical modelling of the rheology of such bodies (see Annen et al. 2006). Similarly, calculated viscosities for melts produced during the dehydration melting of the GHS vary little, between 105 Pa s at 750~ down to 104 Pa s at 820~ (calculated using melt compositions of Patino Douce & Harris (1998) and the viscosity model of Scaillet et al. (1996)). Thus, during melting, the melt viscosity does not increase in parallel to the increase in melt fraction, as could have been anticipated from a simple mass balance approach (i.e. progressive dilution of water with reaction progress). This is due to the near-invariant character of reaction 1: although strictly speaking mineral compositions must vary with melting, the changes remain small (i.e. increase in Mg over Fe in residual Ms). The important point, as far as rheological properties are concerned, is that muscovite buffers the melt water content at values around 5 - 6 wt%, which implies in turn that during melting water fugacity is buffered as well, as long as muscovite persists. Thus during melting of GHS the melt viscosity remains constant, or may even decrease slightly. Evidence for the near-invariant viscosity of anatectic melts has been reported by Scaillet et al. (1998), and has been explained as a natural consequence of dehydration melting reactions (Clemens & Watkins 2001). Another important point is that thermodynamic equilibrium demands that the residue does not become drier during this melting stage (that is, nominally anhydrous minerals do not lose their dissolved H-beating species). The implications of this point are discussed further below.

Melt extraction Models used below rely heavily on knowledge of transport properties of geological materials such as diffusivity, density and viscosity. While both diffusivity and density are relatively well known, the viscosity of rocks, in the solid or partly melted state, is still a matter of debate. This introduces a

Fig. 2. Field pictures showing relationships of leucogranites with host rocks at various scales, in the Garhwal Himalaya. (a) View looking NE of the Bhagirathi peaks, showing a typical HHL lense with top, side and bottom contact. The laccolith thickness is c. 900 m. (b) Base of Thaley Sagar peak, showing large host rock enclaves, more or less rotated relative to regional foliation. Field of view is c. 500 m. (c) Top of Bhagirathi II peak with abundant host rock enclaves near top of the leucogranite laccolith. The height of the cliff is c. 450 m. (d) Top contact of the Shivling-Meru leucogranite laccolith, showing straight contact with little enclaves. At this location petrographic observations show that the contact is intrusive in nature. (e) Top contact of the Brigupanth leucogranite laccolith with a more diffuse contact between leucogranite and host rocks. Note that, compared to case (c), foliation in wall-rock screens appears relatively undisturbed by granite injection. (f) South face of Shivling peak showing the upper 1 km of white leucogranite laccolith resting on a thin layer of black schist which in turn lies on top of a massive orthogneiss. On the right of the Shivling shoulder a dyke swarm is seen, cutting across the orthogneiss level and black schist layer and joining the main body of leucogranite. The left-ward deflection of dykes is interpreted as being due to late normal faulting.

298

B. SCAILLET & M. P. SEARLE

relatively large uncertainty to the values calculated. Perhaps the most contentious issue is the way viscosity of high-grade metamorphic rocks varies with partial melt content and/or the more or less anhydrous character of nominally anhydrous minerals (see Jackson et al. 2003). Several examples relevant to the Himalayan context illustrate this point. In their channel flow model, Beaumont et al. (2004) consider that melt weakening decreases the effective viscosity by one order of magnitude relative to the melt-free case, and assign a viscosity of 1019 Pa s at 750~ to meltweakened crust. On the other hand, Harrison et al. (1998) consider that such a weakening is not supported by existing experimental data and assign the crust a viscosity as given by a standard power law stress-strain relationship. In the original paper on extraction of magma by compaction, McKenzie (1985) considered the matrix to have a constant viscosity of 1018 Pa s. Yet, in a more recent study addressing the role of shearing on melt extraction, Rabinowicz & Vigneresse (2004) considered that matrix viscosity varies widely during melting owing to changes in the nature of minerals during melting as well as to their progressive dehydration, the latter resulting from a dilution of total H20. They calculated variations in matrix viscosity ranging from as low as 1015 up to 10 as Pa s. Although such behaviour may apply once the melting regime leaves the field of hydrous minerals, we contend that in the Himalayan context, both melt and matrix viscosities remain approximately constant as long as the muscoviteout reaction has not been significantly overstepped. The current lack of consensus implies that some arbitrary choice has to be made. In the following we have adopted a constant value of 10 j7 Pa s for matrix viscosity, which is in keeping with the water-rich character of the melting reaction producing the HHL and within the range of matrix viscosities considered by Jackson et al. (2003) in their numerical modelling of magma extraction from partially molten crust (and in which more details on this aspect can be found). Compaction

With the above physical constraints, the feasibility of melt collection via compaction (or pure shear) can be tested using the compaction model derived by McKenzie (1985) which assumes that melt segregation occurs within a plastically deforming matrix, driven by the density difference between melt and matrix, provided melt is interconnected. The permeability threshold of melted crustal protoliths has been inferred to be at around 3 - 4 % (Laporte et al. 1997). The modelled process is one of mass redistribution in response to density

differences, and the rate of the process is controlled by the bulk viscosity of the matrix. The volume left over by removing the melt is taken up by the compacted matrix. Ideally, if the process were to proceed to completion, the end result would be a two-layer system with a melt layer lying on top of a crystalline source layer. In practice, the timescale over which this would happen is well beyond the timescale of the thermal evolution of orogens, and basically compaction under crustal conditions is halted by exhumation (i.e. cooling). An important aspect is that such a process of granite segregation and emplacement is volume-conservative at the scale of the partially melted layer. We reproduce below the key equations. We calculate the compaction length, L,,,, which is the thickness of the compacting region of a partially melted crustal section of thickness H, at the base of the partially molten section that directly overlies a solid lower boundary: Lm = ('rlsko/'rlm) 1/2

(2)

where r/~ and T/m are the matrix and melt viscosities (Pa s), respectively, and ko is the permeability (m2) of the melt network. Permeability is calculated in this work using the equation (see Jackson et al. 2003):

ko = 1/50 (a2~3o)

(3)

in which d is the average grain size (m) and q~o is the melt fraction. The factor 1/50 corresponds to a texturally equilibrated, monominerallic, fluid-saturated mush of isotropic grains (Jackson et al. 2003). For unequilibrated aggregates values as low as 1/2500 would be best suited (see Jackson et al. 2003), yielding permeabilities two orders of magnitude lower than those calculated here and consequently longer times for melt extraction. We adopt the value of 1/50 on the basis that the protracted thermal evolution of GHS prior to melting (see below) has presumably allowed textural equilibration of the source. The parameter th is a characteristic time (s) to reduce the total amount of melt by a factor of e in the partly melted unit. When H >> Lm, as calculations below show, th is given by: th = H / W o ( I -- ~o)

(4)

where w,, is the relative velocity between melt and matrix (m s- ~) calculated as: wo = ko(1 - c~o)(~p)g/(71m~O)

(5)

where tip is the density contrast between melt and solid (400 kg m-3), and g is the)gravitational_ acceleration (taken here at 9.8 m 2 sFinally, the

MAGMATIC PROCESSES IN THE HIMALAYA amount of melt extracted hm (m) after th is given by: hm = H ~ o (1 - e - l )

(6)

Figure 3a shows the characteristic time th needed to segregate by compaction a melt layer on top of a migmatite unit either 5 or 10 lon thick, for grain sizes varying between 1 and 10 mm, and melt fractions of 0.1 and 0.2. For the above conditions, the thickness of the segregated melt layer hmvaries between 300 m and 1200 m. We note that several geophysical investigations carried out north of Mount Everest, have indicated the possible presence of melted crust at depths of 1 5 - 2 0 k m (Pham et al. 1986; Nelson et al. 1996) with local accumulations of fluid/melt pools a few hundred metres thick (Li et al. 2003) lying on top of a partially melted section, which compares well with our calculations (see also

Fig. 3. Characteristic timescales for melt extraction from partially molten crust via compaction, calculated using equations of McKenzie (1985). All calculations assume a density contrast of 400 kg m 3, and a matrix viscosity of 10 t7 Pa s. (a) Effect of grain size shown for partial melting in magma source zone with thicknesses (H) of 5 and 10 km, each for melt fractions (f) of 0.1 and 0.2 and for a melt viscosity of l04 Pa s. (b) Effect of melt viscosity. Calculations asssume a grain size of 4 ram. Also shown is the range of melt viscosities experimentally determined for Himalayan leucogranites (Scaillet et al. 1996). See text for details on assumptions and calculations.

299

Gaillard et al. 2004). For grain sizes larger than 2 mm, melt is collected on top of the partially melted layer in less than 1 million years, and could even collect in less than 100 000 years if high melt fractions are achieved in the source (Fig. 3a). Geochronologic and thermobarometric studies have revealed that in response to the early stages of crustal thickening, the protoliths of HHL have experienced high-grade metamorphism, probably close to melting conditions, during at least 10 million years prior to the melting event that produced the HHL (Hodges et al. 1994; Vance & Harris, 1999; Simpson et al. 2000). We suggest that this protracted thermal history of the metapelites allowed mineral coarsening, leading to a significant increase of permeability when melting started. Although no systematic study has been undertaken to quantify this parameter, field observations show that metamorphic rocks with grains sizes exceeding 3 - 4 m m are common in the Himalayan range (Fig. 4). Another critical factor on melt

Fig. 4. Illustration of grain size variability within Himalayan metapelites. (a) Migmatite from Zanskar. (b) Kyanite-bearing metapelite from Zanskar. Note the coarse grain size of both types of rocks. Large grain size increases rock permeability once melting occurs and facilitates melt extraction processes. See text for additional details.

300

B. SCAILLET & M. P. SEARLE

collection time is melt viscosity (Fig. 3b). Melts having a viscosity lower than 105 Pa s, such as HHL melts, are fluid enough to be extracted from their protoliths via a compaction mechanism in less than 1 million years. For a viscosity of 104 Pa s, production of a 630 m layer on top of a 5 km thick migmatite source layer with a melt fraction of 0.2 requires only 48 000,,years, rising to 480 000 years for a viscosity of 10~Pa s (calculated for an average grain size of 4 ram). The above arguments show that compaction is: (1) a plausible mechanism of magma extraction in partially molten Himalayan crust; and, (2) it operates on very short timescales (see Laporte 1994). However, if compaction is to operate alone, clearly there is a need for an additional mechanism, at least in some areas, since exposed leucogranite bodies can reach thicknesses up to 4 - 5 km, which are difficult to achieve by compaction alone. In the calculations above, the maximum sheet thickness accumulated on top of the migmatite source layer ranges from 300 m (5 km thick source melt zone) up to 1200 m (10 km thick source zone) and requires that the source zone is of similar thickness to the entire GHS. As stressed previously by Cruden & McCaffrey (2001), the relatively low melt fraction in the protolith (7 kin) may require a minimum denudation rate, on the order of 1 mm a 1, to prevent wholesale convection and allow muscovite crystallization. Yet, emplacement of such thick HHL laccoliths during normal faulting implies that the top part of the leucogranite nearly reaches the surface while its base is still fed by active intrusions. Overall, such relatively low denudation rates suggest that, when HHL were intruded, the overlying crustal column was not undergoing vigorous erosion. Within the framework of a crustal channel flow, this suggests that the zone of focused erosion during the Miocene was located to the south of the current exposures of the HHL belt. Our results also show that to explain the steep cooling histories documented in many HHL, denudation must have been active after HHL solidification, especially when they were intruded close to their source region. However, to preserve the HHL from exhumation and erosion until the present time, the average denudation rate after emplacement cannot have exceeded 0.5 mm a -1.

Much of our current understanding on h o w crust thickening and partial melting occur during continental convergence has been gained from study of the H i m a l a y a n - T i b e t a n orogen. A number of conceptual, numerical and analogue models have been proposed to explain the main tectonothermal data and morphological observations gathered over the last 50 years in this orogen (e.g. Le Fort 1975; Molnar et al. 1983; Royden & Burchfiel 1985; Jaupart & Provost 1985; Pinet & Jaupart 1987; England et al. 1992; Harris & Massey 1994; Grujic et al. 1996; Henry et al. 1997; Harrison et al. 1998; B e a u m o n t et al. 2001; Hodges et al. 2001; Kohn & Parkinson 2002). The most recent developments in the field of numerical simulations include fully coupled t h e r m a l - m e c h a n i c a l models which have in particular investigated the hypothesis that the High Himalaya range, and the adjacent Tibetan Plateau, could result from a combination of mid-crustal channel flow and focused

surface denudation, which develop as a consequence of partially molten mid-crust and enhanced surface erosion on the plateau edge, respectively (see B e a u m o n t et al. 2004; Jamieson et al. 2004). Such simulations have shown that models in which no surface denudation occurs contemporaneously with partial melting cannot achieve a tectonic style compatible with H i m a l a y a n - T i b e t a n geology (Beaumont et al. 2004). The rate of surface denudation also influences the style of large-scale deformation: low denudation rate ( < 4 m a-1) at the plateau edge during leucogranite generation favours the generation of extruded domes in the hinterland, whilst denudation rates higher than 10 m m a-1, yield a general tectonic style consistent with the observed H i m a l a y a n - T i b e t a n orogen (Beaumont et al. 2004; Jamieson et aL 2004). Partial melting has been proposed as the main cause of mechanical weakening of the midcrust (e.g. Beaumont et al. 2001, 2004), based

From: Law, R. D., SEARLE,M. P. & GODIN,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 309-326. 0305-8719/06/$15.00 9 The Geological Society of London 2006.

310

C. ANNEN & B. SCAILLET

on: (1) field recognition of Tertiary migmatiteleucogranite in the Himalaya; and (2), the results of INDEPTH geophysical surveys that revealed the probable existence of molten mid-crust beneath Tibet (e.g. Nelson et al. 1996; Brown et al. 1996; Li et al. 2003). The channel flow model predicts that the mid-crust is progressively extruded at the channel tip and is bounded by coeval shear zones with reverse (base) and normal (top) shear senses (Fig. 1). Both types of shear zone have been recognized for many years in the Himalaya, corresponding respectively to the Main Central Thrust (MCT) and South Tibetan Detachment (STD), which mark the lower and upper surfaces of the Greater Himalayan Sequence (GHS). Critical to the evaluation of such a model is the demonstration that motion on both the MCT and STD was synchronous during crustal melting, at least during that time period in which channelized crustal flow operated (see review by Godin et al. 2006b). Synchronicity between Himalayan melting and MCT slip is nowadays generally accepted, based primarily on geochronological data that show overlapping ages between peak metamorphic conditions in the GHS and crystallization ages of some High Himalayan Leucogranite (HHL) bodies (e.g.

_ _

10

Fig. 1. Sketch of the geometrical relationships between the High Himalayan leucoganites (HHL) and normalsense motion on the South Tibetan Detachment (STD), as employed in the numerical model. The maximum possible depth of the magma source in our numerical models is on the MCT. The level of intrusion of the Manaslu pluton corresponds approximately to the final depth reached during the last increment of laccolith building, so that any molten part of the body is still able to crystallize muscovite at a magmatic stage. The laccolith labelled Everest shows the initial stage of magma accretion, in the case where intrusion starts at a depth of around 20 km, as inferred for the Everest area (Searle et al. 2003). Normal-sense slip along the STD during laccolith growth will eventually bring the Everest intrusion to a structural position similar to that displayed by the Manaslu intrusion. Note that the thicknesses of laccoliths are not scaled to their ages. The horizontal dashed line shows the minimum depth for crystallization of muscovite in HHL. GHS, Greater Himalayan Sequence; MCT, Main Central Thrust; TSS, Tibetan Sedimentary Sequence.

Hubbard & Harrison 1989; Coleman 1998; Coleman & Hodges 1998). In contrast, the exact timing of slip on STD, as well as its slip rate, remain largely unconstrained and only at a few locations has the timing relationship between HHL intrusion and motion on the STD been studied in any detail (e.g. Vance et al. 1998; D6zes et al. 1999; Murphy & Harrison 1999; Searle et al. 2003). The STD is located structurally above the HHL bodies and in most cases cross-cuts the HHL bodies, imparting a penetrative, although largely subsolidus, shear fabric to the leucogranites (e.g. Burg et al. 1984; Herren 1987; Burchfiel et al. 1992; Searle et al. 1997; Murphy & Harrison 1999). However, although it is not yet clear that the STD was active during HHL emplacement (e.g. Murphy & Harrison 1999; Harrison e t al. 1999a), such a relative timing relationship is generally implicitly accepted (e.g. Hodges et al. 2001). Since the HHL are widely believed to represent the culmination of orogenic thermal evolution, their crystallization age is taken as the time at which peak metamorphic conditions were reached during convergence. Any successful time-temperature (t-T) model of the metamorphic evolution of the Himalayan range is thus committed to melting of the mid-crust at around 2 4 - 2 0 Ma, or to some melting 30 million years after the onset of the Indo-Tibetan collision. In parallel, the cooling patterns of some well-known HHL have been qualitatively used to infer the mechanism by which exhumation took place (e.g. D6zes et al. 1999; Searle et al. 1999). Fast HHL cooling rates are generally equated with tectonically driven exhumation, i.e. HHL unroofing was mainly driven by normal-sense motion on the STD. However, while the thermal evolution of the whole orogen has been more or less successfully modelled in various studies (e.g. Jaupart & Provost 1985; England et al. 1992; Harrison et al. 1998, 1999a; Beaumont et al. 2001), modelling of cooling patterns for individual HHL bodies has still not been performed, and hence the hypothesis of HHL cooling being accelerated by tectonic unroofing remains largely qualitative. In this paper we evaluate the thermal effect that normal-sense motion on the STD may have had on HHL cooling rates while the HHL were being emplaced in the mid-crust, and compare these model predictions with available t - T data, in an effort to evaluate the role of tectonic exhumation on magmatism in the Himalaya. With numerical simulation, we constrain the maximum rate of vertical movement of the HHL during their emplacement, and thus place limits on the possible rate of mid-crust exhumation due to channel flow during the time of HHL intrusions. To this end we have performed one-dimensional (1D) numerical simulations that model the thermal evolution of instantaneously emplaced small sills, which upon accretion

LEUCOGRANITES AND EXTENSIONAL FAULTS eventually give rise to thick leucogranite lenses emplaced at the top of the GHS (Fig. 2). We build upon previous similar work (Annen et al. 2006 ) in which we attempted to constrain magma injection rates in the upper crust under the assumption that emplaced magma bodies do not reach the condition of wholesale convection, and thus that the thermal regime was dominantly conductive. That HHL bodies did not convect while molten is implied by R b - S r and S m - N d isotopic studies that have shown that leucogranites preserve the heterogeneity of their metasedimentary sources (GHS) down to a scale of a few tens of metres (e.g. Deniel et al. 1987; France-Lanord et al. 1988; Inger & Harris 1993). In our previous work we have systematically explored the effects of physical parameters such as

311

thermal conductivity, magma emplacement rate, depth of intrusion, and pre-intrusion geotherm on the overall cooling rate, assuming that leucogranite intrusion growth was achieved beneath a static overlying rock column. In such a case, thermal modelling requires that the thickest leucogranites, such as the Manaslu laccolith, were emplaced over several million years, approximately 1 million years for every kilometre of emplaced leucogranite, at depths not exceeding 12 kin, in general agreement with available thermobarometric and geochronological constraints (Guillot et al. 1995; Harrison et al. 1999b). However, this provides only one endmember of the possible temporal relationships between the STD and HHL. Here we relax the assumption of a static overlying rock column, and

(a) t = t o Topographic surface .......

6t

~''"'-'J~""-'~.~

'S'

TSS

Slip vertical component -

t I ]

Depth of the first sill at t= to

MCT~ (b) t= tl

GHS Depth of the magma source at t=to

Topographic surface

~

TSS

]~ Depth of the first sill at t= t~ Depth of the first sill at t=to

I I

GHS

(c) t = t n

Depth of the magma source at t = tl ]~ Depth of the magma source at t= t o

Topographic surface Depth of the first sill at t=tn GHS

Depth of the first sill at t= t o Depth of the last sill at t= tn Depth of the magma source at t= t,, Depth of the magma source at t= to

Fig. 2. Geometry of the system as modelled. (a) At t ----to the first sill is emplaced. (b) At t = tl the second sill is emplaced. The time interval between intrusion of the two sills, i.e. between to and q, is equal to the sill thickness (50 m) divided by the magma emplacement rate. The crustal channel is extruded by channel flow, the GHS and the leucogranite are moving along the STD and the MCT. The level of the magma source, which corresponds to our lower boundary condition, is moving too. (c) At t = &, a series of n sills has been emplaced. The position of each sill depends on the respective values of the magma emplacement rate and of the vertical component of the normal-sense slip along the STD, and can be calculated with Equation 3.

312

C. ANNEN & B. SCAILLET

model the leucogranite thermal evolution in the case of a constantly moving magma roof due to normalsense slip on the STD. The thermal effect of normal faulting can be anticipated to be an increase in cooling rates, since the leucogranite at the top of the intrusion is permanently removed and replaced by colder rocks, thereby greatly inhibiting the thermal insulating effects that static host rock has on a cooling magma body.

Geological constraints Our work is centred on the Manaslu leucogranite in central Nepal where the granite exposures constrain the three-dimensional geometry of the laccolith and of its surroundings. As elsewhere in the Himalaya, the main body of the Manaslu leucogranite lies in between the GHS and the Tibetan Sedimentary Series (TSS). Depending on the north-south transect, the thickness of the GHS varies from a few kilometres up to more than 10 km, which is where the leucogranite thickens to more than 4 km (Colchen et al. 1986). The lithostratigraphic thickness of the TSS is estimated to be around 11 km (Colchen et al. 1986), yet folding and stacking of the TSS prior to, or during, leucogranite intrusion may have significantly increased the structural thickness of the sedimentary pile overlying the Manaslu laccolith during its formation (Guillot et al. 1995; Searle & Godin 2003; Godin et al. 2006a). In the Manaslu area the position of the STD is still controversial: early field studies report the detachment as lying beneath the main laccolith (e.g. Colchen et al. 1986) while recent fieldwork concluded that the main detachment occurs above it (Searle & Godin 2003; Godin et al. 2006a). Here we have explored specifically the latter situation. Our modelling efforts are guided by three types of constraints: (1) lack of convection across the laccoliths as stated above; (2) occurrence of magmatic muscovite; and (3) geochronological data. A basic assumption of the modelling is that thermal convection in the magma would have led to magma mixing and homogenization of R b - S r ratios. This assumption is supported by Jellinek et al. (1999) who showed that two magmas with low viscosity contrast are efficiently mixed for any Reynolds number, i.e. even for sluggish convection. Thus the heterogeneity of R b - S r and S m - N d in the leucogranites indicates that the conditions for large-scale convection in the magma were not reached. This is possible if the leucogranite was emplaced by accretion of discrete sills and if the time interval between two sill intrusions was long enough for each sill to significantly cool and crystallize before the next sill was emplaced (Annen et al. 2006). In addition to their isotopic heterogeneity (e.g. France-Lanord et al. 1988), HHL are petrologically

characterized by the ubiquitous presence of magmatic muscovite, which reflects the peraluminous character of their source rock. However, the fact that HHL bodies recognized so far along the 2500 km long Himalayan arc all have magmatic muscovite (e.g. Le Fort et al. 1987), imposes a severe constraint on their pressure of crystallization. The possibility that the main part of muscovite is of hydrothermal origin can be ruled out on textural grounds (e.g. Scaillet et al. 1990) and also from stable isotope constraints which show that HHL have been little, if at all, affected by pervasive hydrothermal circulation in the subsolidus state (France-Lanord et al. 1988). Several experimental and petrological works have shown that the crystallization of F-poor muscovite in leucogranite, as observed for the HHL, demands lithostatic pressures of at least 3 kbar (e.g. Benard et al. 1985; Scaillet et al. 1995). The F-content of muscovite from the Gangotri leucogranite indicates that the limiting pressure-equivalent depth could be dropped by 1.5 km at best, relative to the threshold value of 10 km adopted by Annen et al. (2006). Besides F, the presence of Mg in muscovite could significantly enhance its thermal stability in leucogranites, yet experimental studies have shown that muscovite crystallizing in those magmas is nearly a pure end-member OH-muscovite owing to both the overall low content in Mg of the bulk rock and to the prevailing low oxygen fugacity during crystallization (Scaillet et al. 1995). Here we adopt the conservative value of 8.5 km as the minimum depth of muscovite crystallization. Therefore, crystallization of a muscovite-beating and isotopically heterogeneous HHL at decreasing pressures, as implied by coeval normal-sense motion on the STD, requires that the following two criteria be met: (i) the intrusions must be shallow enough for the heat advected by the magma to be conducted away towards the Earth surface between successive injections so as to prevent onset of convection across the sill thickness (Annen et al. 2006); and (ii) any injected sill must not reach depths shallower than 8.5 km while still molten. The possible vertical component of the slip rates on the STD is strongly constrained by these two criteria. For a given initial depth of sill intrusions at the beginning of the emplacement process (Fig. 2a), the slip rate must be high enough to increase magma cooling and prevent convection, but low enough for the molten part of the HHL to stay below the depth of magmatic muscovite crystallization. We use as an example the Manaslu leucogranite (Le Fort 1981) since it is one of the most intensely studied HHL laccoliths, having in particular a well characterized geochronological pattern (Deniel et al. 1987; Copeland et al. 1990; Harrison et al. 1999b). The geochronological constraints available ( U - T h data) indicate that the 5 - 1 0 k m thick

LEUCOGRANITES AND EXTENSIONAL FAULTS Manaslu laccolith was built by several pulses of magmatic injection in the time interval 24-19 Ma (Coleman 1998; Harrison et al. 1999b). In addition to these U - T h ages, which can be taken as the crystallization age range for the Manaslu laccolith, there is an extensive set o f 4~ ages available for both the main laccolith and its immediate country rocks (Copeland et al. 1990; Guillot et al. 1994). M u s c o v i t e 4~ ages show a regular pattern of cooling ages increasing upwards across the laccolith, from 13 Ma at its base up to 18 Ma at its top (Copeland et aL 1990). The cooling age for samples at the top of the leucogranite is identical to m u s c o v i t e 40Ar/39Ar ages in the adjacent contact aureole (Guillot et al. 1994). Altogether the U - T h and 4~ ages indicate that 1 million years after cessation of magmatic activity in the Manaslu area, the top of the leucogranite was crossing the closure temperature for Ar diffusion in muscovite (generally estimated to be 350 __+50~ while the bottom must have still been at close to magmatic temperatures. In the present study, we therefore also seek to define the conditions that would generate such a geothermal gradient inside the laccolith soon after its building period. We take advantage of the result that the boundary conditions of the numerical model detailed below need not be changed to model the subsolidus thermal evolution of the top part of the laccolith; we regard this as a valid hypothesis as long as the temperature of the source (melting zone) had not significantly decayed after cessation of melting.

Numerical model Full details of the numerical model can be found in Annen et al. (2006) and only the essential features are repeated here. We simulated the growth of a

313

laccolith-shaped leucogranite by the successive emplacement of discrete sills at magmatic temperature (Fig. 2). In the model tested here, the leucogranite is emplaced at the boundary between the GHS and the Tibetan Sedimentary Series (TSS) and this boundary is marked by the STD. Thus the thickness of the TSS varies horizontally as a function of the angle of dip of the STD (Figs 1 & 2). In our 1D model, we explore the possible depth of emplacement of the leucogranite top, which also corresponds to the thickness of the overlying TSS. Since at Manaslu the youngest crystallization ages are structurally located below the oldest (Harrison et al. 1999b), in our model each successive sill is emplaced below the former one (Fig. 2). The leucogranite and the country rock layers are discretized and the temperature evolution is computed with the finite difference expression of the equation of heat balance: 6T pCp ~-[ = k V T + pL + A

(1)

where p is density, Cp is specific heat capacity, T is temperature, t is time, k is thermal conductivity, L is latent heat of fusion or crystallization, and A is radioactive heat production. Parameter values used in the model are reported in Table 1. Values taken for p and Cp are similar to those used in previous thermal modelling studies of the Himalaya (e.g. England et al. 1992; Harrison et al. 1998; Huerta et al. 1999). Values adopted for L and A are specific to the Himalayan context and are discussed in detail by Annen et al. (2006) and England et al. (1992), respectively. Owing to the importance of conductivity in thermal simulations, the value ascribed to this parameter is considered below in some detail (see also Annen et al. 2006).

Table 1. Values of the parameters used in the numerical simulations Parameter Pse pg Psi Cpse Cpg Cpst

kse kg ksl Ase Ag Asl ozv wt% 1120 b

Definition

Value

Unit

Sediment (TSS) density Leucogranite density Slab (GHS) density Sediment specific heat capacity Leucogranite specific heat capacity Slab specific heat capacity Sediment conductivity Leucogranite conductivity Slab conductivity Sediment radioactive heat production Leucogranite radioactive heat production Slab radioactive heat production Magma volumetric coefficient of thermal expansion Melt water content Sill thickness

2700 2300 2700 1000 1600 1000 3.5 3 2 1.4 x 10-6 5.3 • 1 0 - 6 2.7 • 10-6 2 • 10-5 5 50

Kg m -3 kg m -3 kg m -3 J kg-1 K-1 J kg-1 K-1 J kg- 1 K- 1 W m -1 K -1 W m - 1 K- 1 W m - ~ K- 1 W In -3 W m -3

W m -3 K-1 % m

314

C. ANNEN & B. SCAILLET

Conductivities of representative Himalayan metapelites of the HHC have room temperature conductivities in the range 3.2-5.3 W m -1 K -1 with an average of 3.7 _+ 0.8 W m - I K -1 (Annen et al. 2006). Himalayan metasedimentary rocks and leucogranite conductivities vary in the range 2.5-4 W m - a K -1, but we have fixed the thermal conductivities in our model at 3.5 W m - l K -1. Lower metasedimentary rock conductivities restrict the heat flux between the leucogranite and the topographic surface and make it very difficult to satisfy the condition of no magma convection (Annen et al. 2006). As Annen et al. (2006) have shown, since we restrict our calculation mostly to the period of magmatic growth, and owing to the relatively high pressure of magma emplacement which imposes low host-rock perrneabilities, heat advection via hydrothermal convection around the cooling leucogranite can be neglected. We take into account the latent heat released during crystallization, based on phase equilibrium experiments performed on HHL at 4 kbar (Scaillet et al. 1995). Deniel et al. (1987) found that the scale of R b - S r heterogeneities in the Manaslu laccolith was c. 100 m. Therefore, simulations were run with 50 m thick sills, with a node spacing of 25 m. We assume that the magma temperature at its source zone depth is 800~ which represent likely conditions for crustal anatexis as inferred from various lines of evidence (e.g. P~cher 1989; Montel 1993; Inger & Harris 1992; Patino Douce & Harris 1998, Harris et al. 2000; Searle et al. 2003). The top boundary condition of the model (topographic surface) is a fixed temperature of 0~ The bottom boundary condition is an initial magmatic temperature of 800~ located at depths below the topographic surface of either 30 or 40 km at the beginning of the simulation, moving upward as denudation proceeds (Figs 1 and 2). Note that these magma source depths are conservative because: (1) shallower sources results in a steeper geothermal gradient which facilitates magmatic convection; and (2) a starting depth of, say, 15 km would allow a molten intrusion to move upwards by a maximum of 6.5 km in order to allow muscovite crystallization at a magmatic stage in HHL, thus leaving little room for the coupling between tectonic exhumation and HHL formation. The initial geothermal gradient is at equilibrium. Its shape is controlled by the depth of the bottom boundary condition ( T = 800~ and by the respective thicknesses of the TSS and GHS as they have different conductivities. Initial geothermal gradients that are not at equilibrium have been tested, but the geothermal gradient tends to equilibrium rapidly enough that the exact shape of the initial geothermal gradient does not affect the results (Annen et al. 2006).

The onset of convection across sill thickness was determined by calculating the Rayleigh number Ra: R a -- P g a v A Tw3~ /zK

(2)

where av is the volumetric coefficient of thermal expansion, AT is temperature difference between the top and bottom of the fluid layer, Wm is the thickness of the magmatic layer, /x is magma viscosity, and K is thermal diffusivity and is equal to k / p C p . We assumed that the convection begins when R a exceeds 103 (e.g. Turcotte & Schubert 1982). For the calculations of Ra, we used the viscosities experimentally determined by Scaillet et al. (1996) for HHL melts and corrected for the presence of crystals (Roscoe 1952; Lejeune & Richet 1995). The crystal fraction for temperatures between solidus and liquidus are taken from Scaillet et al. (1997). The effect of tectonic unroofing was modelled assuming that either the geotherm above the leucogranite is not altered by the intrusions, or that previous intrusions have affected the overlying geotherm. The first case corresponds to the situation of a sill of small horizontal extent such that the moving overlying roof will be affected by the growing laccolith only when very close to it. The second case corresponds to the situation of an intruded sill of infinite horizontal dimension, in which case the overlying rock sequence brought to the top of the leucogranite by the normal-sense STD slip can be affected by the thermal effect of the intrusion, sometimes before it comes into contact with it. The natural situation must lie between these two end-members. We have taken final laccolith thicknesses of 5, 7.5 and 10 km, which encompass the estimated present-day variation in thickness of the Manaslu pluton (Le Fort 1981; Le Fort et al. 1987; Searle & Godin 2003). Our calculations are performed over 5 million years, which is a likely time interval of crustal melting in the Himalaya (e.g. Noble & Searle 1995), including the Manaslu area (Harrison et al. 1999b). We have performed 1D simulations because HHL have much greater lateral than vertical dimensions (e.g. Le Fort et al. 1987), a situation which, when coupled to the short time span of the calculation, allows us to neglect lateral heat transport (Annen et al. 2006). In some cases, however, we extended calculations over 10 million years. In addition to unroofing arising from normal-sense slip on the STD alone, we have also considered the effect that natural erosion might have on the overall cooling pattern, using an average erosion value of 1 mm a -1. Although both processes decrease the lithostatic load upon the leucogranite,

LEUCOGRANITES AND EXTENSIONAL FAULTS their thermal imprints on it differ: erosion removes cold surface rocks and so it does not affect the thermal gradient in the immediate vicinity of the laccolith, while tectonic unroofing permanently brings colder rocks into contact with the leucogranite, modifying the local geothermal gradient. For each given rate of denudation, the model looks for the maximum depth at which the first intrusion can occur, so as to build a laccolith of 5, 7.5 or 10 km final thickness while preventing convection over more than two sills' thickness. Since the limiting pressure-equivalent depth for magmatic crystallization of muscovite is c. 8.5 km, first injected sills are necessarily located at higher pressure depths than this boundary. The depth of the following sills is determined by the rate of the slip along the STD that moves the whole system upward, and by the erosion rate and the emplacement rate of the leucogranite: Di = Do - (QsTD -q- Qe + Qm)ti

(3)

where Di is the depth of sill i, Do is the depth of the first sill, QSTD is the vertical component of the slip rate along the STD, Qe is the erosion rate, Qm is the magma emplacement rate, and t i is the time elapsed between the emplacement of the first sill and sill i. The model output gives the depth of the last intrusion as well as the t o p - b o t t o m temperature difference at the intrusion cessation and 5 million years later.

Results S T D slip a n d e r o s i o n r a t e s

We first discuss the effect of erosion alone. Figure 3 shows, for each erosion rate we have tested, the maximum depth of the first-emplaced sill and the corresponding maximum depth of the last sill for a heterogeneous leucogranite (no convection in the magma). Without erosion, the top of a 5 km thick laccolith emplaced in 5 million years is constrained to be no deeper than 11 km, so that by the end of the intrusion period the base of the laccoliths lies at 16 km depth. This corresponds to the standard situation explored in greater details by Annen et al. (2006). As the erosion rate increases, the maximum initial depth of the first-intruded sill increases up to a depth of 18 km for erosion rates of 3 m m a -a. In parallel to this increase in depth of the first-intruded sill, there is a continuous decrease in the final depth of the last intrusion as the erosion rate increases, such that the two curves (for depth of first and last intrusions) cross over when erosion rate and magma emplacement

315

E r o s i o n rate ( m m / y r ) 1

0 0

,

I

2

I

3

I

,

I

-

4 -

Minimum emplacement

,,,~" 8 -=

level o

/(9

..._.,

' _

16 1-..

~O--Q

-

20 -

-

9 - -

- - e - -

Initial e m p l a c e m e n t Emplacement

d e p t h of t h e first sill

d e p t h of t h e last sill

Fig. 3. Relationship between the erosion rate and the maximum initial emplacement depth of the first sill and the corresponding emplacement depth of the last sill (see Equation 3). The constraint is no convection in the magma. The leucogranite thickness is 5 km and the total emplacement duration is 5 million years. The horizontal solid line shows the minimum depth of magma injection as constrained by muscovite crystallization. The denudation erosion rate is limited to 3 mm a -1 by leucogranite exhumation. rate are 1 m m a-1 (i.e. 5 km of overburden have been removed in 5 million years such that the last sill is intruded at the same depth as the first one). Beyond this intersection point, the last-intruded sill is always intruded at a shallower structural level than the first sill. For instance, for a slip rate of 2 m m a -1, the first sill is intruded at 16 kin, while 5 million years after, the last sill is intruded at a depth of c. 11 km. For erosion rates higher than 2.75 m m a -1, the last-emplaced sills are intruded at depths shallower than 8.5 km, and thus will not crystallize muscovite (horizontal solid line on Fig. 3). The maximum permissible erosion rate to intrude a 5 km thick laccolith without convection and with magmatic muscovite is 2.75 m m a-l" at such an erosion rate the starting depth for magma injection must be 17 km. Any shallower starting depth for injection at this erosion rate will allow HHL bodies to crystallize in the magmatic muscovite-absent field. We now consider the role of the STD. The case of a 5 km thick laccolith with the geotherm of overlying rocks not affected by formerly injected sills is shown on Figure 4a. Calculations of maximum depth have been performed for various STD slip rates, expressed here via their vertical component. Conversion into actual normal-sense slip rates along the fault plane can be easily done by dividing the vertical component by a sin(c0 factor, c~ being the dip angle of the STD fault (Figs 1 & 2), whose

316

C. ANNEN & B. SCAILLET Laccolith thickness = 5 km

(a)

(b) Vertical component of STD slip (mm/yr) 0

1

o!

,

2

I

,

3

I

,

Vertical component of STD slip (mm/yr)

4

I

,

I

0

1

o

,

I

2

,

I

3

i

4

I

,

I

- -

9

4-

4t

Min. emplacement level

_

- Min. emplacement lev~ ,, s"

8

_

..~ 12"--I K J

e

16

,, ,,,, ~ce

,"

~

a

12 _ i t ,

,,"

. "~t

t

/ 20"

~s, ~

-t

~'~'-'"-1-

24 J

~ ~l'-";-,m-

- -

9

24 -

Laccolith thickness = 7.5 km

(c)

(d) Vertical component of STD slip (mm/yr) 0

1

0

,

2

I

,

I

3

i

I

Vertical component of STD slip (mm/yr)

4

,

I

4--

0

t

2

3

4

, I , 1 , 1 , 1

o 4-

8J

Min. emplacement level

8

t.

2 v

-= 1 2 - -

12-,,I

O 16 '~ J

~

,,

~.

-

,,,

..I~1S

Min. emplacement level

i'_~-'-~- -

--~.-~---~

16--

m.

t"'-.

20

20--

24

24--

Laccolith thickness = 10 km

(e)

(f) Vertical component of STD slip (mm/yr) 0

I

0

1

,

I

2

i

I

3

,

I

Vertical component of STD slip (mm/yr)

4

,

I

01

0

23

,

I

,

I

4

,

I

,

I

~

4-

4--

J 8"

t ~

Min. emplacement level

m Min. emplacement level

8

~

12-~

12 ~

o

16--

16-

20--

20 - -

24--

24 - - - 9- -

Initial depth of first sill (no erosion)

- - 9- -

Initial depth of first sill (erosion=l mm/yr)

- - e - -

Depth of last sill (no erosion)

- - e - -

Depth of last sill (erosion=l mm/yr)

"/

Fig. 4. Relationships between leucogranite emplacement depths and rate of slip along the STD for different laccolith total thicknesses. The constraint is no convection in the magma and no exhumation of the top of the leucogranite while sills continue to be accreted at the bottom. When data are not plotted for a given slip rate, it means that the coupled constraints could not be satisfied. Data are shown with no erosion, and with an erosion of 1 mm a- 1. The total emplacement duration is 5 million years (a, c, e) Sills are assumed to have a small vertical extent and the geotherm above the leucogranite top is not affected by sills. (b, d, f) Horizontal extension of the sill is assumed infinite and the geotherm above the leucogranite top is thermally affected by sills. The horizontal solid line shows the minimum depth of magma injection as constrained by muscovite crystallization. The denudation rate (slip + erosion) is limited to 3 - 4 mm a-1 by leucogranite exhumation.

LEUCOGRANITES AND EXTENSIONAL FAULTS value varies between 10 ~ to 30 ~ along the Himalaya (e.g. Herren 1987; Burchfiel et al. 1992; Harris & Massey 1994; Searle et al. 1999). We recognize that the dip angle may vary with depth. However, quantifying the effect that this variation m a y have on exhumation rates is difficult and so we have considered the rate of vertical m o v e m e n t to be constant in all our simulations. Reported values in Figures 4 to 6 correspond to conditions in which the 5 - 1 0 k m thick laccolith never enters the field of magmatic convection. Another limit is that the leucogranite top is not allowed to breach the surface before the end of intrusion. Again, two sets of curves are shown, one corresponding to the starting depth of the first sill intruded, and one to the final depth of the last sill emplaced at the base of the laccolith, each calculated with and without erosion which is set at 1 m m a-1. We present first the case with no erosion. As the vertical c o m p o n e n t of the slip rate increases, the m a x i m u m starting depth of the first sill also increases down to 23 km, showing that motion on the STD accelerates cooling of the accreted magmatic pile, and thus allows initial intrusion at greater depths than with an overlying static

slip rate (mm/yr) 1

,

0

2

I

i

3

I

4

5

J IiI

iI

4 8

Minimum emplacement level c / /

/

r

/

~_16

8

,~

\

20

/ \

/

\

\

Y

/

/

Y

./

/ /

"x"

24 28 - -

9

- -

9- -

-

Initial depth of first sill (no erosion) Initial depth of first sill (erosion = 1 mm/yr)

- - ~ - -

Depth of last sill (no erosion)

- - ~ - -

Depth of last sill (erosion = 1 mm/yr)

Fig. 5. Relationship between the leucogranite emplacement depths and rate of slip along the STD with and without erosion, when the source of magma is at 40 km depth. The total thickness of leucogranite is 5 km and the total emplacement duration is 5 million years. Sill vertical dimensions are assumed small and the geotherm above the leucogranite is not affected by formerly emplaced magma. The constraint is no convection in the magma and no exhumation of the top of the leucogranite while sills continues to be accreted at the bottom. The horizontal solid line shows the minimum depth of magma injection as constrained by muscovite crystallization.

317

slip r a t e ( m m / y r ) 1 I

0

]

I

4 .-..

M i n i m u m e m p l a c e m e n t level

IL

8

E

12

Q..

! 20 m

24

m

l

m

m

Initial d e p t h o f first sill (no erosion) Initial d e p t h of last sill (no e r o s i o n )

Fig. 6. Relationships between leucogranite emplacement depths and rate of slip along the STD. The leucogranite thickness is 10 km and the total emplacement duration is 10 million years. Sills are assumed to have a small horizontal extension and the geotherm above the leucogranite top is not affected by sills. The horizontal line shows the minimum depth of magma injection as constrained by muscovite crystallization. Here, this constraint cannot be satisfied for a vertical component of the slip rate of more than 1 mm a- 1 or with erosion. The maximum slip rate is more limited than for an emplacement duration of 5 million years, because denudation lasts 10 million years. For a slip rate of 2 mm a -] the total denudation is 20 km. c o l u m n ( 1 5 k i n for the conditions shown on Fig. 4) or with erosion alone (Fig. 3). The increase in depth of the first-intruded sill is not linearly correlated to the slip rate on the STD, however, and for that specific example the m a x i m u m depth allowable by the model is slightly over 23 kin, reaching more or less constant values at slip rates higher than 2 m m a - t . As in the first example, the depth of the last intrusion increases as the STD slip rate increases, and the two curves cross over at a value of 1 m m a -1. For a slip rate of 3 m m a-~, the intrusion/crystallization process occurs within the stability field of magmatic muscovite (depth > 8 . 5 km), while at 4 m m a -1, the first intrusion occurs at around 2 2 - 2 3 k m depth, but the last intrusion occurs at 8 km. In this case the last-intruded sill will not crystallize muscovite, since its pressure depth is shallower than the experimentally constrained threshold value for magmatic crystallization of muscovite (horizontal solid lines on Fig. 4). Note that in such a case, the topmost part of the laccolith would be only 3 k m below the topographic surface at the end of magmatic intrusion. Thus, considering the constraints of muscovite stability and lack of convection only, the example illustrated in Figure 4a shows that a 5 k m thick laccolith emplaced in 5 million years,

318

C. ANNEN & B. SCAILLET

cannot have sustained STD slip rates with a vertical component higher than 4 m m a -1. If an erosion value of 1 m m a -~ is incorporated then the maximum value drops to 3 m m a --1 The case where the geotherm of overlying rocks is affected by former magmatic intrusions is shown in Figure 4b, d and e. Compared to the previous simulation, the effect of an infinite horizontal sill is to slightly decrease the maximum slip rates on the STD, by approximately 0.2-0.3 m m a 1. Increasing the laccolith final thickness to 7.5 km, assuming no temperature effect on overlying rock, decreases the maximum depth attainable by the first intrusion to slightly below 22 km (Fig. 4c). This is due to the fact that, as we maintain a constant emplacement duration at 5 million years but increase the leucogranite thickness, this in turn increases the magma injection rate to 1.5 mm a -1 so that conditions of convection are more easily attainable than in the case of a 5 km thick laccolith emplaced at a rate of 1 m m a - t . However, the maximum slip rates inferred from the necessity of crystallizing magmatic muscovite are not significantly altered relative to the first simulation, being also 3 m m a -1 with an erosion of 1 mm a -~. When the rocks above the leucogranite are thermally affected by the intrusion process (infinite sill case), it seriously restricts the possibility of having no convection while crystallizing muscovite in a thick laccolith (Fig. 4d). At the imposed emplacement rate of 1.5 m m a - ~, if there is no erosion and no slip along the STD, the emplacement depth of the top sill must be shallower than 8.5 km in order to avoid convection, and magmatic muscovite cannot crystallize. Without erosion, a minimum slip rate of c. 0.5 m m a-~ on the STD is required in order to allow muscovite crystallization in the first-emplaced sills, while at a slip rate faster than c. 3 m m a-1, the top of the granite would reach the topographic surface before the end of intrusion. When the erosion component is taken into account, no solution can be found with slip rates greater than 2 rnm a -1 . In addition, it has to be noted that there are no solutions for any laccolith whose first injection starts at depths in excess of 20 km. Beyond that depth, any leucogranite body 7.5 km thick will inevitably convect. The case of a 10 km thick laccolith, corresponding to the maximum inferred thickness for the Manaslu, is shown on Figure 4e and f. For such a HHL thickness, no solution exists for slip rates in excess of 3 m m a - 1, irrespective of small or infinite sill configuration or presence or absence of erosion. In addition, some configurations require a minimum value for slip on the STD. These are as follows: for no thermal effect produced by prior intrusions, the first sill must be injected with a vertical component of slip rate greater than 1 m m a-1 with no erosion,

or 0.5 nun a-1 with erosion. In the case of an infinite sill, there are no solutions with erosion, and without erosion slip rates on the STD only in the range 1.5-3 m m a -1 are possible. In this latter case, by the time intrusion ends, the top part of the laccolith lies only 2 km beneath the topographic surface. When the initial depth of the source region is 40 km (instead of 30 km in previous simulations), it allows us to increase by c. 1 m m a - l (Fig. 5) the threshold in denudation rates derived above, while allowing first-emplaced sills to be intruded at greater depths as well, on the order of 28 km for a 5 km final HHL thickness intruded over 5 million years. We have also simulated the case of a 10 km thick laccolith emplaced over 10 million years, maintaining a constant magma injection rate of 1 m m a -1 (Fig. 6). Solutions can be found only if the overlying country rock is not heated up during intrusion and there is no erosion. Such conditions imply a very narrow range of vertical components for slip on the STD, between 0.2 and 1 m m a -1. At lower slip rates, a significant proportion of the leucogranite body would need to be emplaced outside the stability field of magmatic muscovite. At higher slip rates the top of the leucogranite reaches the surface before the intrusion process ends. This is because the duration of crustal exhumation is twice as long as in previous simulations. For example, with a slip rate vertical component of 2 m m a - l , 20 km of overburden are removed during magma intrusion as compared to 10 km in the former examples, but the maximum depth of sill emplacement that respects the condition is less than 20 km, which results in the leucogranite breaching the surface before the end of the intrusion. Thus the possibilities for injecting a nonconvecting thick HHL body in a rapidly exhumed crustal column are seriously restricted. In summary, the simulations described above indicate that HHL thicknesses in the range 5 - 1 0 km require exhumation rates, due to the combined effects of STD slip and surface erosion, no higher than 4 mm a - l , and more likely below 3 m m a-1 if, as geological evidence suggest, the initial depth of the magma source is located at 30 km depth. At such high denudation rates, the intrusion process needs to start at depths of c. 2 2 k m , except when the final laccolith thickness is 10 km, in which case the depth of the first intruded sills cannot exceed 18 km. Similarly, thicker HHL ( > 7 km) may require a minimum denudation rate, on the order of 1 m m a - 1, to prevent wholesale convection and allow muscovite crystallization. Yet, this implies that the top part of the leucogranite nearly reaches the topographic surface when its base is still fed by active intrusions.

LEUCOGRANITES AND EXTENSIONAL FAULTS

319

Geothermal gradients We now consider the evolution of the temperature gradient across the leucogranite produced by the process of incremental laccolith growth during normal faulting and channel flow. To this end we have performed several simulations in which the vertical c o m p o n e n t of the STD slip rate is fixed at 2 m m a - l , a likely average value based on the simulation results presented above. Figures 7 to 9 show the evolution of the geothermal gradient of the modelled crustal section intruded by the leucogranite for conditions considered as being representative of the Manaslu area, and for final leuco-granite thicknesses of 5, 7.5 and 10 km, respectively. In Figure 7, the first intrusion is injected at 20 k m depth. Subsequent sills are intruded during a period of 5 million years at an intrusion rate of 1 m m a-~ so as to build a laccolith with a final thickness of 5 km. Meanwhile the entire crustal section is e x h u m e d at a rate of 2 m m a-1 owing to normalsense slip on the STD. These are conditions appropriate for preventing large-scale magmatic convection while allowing muscovite crystallization (Fig. 4). The evolving geotherm is shown at 0, 2.5 and 5 million years (Fig. 7). The first sill intruded at 20 k m (0 million years) produces only a small perturbation of the ambient geotherm and instantaneously cools to temperatures largely below the HHL solidus (c. 650~ Scaillet et al. 1995). The temperature evolution of this first sill,

Fig. 7. Evolution of temperature profiles with an initial emplacement depth of the first sill at 20 lan and an initial magma source depth of 30 km. Intrusion starts a t = 0 million years, and last until t = 5 million years. The total leucogranite thickness emplaced is 5 km. The vertical component of slip rate is 2 mm a- 1. Leucogranite emplacement rate is 1 mm a -1. Dotted areas show leucogranite depth and thickness at t = 0, 2.5 and 5 million years.

Fig. 8. Temperature profiles as in Figure 7, but with a total leucogranite thickness of 7.5 km. The total emplacement duration is 5 million years, and the emplacement rate is 1.5 mm a-1. To satisfy the constraint of no convection, the initial emplacement depth of the first sill is 15 km.

which forms the roof of the growing leucogranite body, is shown in Figure 10a. After 2.5 million years, the crustal section has been e x h u m e d by 5 k m and the laccolith thickness reaches 2.5 km. At this stage the temperature of the first-intruded sill is about 520~ (Figs 7 & 10a), assuming that the rocks overlying the leucogranite are heated up by the leucogranite (infinite sill case). After 5 million

Fig. 9. Temperature profiles as in Figures 7 and 8, but with a total leucogranite thickness of 10 km. The total emplacement duration is 5 million years, and the emplacement rate is 2 mm a -1. To satisfy the constraint of no convection, the starting emplacement depth is 10.5 km.

320

C. ANNEN & B. SCAILLET

years of laccolith growth and parallel rock exhumation, the top sill is at 10 k m depth and its temperature is calculated to be slightly below 415~ (Figs 7 & 10a). Figures 8 and 9 show the growth of 7.5 and 10 k m thick laccoliths, respectively, with a similar emplacement rate and denudation rate as in the example described above for the 5 k m body. For thicker plutons, however, the initial emplacement has to be shallower in order to respect the condition of no magmatic convection (Fig. 4), and the final position of the body is close to the topographic surface. After 5 million years of intrusion and denudation, the roof of the 7.5 k m thick leucogranite body is

at 5 k m depth and the roof of the 10 k m thick leucogranite is only 1500 m below the topographic surface. In both cases, however, the temperatures o f the structurally highest levels o f the laccolith at the end of the intrusion are largely below 350~ and therefore such model predictions conflict with available geochronological data (Copeland et al. 1990). Figure 10 shows in detail the thermal evolution of the first intruded sill (top of the laccolith) over a time interval of 10 million years, for a final laccolith thickness of 5 km. Here, the simulation considers that intrusion starts at 24 Ma (Harrison et al. 1999b), occurs during the first 5 million

Fig. 10. Temperature evolution of the first emplaced sill (roof of the leucogranite) for initial magma source depths of 30 (a, b) and 40 km (c, d). The initial intrusion depth is 20 km and the leucogranite emplacement rate is 1 mm a -1. The vertical component of the slip rate on the STD during the ieucogranite emplacement is 2 mm a-1 and there is no erosion. Intrusions start at 24 Ma and stop at 19 Ma (vertical dashed line). After 19 Ma, curves are plotted for an after-emplacement slip rate on the STD of 0, 0.5 and 2 mm a -1. (a, c) Horizontal dimension of sills is assumed infinite and the geotherm above the leucogranite top is thermally affected by sills. (b, d) Sills are assumed to have a small horizontal dimension and the geotherm above the leucogranite top is not affected by sills. The grey area represents the closure temperature of muscovite for At.

LEUCOGRANITES AND EXTENSIONAL FAULTS years, is coeval to STD slip, and that intrusions stop at 19 million years. Results are plotted for initial source depths (bottom boundary condition) of 30 (Fig. 10a, b) and 40 km (Fig. 10c, d), and for either infinite (Fig. 10a, c) or finite (Fig. 10b, d) sill geometries. Starting at 24 Ma, the first sill cools down to about 500~ (475~ if the magma source is at 40 km). In the case of infinite sills, the temperature increases slightly as the first 20 sills are emplaced, then when the injection locus progressively moves away from the top of the growing laccolith, cooling due to denudation becomes dominant and the temperature of the leucogranite roof starts to decrease regularly until 19 Ma, when intrusions stop; at this time the temperature at the top of the leucogranite is about 400~ irrespective of the depth of source. In the case of small sills, the cooling of the first sill is immediate closing temperature of and linear. The 4~ muscovite is attained at about 21 Ma (23 Ma for a source at 40 km), i.e. 3 to 5 million years before the cooling time recorded by muscovites from the top of the Manaslu leucogranite (c. 18 Ma; Copeland et al. 1990). The small sill configuration thus appears unlikely at Manaslu. After leucogranite emplacement, the thermal evolution at the top of the laccolith depends on the denudation rate. Three post-emplacement denudation rates were modelled. Denudation continues at a rate of 2 m m a-1. The cooling proceeds at a high rate. In the case of an infinite sill, the closure temperature of muscovite is reached shortly after 19 Ma. In this case, however the leucogranite roof reaches the topographic surface at 14 Ma. Denudation proceeds at the lower rate of 0.5 man a -1 so that the leucogranite reaches the topographic surface today (0 Ma). The cooling slows down after the end of intrusion. musIn the case of infinite sills, the 4~ covite closing temperature is reached at about 16 Ma (17 Ma for a source at 40 km depth). Both slip on the STD and denudation stop after the end of intrusion. This corresponds to a model where channel flow is driven by the presence of melt and thus stops when melt ceases to be generated. In the case where the country rocks were not heated during leucogranite emplacement (finite sills), these country rocks start to heat up when the normal-sense slip of STD stops and the temperatures at the top of the laccolith start to increase from 19 Ma going back across the 40Ar/ 39Ar muscovite closure temperature (Fig. 10b). This would muslead to a thermal resetting o f 4~ covite cooling ages within the laccolith. In the case of a country rock that has been

321

heated up during intrusion (Fig. 10a), the cooling becomes very slow and, in the case of an initial magma source at 30 km depth, temperatures at the top of the laccolith stays above 400~ until after 14 Ma. On this basis we conclude that, for the Manaslu laccolith: (1) an infinite sill geometry better reproduces the geochronological data; and (2) a simple scenario in which normal-sense motion on the STD occurs during magma intrusion and is then followed by surface erosion at a rate of about 1-0.5 m m a-1, can account for the first-order features of the cooling pattern recorded in the laccolith. Figure 11 shows geotherms calculated at 0, 5 and 10 million years for various conditions, including the static case with no erosion and no movement on a normal fault, which provides a limiting scenario for HHL emplacement. We simulated the conditions corresponding to: (1) a starting emplacement at 10 km depth, no STD slip, and a magma source at 30 km; (2) a starting emplacement at 20 km depth, STD active with a vertical component of 2 m m a - ~, and an initial magma source at 30 km; (3) a starting emplacement at 20 km depth, STD slip rate of 2 m m a-~, but an initial magma source at 40 km depth. In the three cases the leucogranite overlying rocks are heated up during intrusion. The initial heat content of the system is higher in case 1 than in case 2, because the T S S - G H S boundary is at 10 km depth instead of 20 km depth, which results in a higher radioactive heat production (the heat production is higher in the GHS than in the TSS; see Table 1). The lowest heat content and geotherm is in case 3, where the magma source is at 40 km depth (Fig. 1 l a). In all simulations, after 5 million years, the laccolith is 5 km thick with its top surface at 1 0 k m depth. At this time, the magma source is at 30 km depth for cases 1 and 3, and at 20 k m for case 2. The geotherms above the leucogranite in cases 1, 2 and 3 are almost superimposed, while below the leucogranite case 2 is significantly hotter, illustrating the effect of having a short source-to-intrusion distance (5 km in case 2 compared to 15 km in cases 1 and 3). The temperature at the top of the intrusion is slightly below 400~ We allowed the model to run for an additional 5 million years with a slower denudation rate of 0.5 m m a -1, as used in the previous section. For cases 2 and 3 the top of the leucogranite is now at 7.5 km depth and the magma sources are at 17.5 and 27.5 km, respectively. Again, the highest temperatures are for case 2, being still around 350~ after 10 million years. These examples therefore show that it is possible to find a set of coupled values for initial depth of intrusion and STD slip rate that broadly satisfies the geochemical, petrological and geochronological

322

C. ANNEN & B. SCAILLET

Fig. 11. Temperature profile in the crust at (a) time t ~- 0 million years (start of intrusion), (b) t = 5 million years (end of intrusion), and (e) t -- 10 million years (after 5 million years cooling). Models plotted are: case 1, a starting emplacement depth of 10 kin, no slip along STD (static model) and a magma source at 30 kin; case 2, a starting emplacement depth of 20 km, slip along STD with a vertical component of 2 mm a - 1, and a magma source at 30 km; case 3, a starting emplacement depth of 20 km, slip along STD with a vertical component of 2 mm a 1, and a magma source at 40 km. At t = 5 million years (b) the extension of the leucogranite is from 10 to 15 km depth in the three cases. For case 1, the magma source has stayed at 30 km depth, for case 2 and 3 it is at 20 and 30 km depth, respectively. For cases 2 and 3, the denudation rate after 5 million years is 0.5 m m a -1. At t = 10 million years (c) the extension of the leucogranite is soil from 10 to 15 km depth for case 1 (no denudation) and from 7.5 to 12.5 km depth for cases 2 and 3. The depth of the magma source is now at 30 km for case 1, 17.5 km for case 2, and 27.5 km for case 3.

LEUCOGRANITES AND EXTENSIONAL FAULTS constraints available at Manaslu. A scenario with a melt source at 40 km depth, first sill injected at 20 km depth, a STD vertical component of 2 mm a-1 during intrusion, followed by surface erosion rate of 1-0.5 mm a-1 broadly fulfils the geochronological, petrological and geochemical constraints. We stress, however, that it is equally possible to find depth-rate of intrusion conditions that satisfy those constraints with no unroofing (case 1; see also Annen et al. 2006). An important aspect that needs to be outlined is that the structural distance between the source of magma and its level of intrusion exerts a major control on the thermal evolution of the cooling laccolith. In particular, it appears that conditions which differ by the starting depth of intrusion, but for which the source-tointrusion distances are similar, may end with similar geotherms, if STD motion brings the deepest section (case 3) to the same final structural level as the shallowest one (case 1). Another important point is that, for thick HHL bodies intruded close to their source region, a significant reheating event is to be expected if exhumation significantly slows down after the intrusion period. We note, however, that after magma emplacement, keeping a magmatic temperature at the bottom boundary might not be appropriate as melt generation has ceased and we have no control at this time on temperature of the underlying MCT. Thus the simulations may underestimate the cooling rate.

Implications and conclusions Our modelling results shed light on the cooling patterns that are to be expected for leucogranite bodies emplaced close to an active normal fault, and constrain the t-T-P paths that could have been followed by leucogranite bodies emplaced in an active orogenic belt such as the Himalaya, if a crustal channel flow model is adopted. We do not claim, however, that our findings reproduce all details specific to that orogen. The main aspect is that, from the thermal point of view, it is possible to simulate cooling paths pertaining to HI-IL, if they are viewed as being intruded close to a major normal fault. Our results show that if HHL intrusion was coeval with STD motion, then this in turn places severe constraints on the maximum slip rates of that fault zone and on crustal channel exhumation rate in the context of channel flow. For an average HHL thickness of 5 km, it appears that denudation rates cannot have exceeded values much higher than 3 mm a-1 (Fig. 4). Taken at face value, this is significantly lower than the minimum denudation rate of 10 mm a -1 required for surface extrusion of channel material in the model of Beaumont et al. (2004) and Jamieson et al. (2004). However, at the

323

time HHL were being produced, the erosion front was located further south (see Pearson & DeCelles 2006; Robinson & Pearson 2006) and it is likely that the crustal section above the intruding leucogranite was undergoing denudation rates much lower than those prevailing at the eroded wedge of the channel (i.e. the leucogranite magmas were lying beneath the plateau). The actual slip rate on the fault plane will depend on the average dip angle of the fault zone. Field evidence generally indicates STD dip angles of 10-30 ~ (e.g. D~zes et al. 1999; Walker et al. 1999), the lower value being in general considered more representative at the regional scale (e.g. Searle & Szulc 2005). A denudation rate of 2 mm a- 1, if due to STD slip alone with a dip angle of 10~ corresponds to a slip rate of 12 mm a 1. A dip angle of 30 C would require slip rates no higher than 4 mm a-1. These values are somewhat lower than the convergence rates (c. 20 mm a -1" e.g. Harrison et al. 1998; Beaumont et al. 2001) generally used for modelling the thermal evolution of the whole orogen. Given the various uncertainties that surround each type of model, we can only state that such apparently lower values may reflect either lower convergence rates than those generally assumed, or a non-homogeneous partitioning of the convergence rate across the deforming crustal wedge. The very large HHL such as perhaps the Manaslu itself or the Monlakarchung-Pasalum intrusion body in central Bhutan (Dietrich & Gansser 1981), with thickness approaching 10 km, remain difficult to emplace if one wants to avoid largescale convection and preserve magmatic muscovite within a rapidly eroded crustal section. An increase in the total thickness of the leucogranite is attained either: (1) by increasing the magma emplacement rate, which favours convection; or (2) if the magma emplacement rate is kept constant, by an increase of the total duration of emplacement and denudation, which rapidly results in emplacement depths being too shallow to allow crystallization of magmatic muscovite, or in the leucogranite top breaching the topographic surface while the bottom is still fed by magma intrusions. Our modelling results show that such thick plutons require a very subtle combination of controlling parameters (rates of STD slip or erosion, magmatic intrusion and starting depth of injection) in order to preserve their petrological and geochemical characteristics. In general, however, laccoliths of more than 5 km thickness cannot be emplaced at levels deeper than 8.5 km without convecting, if their cooling is not accelerated by erosion or slip along a normal fault, but if the denudation rate is too high the leucogranite breaches the topographic surface before the end of the intrusion process. If this were true, then it would imply that the main parts of the

324

C. ANNEN & B. SCAILLET

thickest HHL bodies currently cropping out in the Himalaya have been so since Miocene times. This would imply that the largest bodies have been preserved from erosional processes during the last 20 million years, which appears to be highly unlikely. Although normal-sense STD slip can accelerate HHL cooling, the upper limit for the slip rate along this fault is bounded by pressure constraints imposed by muscovite. Any factor that could increase muscovite stability in HHL towards lower pressures would be likely to increase the range of permissible slip rates. As discussed above, fluorine can play such a role, yet the actual F content of HHL muscovites does not suggest that HHL were particularly F-rich (B. Scaillet, unpublished data). We suggest, however, that future petrological studies evaluate more thoroughly this important point. The maximum starting depth we infer for leucogranite intrusion lies in the range 18-23 km, for an initial depth of the magma source at 30 kin. It broadly fits with thermobarometric constraints obtained from the country rocks surrounding some HHL plutons, such as in the Everest area (Searle et al. 2003) where the maximum inferred pressure of melting is c. 7 kbar. The modelling indicates that the maximum initial depth of the first-emplaced sills is 23 km for an initial depth of the magma source at 30 kin, or 28 km for an initial depth of the magma source at 40 km, corresponding to a magma s o u r c e - m a g m a emplacement distance of at least 7 km. At the end of the leucogranite emplacement, however, if the denudation rate is several millimetres per year, the distance between the magma source and the bottom of the leucogranite can be as low as 2 km. Thus, when HHL are intruded close to their source region (at or less than 5 km above it), as envisaged in some fieldbased studies (e.g. Visona & Lombardo 2002; Searle & Szulc 2005), our results show that HHL cooling must be accelerated by denudation so as to inhibit convection across the pluton. In contrast, for HHL bodies intruded away from their source ( > 10 km above it), the previous work of Annen et al. (2006) has shown that HHL may be intruded with a static overlying roof. Yet, the present study shows that it is equally possible to reproduce the cooling pattern of HHL using a simple scenario which combines normal sense STD slip during intrusion followed by surface erosion. We stress that STD slip must continue below subsolidus conditions, otherwise it would be difficult to preserve the fast cooling rate documented in many exhumed metasedimentary sections of the GHS and the cross-cutting HHL (e.g. Searle et al. 1999). Altogether, however the average denudation rate after leucogranite emplacement cannot exceed 0.5 m m a -1 in order to preserve the leucogranite from exhumation and erosion until the present day.

The research was funded by an Ernst and Lucie Schmidehiny grant to C.A. The authors thank R.S.J. Sparks for useful discussions. The reviews of N. Harris and of an anonymous reviewer, as well as the careful editorial handling of R. Law, greatly helped us to clarify the manuscript.

References ANNEN, C., SCA1LLET, B. & SPARKS, R. S. J. 2006. Thermal constraints on the emplacement rate of a large intrusive complex: the Manaslu leucogranite, Nepal Himalaya. Journal of Petrology, 47, 71-95. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & MEDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06406. DOI: 10.129/2003JB002809. BENARD, F., MOUTOU, P. & PICHAVANT, M. 1985. Phase relations of tourmaline leucogranites and the significance of tourmaline in silicic magmas. Journal of Geology, 93, 271-291. BROWN, L. D., ZHAO, W. J., NELSON, D. K. g~ AL. 1996. Bright spots, structure, and magmatism in southern Tibet from INDEPTH seismic reflection profiling. Science, 274, 1688-1690. BURCHFIEL, B. C., CHEN, Z., HODGES, K. V., Liu, Y., ROYDEN, L. H., DENG, C. & XU, J. 1992. The South Tibetan Detachment System, Himalayan orogen: extension contemporaneous with and parallel to shortening in a collisional moutain belt. Geological Society of America Special Paper, 269. BURG, J. P., BRUNEL, M., GAPAIS, D., CHEN, G. M. & LIu, G. H. 1984. Deformation of leucogranites of the crystalline Main Central Sheet in southern Tibet (China). Journal of Structural Geology, 6, 535-542. COLCHEN, M., LE FORT, P. & PI~CHER,A. 1986. Annapurna-Manaslu-Ganesh Himal. Centre National de la Recherche Scientifique, Paris. COLEMAN, M. E. 1998. U-Pb constraints on Oligocene-Miocene deformation and anatexis within the central Himalaya, Marsyandi valley. American Journal of Science, 298, 553-571. COLEMAN, M. E. & HODGES, K. V. 1998. Contrasting Oligocene and Miocene thermal histories from the hanging wall and footwall of the South Tibetan Detachment in the central Himalaya from 40Ar/ 39Ar thermochronology, Marsyandi valley, central Nepal. Tectonics, 16, 726-740. COPELAND, P., HARRISON,T. M. & LE FORT, P. 1990. Age and cooling history of the Manaslu granite: implications for Himalayan tectonics. Journal of Volcanology and Geothermal Research, 44, 33-50. DEMEL, C., VIDAL, P., FERNANDEZ,A., LE FORT, P. & PEUCAT, J. J. 1987. Isotopic study of the Manaslu granite (Himalaya, Nepal): inferences on the age and source of Himalayan leucogranites. Contributions to Mineralogy and Petrology, 96, 78-92.

LEUCOGRANITES AND EXTENSIONAL FAULTS DEZES, P. J., VANNAY, J. C., STECK, A., BussY, F. & COSCA, M. 1999. Synorogenic extension: quantitative constraints on the age and displacement of the Zanskar shear zone (northwest Himalaya). Geological Society of America Bulletin, 111, 364-374. DIETRICH, V. & GANSSER, A. 1981. The leucogranites of the Bhutan Himalaya (crustal anatexis versus mantle melting). Bulletin Suisse de Miniralogie et Pitrographie, 61, 177-202. ENGLAND, P., LE FORT, P., MOLNAR, P. & P~CHER, A. 1992. Heat source for Tertiary metamorphism and anatexis in the Annapurua-Manaslu region, Central Nepal. Journal of Geophysical Research, 97, 2107-2128. FRANCE-LANORD, C., SHEPPARD, S. M. F. & LEFORT, P. 1988. Hydrogen and oxygen isotope variations in the High Himalaya peraluminous Manaslu leucogranite - Evidence for heterogeneous sedimentary source. Geochimica et Cosmochimica Acta, 52, 513-526. GODIN, L., GLEESON, T., SEARLE, M. P., ULLRICH, T. D. & PARRISH, R. R. 2006a. Locking of southward extrusion in favour of rapid crustal-scale buckling of the Greater Himalayan Sequence, Nar Valley, Central Nepal. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 269-292. GODIN, L., GRUJIC, D., LAW, R. D. & SEARLE, M. P. 2006b. Channel flow, ductile extrusion and exhumation in continental collision zones: an introduction. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 1-23. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KUNDIG, R., PAVLIS, T. & SCHMID, S. M. 1996. Ductile extrusion of the Higher Himalayan Crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. GUILLOT, S., HODGES, K. V., LE FORT, P. & P~CHER, A. 1994. New constraints on the age of the Manaslu leucogranite : evidence for episodic tectonic denudation in the central Himalayas. Geology, 22, 559-562. GUILLOT, S., LE FORT, P., P~CHER, A., Roy BARMAN,M. & APRAHAMIAN, J. 1995. Contact metamorphism and depth of emplacement of the Manaslu granite (central Nepal). Implications for Himalayan orogenesis. Tectonophysics, 241, 99-119. HARRIS, N. & MASSEY, J. 1994. Decompression and anatexis of Himalayan metapelites. Tectonics, 13, 1537-1546. HARRIS, N., VANCE, D. & AYRES, M. 2000. From sediment to granite: timescales of anatexis in the upper crust. Chemical Geology, 162, 155-167. HARRISON, T. M., GROVE, M., LOVERA, O. M. & CAa'LOS, E. J. 1998. A model for the origin of Himalayan anatexis and inverted metamorphism. Journal of Geophysical Research, 103, 2701727032.

325

HARRISON, T. M., GROVE, M., LOVERA, O. M., CATLOS, E. J., D'ANDREA, J. & MCKEEGAN, K. 1999a. The origin of Himalayan anatexis and inverted metamorphism: Models and constraints. Journal of Asian Earth Sciences, 17, 755-772. HARRISON, T. M., GROVE, M., MCKEEGAN, K., COATH, C. D., LOVERA, O. M. & LE FORT, P. 1999b. Origin and episodic emplacement of the Manaslu Intrusive Complex, Central Himalaya. Journal of Petrology, 40, 3-19. HENRY, P., LE PICHON, X. & GOFFE, B. 1997. Kinematic, thermal and petrological model of the Himalayas: constraints related to metamorphism within the underthrust Indian crust and topographic elevation. Tectonophysics, 273, 31-56. HERREN, E. 1987. Zanskar shear zone-northeastsouthwest extension within the Higher Himalayas (Ladakh, India). Geology, 15, 409-413. HODGES, K. V., HURTADO, J. M. & WHIPPLE, K. X. 2001. Southward extrusion of Tibetan crust and its effects on Himalayan tectonics. Tectonics, 20, 799-809. HUBBARD, M. S. & HARRISON,T. M. 1989.40Ar/39Ar age constraints on deformation and metamorphism in the Main Central thrust zone and Tibetan Slab, eastern Nepal, Himalaya. Tectonics, 8, 865-880. HUERTA, A. D., ROYDEN, L. H. & HODGES, K. V. 1999. The effects of accretion, erosion and radiogenic heat on the metamorphic evolution of collisional orogens. Journal of Metamorphic Geology, 17, 349-366. INGER, S. & HARRIS, N. 1992. Tectonothermal evolution of the High Himalayan crystalline sequence, Langtang Valley, northem Nepal. Journal of Metamorphic Geology, 10, 439-452. INGER, S. & HARRIS, N. 1993. Geochemical constraints on leucogranite magmatism in the Langtang valley, Nepal Himalaya. Journal of Petrology, 34, 345-368. JAMIESON, R. A., BEAUMONT, C., MEDVEDEV, S. & NGUYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research-Solid Earth, 109, B06407. DOI: 10.1029/2003JB002811. JAUPART, C. & PROVOST, A. 1985. Heat focusing, granite genesis and inverted metamorphic gradients in continental collision zones. Earth and Planetary Science Letters, 73, 385-397. JELLINEK, A. M., KERR, R. C. & GRIFFITHS, R. W. 1999. Mixing and compositional stratification produced by natural convection. 1. Experiments and their application to Earth's core and mantle. Journal of Geophysical Research-Solid Earth, 104, 7183-7201. KOHN, M. & PARKINSON, C. D. 2002. Petrologic case for Eocene slab break-off during the Indo-Asian collision. Geology, 30, 591-594. LE FORT, P. 1975. Himalayas, the collided range: present knowledge of the continental arc. American Journal of Science, 275 A, 1-44. LE FORT, P. 1981. Manaslu leucogranite: a collision signature of the Himalaya. A model for its

326

C. ANNEN & B. SCAILLET

genesis and emplacement. Journal of Geophysical Research, 86, 10545-10568. LE FORT, P., CUNEY,M., DENIEL, C., FRANCE-LANORD, C., SHEPPARD, S. M. F., UPRETI, B. N. & VIDAL, P. 1987. Crustal generation of the Himalayan leucogranites. Tectonophysics, 134, 39-57. LEJEUNE, A. M. & RICHET, P. 1995. Rheology of crystal-bearing silicate melts - an experimentalstudy at high viscosities. Journal of Geophysical Research, 100, 4215-4229. LI, S., UNSWORTH, M. J., BOOKER, J. R., WEI, W., TAN, H. & JONES, A. G. 2003. Partial melt or aqueous fluid in the mid-crust of Southern Tibet? Constraints from INDEPTH magnetotelluric data. Geophysical Journal International, 153, 289-304. MOLNAR, P., WANG-PIN, C. & PADOVANI, E. 1983. Calculated temperatures in overthrust terrains and possible combinations of heat sources for the Tertiary granites in the Greater Himalaya. Journal of Geophysical Research, 95, 4833-4856. MONTEL, J. M. 1993. A model for monazite/melt equilibrium and application to the generation of granitic magmas. Chemical Geology, 119, 127-146. MURPHY, M. A. & HARRISON, T. M. 1999. Relationship between leucogranites and the Qomolangma detachment in the Rongbuk valley, south Tibet. Geology, 27, 831-834. NELSON, K. D., ZHAO, W. J., BROWN, L. D. ET AL. 1996. Partially molten middle crust beneath southern Tibet: Synthesis of project INDEPTH results. Science, 274, 1684-1688. NOBLE, S. R. & SEARLE, M. P. 1995. Age of crustal melting and leucogranite formation from U-Pb zircon and monazite dating in the western Himalaya, Zanskar, India. Geology, 23, 1135-1138. PATINO DOUCE, A. & HARRIS, N. 1998. Experimental constraints on Himalayan anatexis. Journal of Petrology, 39, 689-710. PEARSON, O. N. & DECELLES, P. G. 2005. Structural geology and regional tectonic significance of the Ramgarh thrust, Himalayan fold and thrust belt. Tectonics, 24, TC4008, DOI: 10.1029/2003 TC001617. Pt~CHER, A. 1989. The metamorphism in the central Himalaya. Journal of Metamorphic Petrology, 7, 31-41. PINET, C. & JAUPART, C. 1987. A thermal model for the distribution in space and time of the Himalayan granites. Earth and Planetary Science Letters, 84, 87-99. ROBINSON, D. M. & PEARSON, O. N. 2006. Exhumation of the Greater Himalayan rock along the Main Central Thrust in Nepal: implications for channel flow. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 255-267. ROSCOE, R. 1952. The viscosity of suspensions of rigid spheres. British Journal of Applied Physics, 3, 267 -269. ROYDEN, L. H. & BURCHFIEL, B. C. 1985. Thinskinned N - S extension within the convergent Himalayan region: gravitational collapse of a Miocene topographic front. In: COWARD, M. P., DEWEY, J. F. & HANCOCK, P. L. (eds) Continental

Geological Society, Extensional Tectonics. London, Special Publications, 28, 611-619. SCAILLET, B., FRANCE-LANORD, C. & LE FORT, P. 1990. Badrinath-Gangotri plutons (Garhwal, India): petrological and geochemical evidence for fractionation processes in a High Himalayan leucogranite. Journal of Volcanology and Geothermal Research, 44, 163-188. SCAILLET, B., PICHAVANT, M. & ROUX, J. 1995. Experimental crystallization of leucogranite magmas. Journal of Petrology, 36, 663-705. SCAILLET, B., HOLTZ, F., PICHAVANT,M. & SCHMIDT, M. 1996. Viscosity of Himalayan leucogranites: Implications for mechanisms of granitic magma ascent. Journal of Geophysical Research, 101, 27691-27699. SCAILLET, B., HOLTZ, F. & PICHAVANT, M. 1997. Rheological properties of granitic magmas in their crystallization range. In: BOUCHEZ, J. L., HUTTON, D. & STEPHENS, W. E. (eds) Granites: From Segregation of Melt to Emplacement Fabrics, Kluwer Academic, Boston, 11-29. SEARLE, M. P. & GODIN, L. 2003. The South Tibetan Detachment and the Manaslu Leucogranite: A structural reinterpretation and restoration of the Annapurna-Manaslu Himalaya, Nepal. Journal of Geology, 111, 505-523. SEARLE, M. P. & SZULC, A. G. 2005. Channel flow and ductile extrusion in the High Himalayan slab-the Kangchenjunga-Darjeeling profile, Sikkim Himalaya. Journal of Asian Earth Science, 25, 173-185. SEARLE, M. P., PARRISH, R. R., HODGES, K. V., HURFORD, A., AYRES, M. W. & WHITEHOUSE, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya: field relations, geochemistry, age, origin and emplacement. Journal of Geology, 105, 295-317. SEARLE, M. P., NOBLE, S. R., HURFORD, A. J. & REX, D. C. 1999. Age of crustal melting, emplacement and exhumation history of the Shivling leucogranite, Garhwal Himalaya. Geological Magazine, 136, 513-525. SEARLE, M. P., SIMPSON, R. L., LAW, g. D., PARRISH, R. R. & WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal-South Tibet. Journal of the Geological Society, London, 160, 345-366. TURCOTTE, D. L. & SCHUBERT, G. 1982. Geodynamics. John Wiley, New-York. VANCE, D., AYRES, M., KELLEY, S. • HARRIS, N. 1998. The thermal response of a metamorphic belt to extension: constraints from laser Ar data on metamorphic micas. Earth and Planetary Science Letters, 162, 153-164. VISONA, D. & LOMBARDO, B. 2002. Two-mica and tourmaline leucogranites from the Everest-Makalu region (Nepal-Tibet). Himalayan leucogranite genesis by isobaric heating? Lithos, 62, 125-150. WALKER, J. D., MARTIN, M. W., BOWRING, S. A., SEARLE, M. P., WATERS, D. J. & HODGES, K. V. 1999. Metamorphism, melting, and extension: age constraints from the High Himalayan Slab of southeast Zanskar and northwest Lahaul. Journal of Geology, 107, 473-495.

4~ thermochronological constraints on the cooling and exhumation history of the South Tibetan Detachment System, Nyalam area, southern Tibet Y U W A N G a, QI LI 2 & G U O S H E N G Q U 2 G e o l o g i c L a b o r a t o r i e s C e n t e r a n d D e p a r t m e n t o f Geology, China University o f Geosciences, Beijing 100083, China (e-mail: w a n g y u 1 9 6 6 0 1 @ s o h u . c o m ) 2Institute o f Geology, China S e i s m o l o g i c a l Bureau, Beijing 100029, China Abstract: The Nyalam detachment is part of the east-west striking South Tibetan Detachment

System exposed in the Nyalam area, southern Tibet. Seventeen muscovite and biotite 4~ age spectra and three K-feldspar multidiffusion domain modelling and cooling ages are presented for metamorphic rocks, leucogranite, granite and mylonite, collected from the Nyalam detachment and surrounding areas. The majority of the 4~ results are cooling ages related to exhumation, which therefore place important constraints on formation of the Nyalam detachment and exhumation history of the region. Muscovite 4~ ages from mylonite within the normal fault system and from the footwall of the fault are 16.1-15.2 Ma. Biotite 4~ ages from the same samples are 15.6-14.8 Ma, slightly younger than the muscovite cooling ages. K-feldspar multidiffusion domain modelling suggests that samples collected from both mylonite on the fault surface and from footwall rocks underwent rapid cooling between 16.1 Ma and 11.7 Ma. Ages and cooling histories in the Nyalam detachment and Greater Himalayan metanaorphic sequence have similar characteristics and time constraints: the K-feldspar modelling indicates a sudden change in cooling rates for these regions during 15.5-14.0 Ma and c. 12 Ma, respectively. Taking the regional thermal history into account, cooling could be associated with significant northward surface movement triggered by detachment normal faulting in the Nyalam area. The Nyala~ detachment and Greater Himalayan metamorphic sequence experienced similar cooling and exhumation histories dtu-ingc. 17.0-11.7 Ma. Formation of the Nyalam detachment may have accompanied the southward extrusion of the Greater Himalaya zone along shear zones formed in response to underthrusting of the Indian plate beneath southern Tibet.

Miocene normal faulting along north-dipping detachment zones in southern Tibet is manifested as a nearly east-west striking fault system known as the South Tibetan Detachment system (STDS). The STDS is situated between the Tibetan sedimentary sequence and the Greater Himalayan metamorphic sequence (Fig. 1). The Nyalam detachment was originally named and mapped by Burchfiel et al. (1992) as a strand of the South Tibetan fault system. The structures and petrologic characteristics of the Nyalam detachment and surrounding areas have been studied by numerous workers (e.g. Burg & Chen 1984; Burg et al. 1984; Burchfiel et al. 1992; Hodges et al. 1993). Significant research effort has been invested in quantifying the timing and cooling rate of tectonic exhumation along the STDS and Greater Himalayan metamorphic zone, as well as identifying the driving mechanism of the STDS (Scharer et al. 1986; Maiuski et al. 1988; Hubbard & Harrison 1989; Copeland & Harrison 1990; Hubbard et al. 1991; Burchfiel et aL 1992; Harrison et al. 1992; Sorkhabi et al. 1996; Edwards & Harrison 1997; Searle et al. 1997, Searle 1999a, b, 2002,

2003; Murphy & Harrison 1999; Dezes et al. 1999; Walker et al. 1999; Simpson et al. 2000; Catlos et al. 2002). Although the STDS is well exposed in southern Tibet, the underlying mechanics are not understood clearly. In recent years, contrasting mechanisms have been proposed for the formation of the STDS and its relationship with the Main Central Thrust (MCT) and Greater Himalayan metamorphic sequence, such as southward extrusion, orogenic channel flow, and compression folding (Grujic et al. 1996, 2002; Carosi et al. 1998; Dezes et al. 1999; Grasemann et al. 1999; Searle 1999a, b; Hodges 2000; Beaumont et al. 2001). What controlled formation of the STDS? Was it gravitational collapse (Burchfiel et al. 1992) or synthrusting southward extrusion (Burg & Chen 1984; Burchfiel & Royden 1985; Burg 1994)? The formation mechanism of the STDS, combined with dynamic constraints on its evolution, remain some of the key tectonic problems unresolved in the region. Along the east-west trending normal fault system of southern Tibet, in the Yadong area, a U - P b monazite age of 10-11 Ma for leucogranite was suggested as constraining the time of movement

From: LAW,R. D., SEARLE,M. P. & GODIN,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 327-354. 0305-8719/06l$15.00 9 The Geological Society of London 2006.

328

Y. WANG E T A L .

Fig. 1. Simplified tectonic diagram (a) and regional geology of the central Himalaya and southern Tibet (b) (simplifiedfrom Burg et al. 1984 and Burchfieiet al. 1992). Abbreviations:K, Kangmar dome; LK, Largikangri dome; M, Mabja dome. (9, Girong detachment; | Nyalam detachment; | Qomolangma (Everest) detachment; | Dingjie detachment; | Yadong detachment. ITSZ, Indus-Tsangpo suture zone; MBT, Main Boundary thrust; MCT, Main Central thrust; STDS, Southern Tibetan Detachment System.

along the Yadong detachment (Wu e t al. 1998), Sch~er e t al. (1986) and Burchfiel et al. (1992) suggested an initial age of c. 16.8 Ma for the formation of the STDS in the Nyalam area, based on monazite collected from leucogranite. Approximately 30 km west of Nyalam at Shisha Pangma, and also in the Everest area to the east, Searle e t al. (1997, 2002) considered that normal faults post-dated 17.3 __+0.2Ma, and were active at 17 Ma, respectively. Edwards & Harrison (1997) and Murphy & Harrison (1999) proposed that the

STDS developed later than c. 20 Ma, as constrained by intrusion of leucogranite. The first clear statement that the Nyalam detachment was active at c. 17-15 Ma was by Dougherty e t al. (1998). With the exception of the above dates, there are no direct isotopic ages available to constrain the age of formation of the STDS on the Chinese side of the Himalayan belt. In this study, we address the timing of STDS development and present relevant cooling histories. Formation of the STDS clearly affected the thermal history in southern Tibet,

SOUTH TIBETAN DETACHMENT SYSTEM especially the Himalayan orogenic belt. Maluski et al. (1988) obtained leucogranite muscovite and biotite 4~ cooling ages of 14.4Ma and 16.3 Ma, respectively, in the Nyalam area (sample locality not given). Both at Shisha Pangma and in the Everest area, rapid cooling occurred during 17-14 Ma (Searle 1999a, b; Searle et al. 1997; 2002, 2003). This study aims to constrain the timing of extension and cooling history of the Nyalam detachment using 4~ dating of structurally controlled samples collected along the fault surface and footwall of the Nyalam detachment. Cooling histories for these rocks were obtained by 4~ stepheating of muscovite and biotite, and multidiffusion domain modelling on K-feldspar.

Geological setting and structural framework of the Nyalam detachment Tectonic setting of the Nyalam detachment and its features In the Nyalam area of southern Tibet and adjoining Nepal, from south to north, three main geological units can be recognized: the Main Central Thrust (MCT) system, the Greater Himalaya zone, and the South Tibetan Detachment System (STDS) (Fig. 1). The MCT is a north-dipping, low-angle thrust fault separating the Lesser Himalaya and Greater Himalaya. The STDS is a N-NNE-dipping zone of normal-sense faulting between the Greater Himalaya and Tibetan sedimentary sequence (Fig. 2). Several distinct areas were selected for sample collections along the China-Nepal highway, between the STDS and the MCT (Figs 2 & 3). The Nyalam detachment, a part of the STDS, is located 30 km north of Nyalam (Fig. 2). The footwall of the Nyalam detachment fault is composed of metasedimentary rocks, granite and leucogranite of the Greater Himalaya zone. The most abundant metasedimentary rocks are dark grey quartz-biotite schist and gneiss, quartzite, and psammitic schist. The common mineral assemblage is garnet + biotite § microcline -t- plagioclase -t- silli manite, indicating upper amphibolite facies metamorphic conditions (Burchfiel et aL 1992). More detailed petrologic features of these rocks were described by Hodges et al. (1993).

Deformational features of the Nyalam detachment Mylonite was formed and leucogranite was intruded (Figs 4 & 5a) along the detachment fault surface and its footwall. Asymmetric folds in quartzite and schist below the Nyalam detachment indicate

329

a normal (top-down-to-the-north) sense of shear (Fig. 5b). Associated S-C fabrics, north- and NNE-trending stretching lineations, and NNEdipping foliation are also exposed. The most spectacular shear-sense indicators and S-C mylonitic fabrics were developed below the fault surface in leucogranites (Fig. 5c). Muscovite in these rocks is orientated parallel to the mylonitic lineation (Fig. 5d). The geometry of abundant mica 'fish' and asymmetric augen structures universally indicate top-down-to-the-north shear, and northdipping C-surfaces always indicate a normal sense of motion (Fig. 5e-j). Other shear-sense indicators include outcrop-scale shear bands, asymmetric plagioclase augen, boudinage, and asymmetric K-feldspar clasts with recrystallized tails. Many microstructures in thin sections confirm this north-dipping normal shear sense, and formed at medium-temperature deformation conditions (c. 400-350~ (using microstructural criteria of Passchier & Simpson (1986) and Passchier & Trouw (1996)). Deformation and metamorphism in the footwall and hanging wall of the Nyalam detachment are very different. The hanging wall of the Nyalam detachment fault consists of unmetamorphosed Cambrian-Permian Tibetan sedimentary strata. The Nyalam detachment locally cuts early southvergent folds and thrust sheets in the hanging wall and is parallel to mylonitized units on the detachment surface and in the footwall, to the south of the Yalai thrust fault (Figs 3 & 4), indicating that normal faulting formed later than the southverging thrusting system in the region.

Deformational features in the footwall of the Nyalam detachment In the footwall of the Nyalam detachment, asymmetric boudin trains and folds indicating a normal sense of shear occur within foliated metasedimentary rocks near the detachment fault (Fig. 5k-l), but disappear at Jiangdong to the south (position of Jiangdong is shown in Fig. 2). These structures are indicative of top-to-the-north shear orientated at N8-20~ parallel to the stretching lineation. Normal-sense shearing is limited to the detachment and its footwall area, and dies out downwards within the upper part of the Greater Himalaya zone. Top-down-to-the-north shearing was therefore limited to the uppermost (northern) part of the Greater Himalayan metamorphic zone. An undeformed, unmetamorphosed granitic intrusion crops out in the footwall of the Nyalam detachment fault (Fig. 3). No direct intrusive relationship between detachment surface and undeformed granite has been found, but the intrusion

330

Y. WANG E T A L .

Greater Himalayan

NL-2S- \

Metamorphic Sequence I 7.5kin t

Metamorphic Sequence

~'.~NL-29 g .~

~

Zhangmu

28 ~

(a)

N E20 ~

(b)

NEso

~. Nyalam

Z hangm u

~

~

.

0 I

~

..- -.,

5 I

10kin 1

Ru.i i

N L.-2~.~

N I.- 53

"

myionite

\~

szDs

granite

schist

gneiss

.

-

sedimentary rocks

Fig. 2. Geological sketch map of the Nyalam-Zhangmu area: (a) sampled sites; (b) cross-section through the Nyalam detachment and Greater Himalayan metamorphic sequence. Abbreviations: MCT, Main Central thrust; STDS, Southern Tibetan Detachment System. Site of Figure 3 is indicated by box outline in (a). Cross-section in (b) extends from Ruji through Zhangmu to footwall of MCT; locations of samples NL-25 and NL-53 are indicated.

cuts foliation with top-to-the-north shear sense in the footwall of the detachment fault. A suite of gneiss, upper-greenschist-facies metamorphic rocks, migmatite, mica schist, and marble make up the south-central part of the Greater Himalaya zone in the region. Near Nyalam, the

Greater Himalayan metamorphic sequence consists of metasedimentary units intruded by leucogranite and granite (Burchfiel et al. 1992). From this area south towards the China/Nepal boundary, gneiss and sillimanite-bearing metamorphic rocks in the Greater Himalaya zone are exposed. The most

SOUTH TIBETAN DETACHMENT SYSTEM

Carboniferous-Permian Ordovician-Devonian Cambrian (?)

.a_25~ dipofbedding .)1.3~176foliation and lineation

N

NL-15 sampled site

injection complex

I

tin-deformed granite

.a..~ / it

detachment fault

// / /

thrust fault

[~ ~/

g . . . . . , . . . . 'fault fault and inferred fault

valle floors

~ \ / ~

-

e.~ / J']" ~ t

/ / /

[I

"-" /

anticline /.."

331

\

I ~

\\

C-P ~'~-~///

\ \

~""~"J/~

32~

-

5 • 106 km, is an arid plateau with low relief and low erosion rates, and forms the largest area of high elevation (average elevation of 5023 m) and thick crust ( 6 5 - 8 0 km) on the planet (Fielding et al. 1994). The plateau is bordered to the south by the Himalayan range, which has a similar average elevation, but much greater topographic relief, ranging from 2000 to 8850 m, intense fluvial and glacial erosion, and high erosion rates (Duncan et al. 2003). The horizontal gradients in lithostatic pressure between the Tibetan plateau and the Indian shield south of the Himalaya, with its normal crustal thickness (c. 3 5 - 4 0 km) and low elevation ( < 1.5 km), is the driving force behind deformation involving crustal flow away from the plateau. Two major types of crustal flow have been proposed for Tibet: firstly, lower crustal flow

in eastern Tibet extruding east and SE around the eastern Himalayan syntaxis; and secondly, middle crustal flow in a layer or channel extruding southward from beneath the southern part of the plateau to the Greater Himalayan range (Fig. 1). The concept of planar channel flow (combined 'Couette flow' (simple shear) and 'Poiseuille flow' (also known as the 'pipe-flow effect') of the weak lower crust, which removes crust from beneath mountains and levels out topography, was first proposed by Bird (1991). The crustal structure of the eastem part of the Tibetan plateau has been interpreted in terms of lateral flow of a ductile layer of lower crust by several authors (Royden et al. 1997; Clark & Royden 2000). This appears to be supported by GPS observations that show the ground surface of north Tibet rotating around the eastern syntaxis as far as Yunnan and Burma

From: LAW, R. D., SEARLE,M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 355-378. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

356

M.P. SEARLE ET AL.

Fig. 1. Sketch map of the Tibetan plateau region comprisingthe Songpan Ganze, Qiangtang and Lhasa blocks accreted to the southern margin of Asia. Major suture zones and strike-slip faults are also shown. The Himalaya (shaded) forms the southern margin of the Tibetan plateau. The Greater Himalaya is the shaded and stippled area between the Main Central Thrust (MCT) and the South Tibetan Detachment(STD); SS, Shyok suture; ISZ, Indus-Tsangposuture zone.

(Wang et al. 2001). Velocity vectors show that the south Tibetan crust and Himalaya (south of 32~ are moving along an azimuth of 020~ with respect to stable Asia. North of this, the velocity vectors swing around to ENE in northern Tibet, east along the eastern margin of the plateau and continue to the south in Yunnan and northern Burma (Wang et al. 2001). This eastward flow of material is restricted to northern Tibet (Qiangtang and Kunlun terranes), north of the Jiale fault and south of the Kunlun fault (Fig. 1). Northern Tibet has a thin, hot lithosphere, high Poissons' ratio (ratio of latitudinal to longitudinal strain), a seismic anisotropy, possibly developed by lateral flow in the mantle, a probable granulitic lower crust and abundant Tertiary-Quaternary volcanic activity (Owens & Zandt 1997; Kosarev et al. 1999; Hacker et al. 2000). Southern Tibet has a thicker, cooler lithosphere consistent with underthrusting of crust and mantle of Indian plate affinity north as far as the latitude of the Bangong suture at c. 32~ (Nelson et al. 1996; Owens & Zandt 1997). The lower crustal flow of the Qiangtang and Songpan Ganze terranes of northern Tibet contrasts

with the channel flow model for the Greater Himalaya-south Tibet crust, in which a layer of hot, ductile deforming and partially molten middle crust was extruded south from beneath the southern part of Tibet (e.g. Grujic et al. 1996, 2002; Beaumont et al. 2001, 2004; Searle et al. 2003; Jamieson et al. 2004). Here, the mid-crustal layer of Indian plate rocks is represented by the Greater Himalayan Sequence (GHS) which is exposed in areas that form most of the highest topography, and the deepest exhumed rocks, along the Himalayan arc. The lower crust, comprising Archean Indian shield granulites, has been underthrust to the north, and is relatively rigid and melt-free. Since the crust beneath the Tibetan plateau is up to 8 0 k m thick (c. 2 0 - 2 2 k b a r pressure, assuming average crustal densities) it is possible that the lower crust has transformed into eclogite beneath the Lhasa block of southern Tibet. However, whereas granulites are dry, strong and capable of supporting thick crust like Tibet, eclogites may be hydrous, and have a density similar to, or possibly even greater than, that of the upper mantle. Ultra-high temperature and near

EVEREST STRUCTURE, RESTORATION AND EVOLUTION ultra-high pressure xenoliths in young alkali basalts in central Tibet include felsic granulites and both felsic and mafic eclogites, interpreted as tapping deep levels of the Tibetan crust (Hacker et al. 2000; Ducea et al. 2003). This paper synthesizes data from the Everest transect, Tibet-Nepal in the context of the channel flow model. For the subsurface structure, we rely on the various seismic profiles from project INDEPTH (e.g. Nelson et al. 1996; Alsdorf et al. 1998; Hauck et aL 1998). For the surface geological mapping we use our data (e.g. Searle 1999a, b, 2003; Searle et al. 2002, 2003) as well as that of Lombardo et al. (1993), Carosi et al. (1998, 1999a, b), and for the microstructural interpretation we use the data from Law et al. (2004). We attempt a restoration of the entire Nepal-south Tibet GHS and use this restored section as a template to determine the protoliths of GHS rocks, the depth of melting of the Everest leucogranites, and the thrust and normal fault trajectories. Finally, we discuss the role of the South Tibetan Detachment (STD) normal faults as a passive roof fault during channel extrusion, and the role of the Main Central Thrust (MCT) in the extrusion and exhumation of the GHS.

Greater Himalayan Sequence The GHS, as exposed along the High Himalaya, is composed of a mid-crustal layer of high-grade metamorphic rocks and migmatites with sheets of crustal melt leucogranites prominent along the higher structural levels. The upper contact of the GHS is a low-angle normal fault system, the STD, which generally dips north (5-20~ beneath the Tibetan plateau (Burg 1983; Burg et al. 1984; Burchfiel et al. 1992; Searle et al. 1997). The lower boundary of the GHS is a high strain ductile shear zone, the MCT, which is coincident with a zone of inverted metamorphic isograds from kyanite grade down to biotite-chlorite grade (Hubbard 1996; Stephenson et al. 2000, 2001). The MCT also dips at low angles to the north and places Proterozoic and Palaeozoic rocks, metamorphosed during the Himalayan orogeny, south over Lesser Himalayan thrust sheets, which were largely unaffected by Himalayan metamorphism (Gansser 1964, 1983; Hodges 2000; Searle & Godin 2003). The large-scale structure and pressuretemperature (P-T) constraints across the GHS have led to numerous thermal-mechanical models (e.g. Searle & Rex 1989; Hodges et al. 1992; Harrison et al. 1998; Searle et al. 1999b). Several recent papers interpreted the structural and thermal data from the GHS in terms of a channel flow model, whereby the middle crust was extruded

357

southwards between coeval thrust- and normalsense shear zones (Grujic et al. 1996, 2002; Beaumont et al. 2001, 2004; Searle et al. 2002, 2003; Searle & Szulc 2005). Although these models differ in some aspects along the strike of the mountain range, several first-order field- and laboratory-based data sets form the basis for meaningful models (see next section). Figure 2a and b show a map of the Nepal-South Tibet Himalaya, and a cross-section across the STD in the Everest region. Figure 2c is a restoration of the profile in Figure 2b using pressure-depth constraints from thermobarometry of footwall gneisses and assuming a constant dip of the STD (Searle et al. 2002, 2003). These authors suggested horizontal transport or flow of at least 100 km (possibly as much as 200km) of GHS rocks along the footwall of a passive roof fault, the STD, during ductile channel flow during the Miocene. The thermal structure of the Greater Himalaya has been interpreted with the help of petrologic and thermobarometric data. P - T profiles across the GHS generally show an inversion of temperature and pressure across the narrow MCT zone (e.g. Hubbard 1996; Searle et al. 1999b, 2003; Stephenson et al. 2000, 2001). Structurally above the kyanite-grade rocks, the Everest section shows a 50km horizontal width of sillimanite-grade gneisses (Hubbard 1989, 1996; Simpson 2002; Searle et al. 2003). Temperatures remain high for 5 0 k m along a north-south profile from the footwall of the STD at Rongbuk south as far as Karikhola, south of Lukla. Pressures are highest (7-8 kbar) in the south above the MCT zone and drop off to the north to 4 - 5 kbar (Searle et al. 2003, fig. 8). Pressures of Everest greenschist rocks above the Lhotse detachment are poorly constrained, but they must be relatively low ( < 3 kbar) because of the low metamorphic grade (upper greenschist facies). Recent discoveries of staurolite-grade schists from the south face of Lhotse (Jessup et al. 2004) suggest that metamorphism decreases up-structural section along the top of the GHS slab in a structural style similar to that observed in Zanskar (Searle et al. 1992, 1999b). Several authors have suggested that metamorphic isograds were affected by crustal-scale post-peak metamorphic folding and shearing (Searle & Rex 1989; Hubbard 1996; Walker et al. 1999; Stephenson et al. 2000, 2001). At least in the NW Himalaya, it can be demonstrated that isograds were inverted by south-vergent folding and associated simple shear along the MCT zone at the base of the extruding layer. Metamorphic isograds at the top of the GHS extruding layer are right-way-up and affected by post-metamorphic normal-sense shearing and flattening (Searle & Rex 1989; Dezes et al. 1999; Walker et al. 1999;

358

M.P. SEARLE E T A L .

Fig. 2. (a) Geological map of the central Himalaya in Nepal, Sikkim, Bhutan and south Tibet. K, Kathmandu; MBT, Main Boundary Thrust; MCT, Main Central Thrust, (b) Geological section across the Everest-Rongbuk profile, marked a and b on (a) (after Searle 2003); vertical and horizontal scales are equal. (c) Restoration of the Everest profile along the line of section marked on (a) after Searle et al. (2002, 2003). The Qomolangma Detachment (QD) and Lhotse Detachment (LD) are part of the South Tibetan Detachment system of low-angle normal faults. Crosses mark the restored positions and depths of three samples from Kala Patar and Pumori, 2 km SW of Everest, which record pressures of 4.0, 4.1 and 4.9 kbar. Using a mean 10 ~ dip of the STD passive roof fault, the relative southward displacement of footwall Greater Himalayan sequence (GHS) rocks is 90-108 km. The true dip of the STD on the north face of Everest is between 3 and 5 ~ N; at this angle the relative displacement along the STD would be 180-216 km. The approximate depth and original position of Everest leucogranites prior to emplacement beneath the passive roof fault is also shown.

EVEREST STRUCTURE, RESTORATION AND EVOLUTION Jessup et al. 2004). In many parts of the Garhwal and Nepal Himalaya, late-stage brittle normal faulting has cut out the flattened and sheared right-wayup isograds at the top of the GHS layer, and the STD fault places unmetamorphosed Palaeozoic sediments directly on top of sillimanite-grade gneisses and/or leucogranites (e.g. Shisha Pangma (Searle et al. 1997), Garhwal (Searle et al. 1999a), Rongbuk valley (Hodges et al. 1998; Murphy & Harrison 1999; Searle 1999a; Searle et al. 2003) and Annapurna (Godin et al. 2001)). Uniform P - T conditions of the sillimanite-grade gneisses along the Everest transect between the MCT zone at the base and the Lhotse detachment at the top, together with the lack of major structural discontinuities (except the Khumbu thrust; Searle 1999a, b) suggest that the GHS acted as a relatively homogeneous high-temperature, partially molten slab during the early Miocene. This thermal structure supports the channel flow model (Fig. 3), in which the ductile deforming channel is flowing south, bounded by major crustal-scale, low-angle ductile shear zones both below (thrust-related) and above (normal sense of shear). The timing of motion along the MCT zone along the base, and the STD zone along the top of the channel, has been extensively reviewed (Hodges et al. 1992; Hodges 2000; Searle et al. 2003; Searle & Godin 2003; Godin et al. 2006a). Both ductile shear zones were active during the early Miocene (23-15 Ma) and hornblende and mica ages are consistently older than 16-14 Ma, providing a minimum time constraint on when both shear zones were exhumed from the ductile into the brittle regime. It is possible that both shear zones are still active (e.g. Hodges et al. 2001). Evidence for this is the major difference in geomorphology and relief across both the bounding shear zones, evidence for young and recent motion along the lower MCT thrusts coincident with knick-points in the rivers, and the fact that the highest topography occurs largely across the GHS.

with partial melting at c. 22-20 Ma (Noble & Searle 1995). The resulting 5 - 1 0 km thickness of migmatites, with abundant leucogranite sills, formed by dehydration melting of muscovite in the underlying metapelites at 4 - 7 kbar and 650750~ (Searle et al. 1992, 2003). The fight-way-up isograds along the footwall of the ZSZ were mapped, linking with the inverted isograds along the MCT hanging wall via a large-scale, SWvergent, NW-plunging recumbent fold in western Zanskar (Searle & Rex 1989; Searle et al. 1992, 1999b; Stephenson et al. 2001). These data are consistent with extrusion of the GHS by channel flow and imply at least 60-100 km of motion occurring along both the bounding shear zones, the MCT and the ZSZ. In the eastern Himalaya, a similar model was developed for the Bhutan Himalaya, where Grujic et al. (1996, 2002) proposed that the GHS formed the core of a low-viscosity channel extending under the Tibetan plateau. Lateral lithostatic pressure gradients produced a 'pipe-flow' effect with the highest velocities in the centre of the channel. Beaumont et al. (2001, 2004) presented a quantitative model in which this crustal channel was coupled to surface denudation at the High Himalayan front. In contrast, Grasemann et al. (1999) and Vannay & Grasemann (2001) proposed a general shear model due to the requirements of strain compatibility, whereby the centre of the channel extruded by pure shear flow, rather than the heterogeneous simple shear model invoked by Grujic et al. (1996). We propose the following points as fundamental boundary conditions for the development of an orogenic channel flow model in the Himalaya (Fig. 3). 1.

2.

Himalayan channel flow model In the western Himalaya, several studies combining structural mapping with thermobarometry show that metamorphic isograds were folded after peak metamorphism, resulting in a right-way-up metamorphic sequence along the top of the GHS beneath the Zanskar shear zone (ZSZ), and an inverted metamorphic sequence along the MCT zone at the base of the slab (Searle & Rex 1989; Searle et al. 1992, 1999b; Walker et al. 1999; Dezes et al. 1999; Stephenson et al. 2000, 2001; Robyr et al. 2002; Vannay et al. 2004). Peak sillimanite + K-feldspar-grade metamorphism was concomitant

359

3.

4.

Channel flow is driven by the topographic and crustal thickness variation between Tibet (c. 5 km elevation; 7 0 - 8 0 k m thick crust) and the Indian foreland ( < 1 km elevation; c. 35 km thick crust). Presence of a mid-crustal layer, between c. 15 and 4 0 k m depth, composed of ductilely deforming rocks with in situ partial melting (Zhao et al. 1993; Nelson et al. 1996). A zone of inverted, telescoped and flattened metamorphic isograds and isotherms is present along the basal contact of the channel (Main Central Thrust zone). A ductile shear zone within kyanite-grade rocks propagates down-section to a later brittle thrust with time (Stephenson et al. 2000, 2001). A zone of right-way-up, telescoped and flattened isograds and isotherms is present along the top of the extruding channel (STD). A ductile, normal-sense shear zone propagates upwards to a later brittle normal fault with

360

M . P . S E A R L E ETAL.

~"o

r

~2 _

~o

~ - ~ ~92

0

0

9

~

~

~00

~

.~~o _~ ~ ' ~ . ~ ~ ..~ ~

,,,~.~

~ = ~

~

oo'~

o

c~ .~

~s

~

~f

o

o

~

~

-~

~.~

_~ ~ ~

.~

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

5.

6.

7.

time (Searle et al. 1992, 1997, 1999a, 2003; Searle & Godin 2003). The two shear zones bounding the extruding channel above (STD) and below (MCT) must be structurally linked and must move synchronously (Searle & Rex 1989; Hodges et al. 1992, 1993, 1996). Presence of partial melts (migmatites) or leucogranite magmas lubricating the flow within the extruding layer, but not cutting across the bounding brittle fault zones (Searle et al. 1992, 1999a, 2003). Coupling between channel flow and surface erosion. Erosion is highest where the channel reaches the margin of the plateau in the High Himalaya (Hodges et al. 2001).

Figure 3 shows our interpretation of the channel flow model, based on geological (Searle 1999a, b; Searle et al. 2003; Law et al. 2004; Jessup et al. 2006) and geophysical (Zhao et al. 1993; Nelson et al. 1996; Hauck et al. 1998) data from the Nepal and south Tibet Himalaya. The shaded part of the middle crust in Figure 3 represents the channel of high-grade gneisses, migmatites and leucogranites that, during the Miocene, acted as a ductile layer sandwiched between the brittle deforming upper crust and a rigid, anhydrous, granulite-facies lower crust.

Deep crustal structure from the INDEPTH profile The subsurface structural interpretation beneath south Tibet is based entirely on deep seismic reflection profiling from project INDEPTH (Zhao et al. 1993; Nelson et al. 1996; Brown et al. 1996), combined with broadband earthquake and magnetotelluric data (Kind et al. 1996; Wei et al. 2001). The crustal structure of southern Tibet shows several reflectors, which have been successfully matched to the major faults of the Himalaya to the south, notably the STD and the Main Himalayan Thrust (MHT; Hauck et al. 1998; Alsdorf et al. 1998). The MHT is the basal detachment, which dips at about 10 ~ north and progressively ramps up-section to the south, with both the MCT and Main Boundary Thrust (MBT) splaying off it at depth. Magnetotelluric data (Wei et al. 2001) suggest that an electrically conductive layer at 15-20 km depth corresponding to zones of seismic attenuation, may reflect partial melts and/or aqueous fluids in the middle crust of southern Tibet (Nelson et al. 1996; Alsdorf & Nelson 1999). T h e 'bright spots' imaged on the INDEPTH profile beneath south Tibet have been interpreted by

361

some as leucogranitic melts forming today, at a similar structural horizon and depth to the High Himalayan leucogranites which crystallized during the Miocene and were extruded southward to their present position along the High Himalaya (Searle et al. 1993, 1997, 2003; Searle 1999a, b; Gaillard et al. 2004). It has further been suggested that the ages of crustal melt leucogranites in the footwall of the STD decrease to the north (Wu et al. 1998) and that the Greater Himalayan crystalline rocks, bounded by the STD above and the MCT below, have been effectively extruded southwards from beneath the middle crust of south Tibet (Grujic et al. 1996, 2002; Searle 1999a; Searle et al. 2003; Searle & Szulc 2005). Several factors such as radiogenic heating, crustal shortening and concomitant thickening could have induced the high temperatures at these mid-crustal levels beneath southern Tibet. Francheteau et al. (1984) measured unusually high heat flow in the lake sediments beneath Yamdrock Tso SE of Lhasa (Fig. 1), which they attributed to the presence of a magma body at 1 0 - 2 0 k i n depth. Satellite data reveal a magnetic low over Tibet indicating hot crust, and the deeper levels of the high-conductivity zone have been interpreted as a zone of in situ partial melting (Alsdorf & Nelson 1999). Seismological data show that relatively fast (cool) upper mantle extends from the Himalaya northwards to roughly the centre of the Tibetan plateau around the latitude of the Bangong suture (Jin et al. 1994; Owens & Zandt 1997; Kosarev et al. 1999). This has been interpreted as the northern limits of cold Indian lithosphere underthrust beneath the plateau.

Crustal structure of the Everest transect The surface structure of the GHS in the Everest transect is shown in Figure 4, together with U - P b ages of metamorphic rocks and leucogranites (Simpson et al. 2000; Viskupic & Hodges 2001; Searle et al. 2003) and Th-Pb ages of Murphy & Harrison (1999) from Rongbuk, and Catlos et al. (2002) from the MCT zone. The structure of the profile has been presented previously (Searle 1999a, b, 2003; Searle et al. 2003; Law et al. 2004; see also Jessup et al. 2006, fig. 3) and only the main points pertinent to the discussion on channel flow need be repeated here. The upper part of the GHS slab in the Everest region is marked by two low-angle normal faults, which merge towards the north along the Rongbuk and Kharta valleys into one major shear zone (Searle et al. 2002, 2003). The structurally lower Lhotse detachment is a ductile shear zone that separates sillimanite-grade gneisses with abundant leucogranite sills beneath from lower amphibolite

362

M.P.

SEARLE

ET AL.

o

o

o

~,eo

o,-~

r,.) ,~

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

363

C

364

M.P. SEARLE E T AL.

(Jessup et al. 2004)-greenschist-facies pelitic and calcareous rocks (Everest series) with few, if any, leucogranites above. The structurally higher normal fault, the Qomolangma detachment, separates Everest Series metapelites beneath from unmetamorphosed Ordovician sedimentary rocks above. Beneath the Lhotse detachment a coherent leucogranite sill, up to 3.5 km thick, composed of tourmaline + biotite + muscovite + garnet leucogranite, extends at least from the Makalu massif westwards to Cho Oyu across the lower edifice of Mount Everest and the Nuptse-Lhotse massif (Searle 1999a, b, 2003). Numerous leucogranite sills extend from the Everest massif north for at least 57 km along the Kharta and Rongbuk valleys as parallel sheets within the gneisses. Early leucogranite sills are parallel to the ductile fabrics within the gneisses, whilst a few later dykes crosscut the fabric, but are themselves truncated by the brittle Qomolangma detachment above (Murphy & Harrison 1999; Searle et al. 2002, 2003). Most of the high peaks in the Khumbu Himalaya (including the base of Cho Oyu, Gyachung Kang, Nuptse, Everest, and most of the peaks Chomolonzo, Makalu, Baruntse, Ama Dablam, Kangteiga, Tramserku) are composed of flat-lying leucogranite sills within the high-grade sillimanite gneisses (Searle, 1999a, b; Weinberg & Searle 1999; Visona & Lombardo 2002). Some sills are very narrow (< 1 m), whilst others reach 1-2 km in thickness. A few interconnecting dykes link the foliation-parallel sills, but the rarity of discordant dykes implies that magma transport was dominantly along the foliation planes. In the EverestLhotse-Nuptse massif, the leucogranites lie above the north-dipping Khumbu Thrust (Searle 1999a, b), well exposed along the south face of Nuptse and Lhotse, which has transported these rocks south over sillimanite gneisses that contain very few leucogranites. Along the Everest-Lukla transect in Nepal, sillimanite remains the stable aluminium silicate over the 45km horizontal width (and > 1 5 - 2 0 k m structural depth) of the GHS, indicating that the extruding slab of middle crust had a high-temperature core during the early Miocene. Although pelites are dominant throughout this section, there are numerous horizons of calc-silicates and K-feldspar augen gneisses throughout the whole sequence down as far as the MCT zone. We apply the original definition of the MCT, and map it as the main zone of high strain separating rocks metamorphosed during the Tertiary above (Phaplu augen gneiss and structurally higher gneisses), from underlying rocks largely unmetamorphosed during the Himalayan orogeny (Lesser Himalaya). This corresponds to the lower MCT-1 position of Arita (1983) and not to the structurally higher 'MCT' (north of Karikhola, south of

Lukla; Catlos et al. 2002). Thus, rocks previously described as 'upper Lesser Himalayan crystallines' (e.g. Catlos et al. 2002; Vannay et al. 2004) at kyanite or sillimanite grade that have P - T conditions similar to rocks higher up the section throughout the GHS slab, are here considered part of the GHS, above the MCT. The entire base of the GHS is a large-scale shear zone, corresponding to the zone of inverted metamorphic isograds, but the precise distribution of strain has yet to be determined by a detailed kinematic investigation.

Restoration of the Everest Himalaya Figure 5 is a schematic restoration of the Everest transect through the GHS, showing the trajectories of a number of the major faults. The restoration shows the position of the two major low-angle normal faults that cut through Everest itself: the upper Qomolangma detachment which flattens out along the top of the Ordovician 'Yellow Band'; and the lower Lhotse detachment, which bounds most of the large leucogranite sills (Searle 1999a, b; Searle et al. 2003). The Everest granites must have been sourced within the Proterozoic black shales that formed the protolith of the sillimanite gneisses comprising Unit 1 in the GHS, also called the Barun gneiss (Pognante & Benna 1993; Lombardo et al. 1993). We discard the previous 'stratigraphic' nomenclature of the GHS (e.g. Le Fort 1981; Colchen et al. 1986) because we believe that it oversimplifies the complex metamorphic lithologies and structure of the GHS (Searle & Godin 2003). For example, there are numerous horizons of augen gneiss (previously called Formation III) throughout the GHS from the base of the slab (Ulleri and Phaplu augen gneisses) up to the highest levels beneath the leucogranites (e.g. Pumori augen gneiss; Searle 2003). Likewise, calc-silicate bands (previously Formation 11) and less common amphibolites are present at several levels throughout the GHS sequence. The Namche K-feldspar orthogneisses, together with similar orthogneisses higher up the slab near Pumori, were intruded into the sillimanite gneisses and the metamorphosed Cambrian-Ordovician limestones-calc-silicates. The orthogneisses contain biotite, sillimanite, cordierite, quartz, plagioclase and K-feldspar. Monazite-xenotime thermochronology suggests that the original protolith crystallized at, or before, 466 Ma; metamorphism began at, or before, 28.37 Ma, and anatectic melting occurred at 25.4-24.75 Ma (Viskupic & Hodges 2001). The Namche orthogneisses are therefore interpreted as part of the suite of late Cambrian-Ordovician granitoids that intruded the Indian plate (Miller e t al. 2001),

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

365

o.=

=~,

. ".~ ~

9,--~

~

&o

~

o

N~

_~ ~ ~ . ~ o ;-,~

~-"o

9

~ . ~

9

~ . = ~ ~.~ ~ n9 .u ~ ' ~ -uo=Om~=

~= ~ u

. ~ 6 ~

~

366

M.P. SEARLE E T A L .

and have subsequently been metamorphosed during the Oligocene Himalayan event, resulting eventually in Miocene partial melting. Farther south within the GHS, another prominent level of K-feldspar augen gneisses crops out near the village of Phaplu (Fig. 4). These gneisses contain K-feldspar, quartz, plagioclase, biotite and hornblende, with a metamorphic assemblage including muscovite, biotite, garnet and sillimanite. A T h - P b monazite age of 1219 ___ 9 Ma is interpreted as the age of the protolith (Catlos et al. 2002), significantly older than the protolith ages of K-feldspar augen gneisses higher in the section. Restoration of the section strongly suggests that the protolith of the sillimanite-grade metapelites lying structurally above the Phaplu augen gneiss was the Proterozoic Haimanta Group shales and greywackes (Fig. 5). A final requirement from the restored section through the GHS is that the footwall and hanging wall rocks across the MCT must match up. The GHS has a sedimentary provenance, dominantly of late Proterozoic age, based on detrital zircon ages of 0.8-1.0 Ga, whereas Lesser Himalayan zircons have older 1.87-2.60 Ga ages (Parrish & Hodges 1996). However, there is no evidence of a Palaeozoic suture and no evidence of any pre-Himalayan displacement on the MCT as required by the model of DeCelles et al. (2000). The MCT is a thrust fault, which, upon restoration, must match footwall-hanging wall cut-off points. The difference in Nd isotope characteristics can be explained by the stratigraphic level of protolith rocks either side of the MCT seen on the restoration. The Greater Himalayan rocks above the MCT have end values of -- 16 and are clearly Proterozoic shales, not basement rocks, whereas the Lesser Himalayan rocks beneath the MCT have an average eNa value of --21.5, derived from Indian basement (Robinson et aL 2003). There is, however, no need to invoke a separate terrane for the GHS as in the Robinson et al. (2003) model; it is merely a higher stratigraphic level exposed in the restored GHS than the more southerly Lesser Himalaya.

STD as a passive roof fault to the extruding channel The upper boundary of the GHS channel is marked by the low-angle normal faults of the South Tibetan Detachment (STD). This detachment is an orogenscale fault present along 2000 krn of the High Himalaya from the Zanskar region of NW India to Arunachal Pradesh in the east. It places the unmetamorphosed or anchizone sedimentary rocks of the Tethyan zone above the high-grade metamorphic

rocks, migmatites and leucogranites of the Greater Himalayan Series. The STD zone includes a structurally lower ductile shear zone, the Lhotse detachment, and a higher-level brittle fault, the Qomolangma detachment. In the Everest transect massive leucogranite sills occur below the structurally lower Lhotse detachment (Searle 1999a, b, 2003). The two shear zones merge towards the north to form one very large high-strain zone with leucogranite sills and ductile shear decreasing structurally upwards towards the Tethyan sedimentary rocks (Fig. 6a). Along the Rongbuk and Kharta valleys, thin leucogranite sills are also present up into the lower-grade rocks below the Qomolangma detachment (Fig. 6b). This suggests that the STD system cuts progressively down-section to the north as shown on the restoration in Figures 2c and 5. Leucogranites in the GHS are mostly sill complexes intruded parallel to the foliation, with less common late dykes that cross-cut the ductile fabrics, but are truncated by the brittle fault above. Searle et al. (2003) defined three major sets of leucogranite sills and dykes, which are broadly comparable across the entire region. These three sets are illustrated from outcrops in the Hermit's Gorge, above the Rongbuk Base camp in Tibet (Fig. 7a,b). Set 1 leucogranites are parallel to the ductile foliation, folded with the metamorphic schistosity and have a weak fabric. Set 2 leucogranite dykes cross-cut the metamorphic fabric, and both sets 1 and 2 are cut by the later set 3 dykes. Set 3 dykes are undeformed garnet+ tourmaline + muscovite + biotite leucogranites that cross-cut the ductile fabrics. No leucogranites cut across the brittle STD fault anywhere along the Himalaya, despite earlier reports to the contrary (Manaslu granite (LeFort 1981; Harrison et al. 1999), and the Rongbuk granite (Burchfiel et al. 1992; Hodges et al. 1998)). Along the East Rongbuk glacier the Lhotse detachment truncates foliation in the footwall gneisses (Fig. 8a). The Yellow Band calc-silicates thin towards the north, but form a prominent layer between the highgrade gneiss below and the unmetamorphosed sediments above (Fig. 8b and c). Rocks above the Qomolangma detachment are sedimentary and composed dominantly of calcite with minor amounts of dolomite, quartz and illite-muscovite (Fig. 8d). Rocks beneath the Lhotse detachment show ductile deformation with lower set 1 leucogranite sills folded in with the pelitic and calc-silicate gneisses (Fig. 8e). Structurally higher in the section, flattening fabrics are common with thin leucogranite sills showing boudinage structures (Fig. 8f). Microstructures within the zone indicate top-tonorth kinematics, compatible with extrusion of footwall gneisses to the south beneath a static

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

367

Fig. 6. (a) Northernmost outcrop of the South Tibetan (STD) along the Rongbuk valley in Tibet, approximately 38 km north of the summit of Everest. Massive cliffs of layered leucogranite (Lg) pass up into ductile deformed gneisses and calc-silicates with fewer leucogranite sills. Metamorphic grade decreases up-section to the Qomolangma Detachment, which places unmetamorphosed limestones and shales above the Greater Himalayan sequence (GHS) metamorphic rocks and leucogranites. (b) STD shear zone approximately 60 km north of Everest along the Kharta valley showing upward decrease in volume of leucogranite. The entire section is a large-scale shear zone with ductile strain decreasing upwards towards the base of the Tethyan sedimentary sequence. Cliff section is approximately 80 m high.

368

M.P. SEARLE E T A L .

Fig. 7. (a, b) Hermit's Gorge, 5430 m altitude, east of Everest Base camp, Rongbuk glacier. Early, set 1 leucogranite sills are parallel to, and folded in with, the metamorphic schistosity. Set 2 dykes cut the metamorphic schistosity and some of the folds. Both sets 1 and 2 sills and dykes are cut by later, vertical leucogranite dykes (set 3) that abruptly cut folds, metamorphic schistosity and ductile fabric in the gneisses. These late dykes have little or no fabric within, but are cut by the flat-lying brittle Qomolangma Detachment at higher structural levels. Cliff section in (b) is approximately 20 m high.

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

369

Fig. 8. (a) Unmetamorphosed Ordovician limestones and mudstones lying above a thin band of calc-silicate (Yellow Band) and high-grade gneisses with leucogranite sills; cliffs east of the East Rongbuk glacier above 6000 m. (b) Cliffs of white leucogranite sills intruding gneisses above the East Rongbuk glacier passing up into thin banded calc-silicate and finally into unmetamorphosed sedimentary rocks at top of the cliffs. (c) Cliff section above ice penitents of the East Rongbuk glacier. Unmetamorphosed Ordovician sedimentary rocks at the top overlie layered calc-silicates and leucogranite sills below. (d) Highest level of the Qomolangma Detachment at c. 6000 m above Rongbuk monastery showing the structurally lowest outcrop of sedimentary rocks. (e) Folded set 1 leucogranite sills with pelites and calc-silicate gneisses immediately above Rongbuk Base Camp site. (f) Flattening boudinage structures in leucogranite sills within the upper level of the Greater Himalayan sequence (GHS) beneath the Lhotse Detachment, lower Rongbttk valley.

hanging wall. Kinematics, vorticity and restoration of the GHS (Searle et al. 2002, 2003; Law et al. 2004; Jessup et al. 2006) show that the STD acted as a 'passive roof fault' (cf. Banks & Warburton 1986), or a 'stretching fault' in the sense of Means (1989). The footwall rocks were extruded

out to the south beneath the fault, whilst the hanging wall rocks remained static. Using pressure-depth constraints from footwall gneisses and leucogranites, Walker et al. (1999) estimated at least 50 km of southward extrusion of GHS rocks in Zanskar, and Searle et al. (2002, 2003)

370

M.P. SEARLE E T AL.

estimated at least 100 km of southward motion may have occurred in the footwall GHS rocks along the Rongbuk valley, north of Everest. This supports the channel flow model of ductile extrusion of the middle crust. Geochronology suggests that localized partial melting may have triggered channel flow, and that the rapid cooling of the entire GHS slab may have caused channel flow to end abruptly at around musco16-14 Ma (Godin e t al. 2006a). 4~ vite ages, interpreted as a proxy for timing of the ductile to brittle transition, across the entire slab also lend support to this. It is possible, however, that the GHS channel may still be active to the north, at depth beneath south Tibet, and the rocks along the High Himalaya record the 20-16 Ma palaeo-channel.

Timescales of metamorphism, melting and channel flow Figure 9 shows geochronological data for the Everest transect plotted on a temperature versus

time diagram, with the known closure temperatures for individual minerals in each system on the right. U - P b dating of monazites growing in equilibrium with garnet, kyanite and sillimanite-cordierite shows that peak metamorphism spanned at least 14 million years, from 32.2 _ 0.4 Ma to 17.9 _ 0.5 Ma (Simpson et al. 2000; Searle e t al. 2003). Large-scale melting to form the Everest leucogranites occurred between 21.3 and 20.5 Ma (Simpson e t al. 2000) corresponding to the timing of peak high-temperature sillimanite-cordieritegrade metamorphism along the top of the GHS slab. Rapid cooling, and therefore rapid exhumation and probably a period of high erosion rates, followed immediately after crustal melting. If crustal thickening and regional metamorphism started soon after the India-Asia collision at c. 50 Ma then approximately 15-20 million years were required to thicken the crust and metamorphose it to kyanite-sillimanite-grade P - T conditions. Temperatures within the GHS slab must then have remained very high for at least 16 million years (from 32 to 16 Ma), with sporadic episodes of

Fig. 9. Temperature-time plot showing all geochronological data from the Everest region. Closure temperatures of individual minerals in various isotopic systems are shown on the right-hand side. Monazite ages are interpreted as timing the growth of metamorphic assemblages at or close to the peak, and in granites as timing the crystallization of the granite. Sources of data are Hubbard & Harrison (1989), Bergman et al. (1993), Hodges et al. (1998), Murphy & Harrison (1999), Hubbard & House (2000), Simpson et al. (2000) and Searle et al. (2003). The steep part of the cooling path at c. 16 Ma is interpreted as the timing of rapid motion of STD footwall rocks as they were exhumed towards the surface.

EVEREST STRUCTURE, RESTORATION AND EVOLUTION crustal melting resulting in partial anatexis of the Namche migmatite at 25 Ma (Viskupic & Hodges 2001), and the main pulse of crustal melting forming the Everest-Nuptse-Makalu leucogranites at 21 Ma (Simpson et al. 2000). Temperatures within the slab remained high enough for leucogranite generation until 16 Ma when the youngest granite dykes were formed in the Rongbuk area (Murphy & Harrison 1999). A l l 4~ and K Ar muscovite and biotite ages of GHS metamorphic rocks are older than 14 Ma, suggesting that by that time, the entire GHS slab had cooled to below 350-300~ as the channel cooled, ductile strain had turned to brittle faulting, and thrusting propagated down-structural-section into the Lesser Himalaya. The timing of metamorphism and melting in the Everest transect closely matches other profiles along the Himalayan chain. In Zanskar, the GHS was metamorphosed to kyanite grade at 33-31 Ma as evidenced by S m - N d dating of garnet (Vance & Harris 1999) and U - P b dating of metamorphic monazite (Walker et al. 1999). Temperatures remained high during active crustal shortening and thickening for at least a further 12 million years until 22-20 Ma, when peak sillimanite-grade metamorphism across the GHS slab was coincident with widespread migmatization and crustal melting producing tourmaline, garnet and muscovite-bearing leucogranites (Noble & Searle 1995; Dezes et al. 1999; Walker et al. 1999). Protoliths of the GHS metamorphic rocks are Proterozoic to early Mesozoic sediments and Permian volcanic rocks (Panjal Traps), partly lateral equivalents to the unmetamorphosed rocks in the Tethyan zone to the north (Searle 1986; Walker et al. 2001). T h - P b monazite ages within the sillimanitegrade gneisses of the GHS from Pheriche south to Lukla in the Everest transect (Fig. 4) are all within the range 25.3-20.7 Ma (Catlos et al. 2002). Only in the far south around the village of Karikhola do the T h - P b monazite ages become younger, from 15.2 to 11.8 Ma. The monazites are inclusions within garnet that was apparently growing at 530 ___50~ (Catlos et al. 2002), a temperature similar to that required for monazite growth in pelite (525~ Smith & Barreiro 1990). Catlos et al. (2002), following Harrison et al. (1997), interpreted these ages as representing late Miocene reactivation of the MCT. Even younger T h - P b monazite ages (c. 7 - 3 Ma) have been dated in rocks from a similar position along the MCT in central Nepal and Garhwal (Harrison et al. 1997; Catlos et al. 2002). Young T h - P b monazite ages from the MCT present a conundrum for interpreting the tectonic evolution of the MCT. 40 Ar/ 39Ar hornblende ages

371

of c. 21 Ma from the upper part of the MCT zone south of Lukla, record the timing of cooling through the 500~ isotherm (Hubbard & Harrison 1989). Along most of the Greater Himalaya, 4~ and K - A r muscovite and biotite ages, which record timing of cooling through 300350~ isotherms, are always older than 14 Ma across the entire slab (Godin et al. 2001). How is it possible, therefore, that 530~ metamorphic temperatures in the MCT samples with young monazite ages do not reset the 4~ and K - A r systematics? Recent studies have shown that monazite and even zircon are not only mobile, but can crystallize at temperatures below 350~ in slates (Rasmussen et al. 2001" Dempster et al. 2004). Rubin et al. (1993) showed that zirconium can be a mobile phase during hydrothermal alteration. Young matrix monazites within the MCT zone could therefore have grown during post-metamorphic hydrothermal alteration, but it is harder to explain the monazite inclusions within armouring garnet, unless minute cracks in the garnet allowed access of fluids during hydrothermal activity. The MCT zone has been, and still is, a conduit for boiling fluids, with numerous hot springs located along the thrust (the location of many villages called Tatopani - Nepali for 'hot water'). If the young monazite ages along the MCT zone do record a period of elevated temperatures, it is a very restricted metamorphism solely along the MCT and has no record higher up throughout most of the GHS. The T h - P b monazite ages south of Kharikhola span 15.2-11.8 Ma (Catlos et al. 2002) and may indicate the youngest period of garnet growth in the Phaplu gneiss, immediately prior to cooling below the ductilebrittle transition.

Discussion In the channel flow model (Fig. 3) for the Himalaya and Tibet, hot, weak, partially molten and thickened crust flows from the Tibetan plateau towards the plateau margins in response to lithostatic pressure gradients between the high elevation and thick crust of Tibet and the low elevation and normal thickness of the Indian plate crust. Channel flow processes have been incorporated into both linear viscous models and coupled thermal-mechanical models (Royden et al. 1997; Beaumont et al., 2001, 2004). Previously, the deformation has been thought of in terms of a 'jelly sandwich' model in which the upper crust and upper mantle are strong, and are separated by a weak and ductile lower crust (Jackson et al. 2004). However, in the Himalaya it is the middle crust that is hot, weak and flowing, whereas the

372

M.P. SEARLE E T A L .

granulitic lower (Indian) crust is relatively strong. The lower crust of Tibet is not exposed so we have to rely on gravity data (Jin et al. 1994), seismic data (Nelson et al. 1996) and volcanic xenoliths (Hacker et al. 2000) to interpret the structure. The INDEPTH seismic experiments in southern Tibet indicate the presence of a fluid-like layer of middle crust (Nelson et al. 1996; Brown et al. 1996) interpreted as containing a degree of partial melt. This mid-crust layer can be followed south to link with the GHS and the bounding shear zones of the MCT and STD. We suggest that the crust beneath the Himalaya and southern Tibet must be rheologically layered, with a major decoupling of the upper, middle and lower crust (Fig. 10). The geotherm is not linear with depth but shows a peak temperature at high levels in the middle crust, corresponding to the location of the 'bright spots' imaged by magnetotelluric data beneath

Fig. 10. Proposed geotherm beneath the southern part of the Tibetan plateau, using structural depths and temperatures from the Everest transect across the Greater Himalayan sequence (GHS) (Hubbard 1996; Searle et al. 2003) as a reverse uniformitarian analogue (Miocene P-T profile across the GHS is used as a guide to the present temperature-depth profile beneath the plateau). The positions of the 'bright spots' and temperatures in the lower crust and upper mantle are after Kola-Ojo & Meissner (2001). The mid-crustal channel, bounded by the Main Central Thrust (MCT) below and the South Tibetan detachment (STD) above, is shaded with the location of the 'bright spots' imaged on the INDEPTH deep seismic profile, and interpreted as leucogranite magmas forming today, also shown. The lower crust, between about 35 and 75 km, is composed of Indian shield granulites underthrust beneath the southern margin of the Lhasa block.

the Lhasa block (Kola-Ojo & Meissner 2001). The lower crust is composed of the more rigid, dry granulites of the underthrusting Indian shield. Earthquakes in southern Tibet are abundant in the upper crust down to c. 18 km and again in the lowermost crust at 60-80 km, but not in the middle crust (Jackson et al. 2004), in agreement with the non-linear isotherm shown in Figure 10. There is however, some dispute as to whether the deep earthquakes beneath the Himalaya and south Tibet are actually in the lower crust or in the upper mantle (Chen & Yang 2004). The obvious question now is why do melting temperatures exist at relatively shallow levels of the thick crust beneath the south Tibetan plateau? The restoration of the Everest transect (Fig. 5) shows that the leucogranites were derived from the Proterozoic shales along a stratigraphic horizon that must have formed the protolith of the migmatites and leucogranites. The heat source for melting cannot have been frictional heating along the MCT because the site of anatexis is along much higher crustal levels than the MCT, and in situ melting is not seen along the MCT. We suggest that the shallow level of high temperatures producing leucogranite melts must have been the consequence of a high concentration of radioactive heat-producing elements at that unique stratigraphic horizon in the crust. The concentration of U, Th and K could have been the result of sedimentary processes within the late Proterozoic basin that accumulated debris off the eroding Archaean mountain belt of peninsula India. Another intriguing question remaining is: why did channel flow operate within a relatively narrow time span (24-17 Ma), and why did it apparently end at c. 16 Ma? High temperatures, close to leucogranite melting temperatures (680720~ are required to maintain ductile channel flow. Numerous 4~ and K - A r cooling ages on micas within the GHS are almost entirely older than 15-14 Ma (see Searle & Godin (2003) and Godin et al. (2006a,b) for a summary). Temperatures in the GHS metamorphic rocks and leucogranites must therefore have been below 350~ by 15-14 Ma, and the whole channel must have cooled through the ductile-brittle transition. This fact makes the subsequent late Miocene-Pliocene monazite ages from the MCT zone (Harrison et al. 1997; Catlos et al. 2002) hard to interpret as indicators of a regional metamorphic event. It is possible that the Himalayan mid-crustal channel could still be active today (Hodges et al. 2001), although there is no seismic evidence for recent activity on the MCT or the STD. The GHS along the High Himalaya corresponds to the zone of highest topography, although some of the highest peaks along the range (e.g. Dhaulagiri-Annapurna

EVEREST STRUCTURE, RESTORATION AND EVOLUTION

373

Fig. 11. Sketch section across the Himalaya and south Tibet showing how the channel flow model could work if channel flow was operating today. The folded metamorphic isograds linking the STD footwall and MCT hanging wall represent 'frozen' 20-15 Ma isograds adjacent to a younger and more active channel of partial melting present beneath southern Tibet today. This partially molten mid-crustal channel separates the seismogenicupper crust, dominated today by east-west extension, and the granulite facies lower crust dominated by north-south compression. Indian Shield granulites subducted to 50-75 km depth beneath south Tibet would remain at granulite facies unless hydrated, in which case they would convert to eclogite facies rocks. The deepest earthquakes occur along the base of the crust as it flexes down beneath the Himalaya. peaks) are composed of the lower Tethyan sedimentary rocks above the STD. Hurtado et al. (2001) suggested that one branch of the STD north of Annapurna may have been active during the Quaternary. The MCT is coincident with prominent knick-points in Himalayan river profiles (Seeber & Gornitz 1983), but the hanging wall uplift could be due either to MCT reactivation (Harrison et al. 1997, 1998) or to uplift above a ramp along the deeper and younger Main Himalayan thrust (MHT) fault (Cattin & Avouac 2000). If the high topography of the Himalaya is due to channel flow during the period 24-16 Ma, why does the range remain so high today? One solution would be to have the Miocene channel preserved along the Himalaya today, as recorded by the geological structure and pressuretemperature-time conditions, as an outer rind or crust surrounding the hot, active channel presently at mid-crustal depths beneath the southern part of the Tibetan plateau (Fig. 11). This could be envisaged as analogous to a cooling Hawaiian-type lava flow with a hot, molten core surrounded by a cooling brittle carapace.

Conclusions Field mapping, macro- and micro-structural observations, and thermal modelling all support the

conclusion that the channel flow model is appropriate for the GHS. A slab of Indian plate middle crust, between 5 and 20 km thick, was at fairly uniform high temperatures (c. 600-700~ for a period of about 12-16 million years during the late Oligocene and early Miocene (between 30 and 17 Ma), sandwiched between a brittle-deforming upper crust (Tethyan sedimentary zone) and a rigid granulitic lower crust (subducting Indian shield). The mid-crustal channel was metamorphosed at sillimanite-grade temperatures and contained a zone of in situ early Miocene migmatization, up to 5 - 8 km thick. The GHS channel was bounded by two major ductile shear zones with thrust sense of shear along the base (MCT), and a normal sense of shear above (STD). Structural geometry, thermobarometry and geochronology suggest that there was a mid-crustal layer composed of high-temperature sillimanitegrade rocks that were close to partial melting temperatures during the period 32-17 Ma. Migmatites and crustal melt leucogranites are widespread at the upper structural levels of this layer. U - P b dating of metamorphic monazites crystallizing in equilibrium with garnet and kyanite, indicates that early metamorphism in the Everest area peaked at 32.2 _+ 0.4Ma (Simpson et al. 2000). Similar ages were obtained from S m - N d dating of garnets in the Zanskar and Garhwal regions in the western Himalaya (Vance & Harris 1999;

374

M.P. SEARLE ET AL.

Prince et al. 2001). It appears that, although high temperatures were attained by 32 Ma, melting did not occur along the top of the GHS until about 24-21 Ma, at least in the Everest-Makalu region (Scharer 1984; Simpson et al. 2000). These ages suggest that it could have been partial melting that triggered channel flow and ductile extrusion. The timing of granite melting also coincides with the timing of increased cooling rates, erosion rates and exhumation of the GHS, and increased sedimentation rates in the Siwalik foreland basin. As the granites cooled, flow ended and lateral transport was achieved by brittle faulting along the outer margins of the channel, both below (MCT) and above (STD). Melting resulted in several large leucogranite bodies that reach a maximum thickness of c. 3 km in the Makalu-Chomolonzo leucogranite body. Himalayan leucogranites contain tourmaline, garnet, muscovite and biotite and are pure crustal melts, derived from a highly radiogenic Proterozoic sedimentary source, corresponding to the Haimanta Group shales (or Cheka Group in Bhutan). The shallow level of crustal melting within the Himalayan crust must have been the result of a high concentration of heat-producing elements within the stratigraphic horizon of the Proterozoic black shales. The Himalayan granites were intruded as sill complexes within the sillimanite gneisses for large distances (over 70-100 km north of Everest) laterally along the foliation planes. Microstructures within the GHS channel indicate that there was a significant component of penetrative vertical pure shear, as well as the dominant south-directed simple shear (Law et al. 2004; Jessup et al. 2006). However, despite the evidence for a reduction in the thickness of the channel through flattening, there is also good evidence that the active channel expanded with time, during a change from ductile to brittle deformation, as MCT zone thrusts propagated down-structuralsection with time, and STD normal faults propagated up-section with time. Thrust faults along the base of the channel and normal faults along the top must have been synchronous, as previously suggested (Searle & Rex 1989; Hodges et al. 1993, 1996; Searle et al. 1997, 2003). STD normal faults were active in a wholly compressional environment, and acted as a passive roof fault during crustal shortening and thickening in the footwall. Although pure shear flattening did compress the sequence, new material was constantly being fed into the channel from crustal subduction beneath the active MCT during the late Oligocene and Miocene. Folding and thrusting in the Tethyan sedimentary rocks above the STD occurred earlier than shortening and thickening in the GHS beneath, although some breakback

thrusting and north-directed backthrusting did occur late in the sequence (Searle 1986; Corfield & Searle 2000; Murphy & Yin 2003). The STD normal faults do not relate to wholescale crustal thinning, whole crustal extension or topographic collapse. This work was funded by NERC grant NER/K/S/2000/ 0951 to M.P.S and NSF grant EAR 0207524 to R.D.L. We are grateful to Randy Parrish, Dave Waters and Rob Simpson for extensive discussions, and to Brad Hacker and Mike Murphy for reviews. We also are very grateful to Tashi Sherpa, Ting Lei, Sonan Wangdu, Rene Schrama and Shiva Dhaktal for logistics. This paper is dedicated to Doug Nelson in memory of his enthusiastic and scholarly input into tapping the mysteries of Tibetan structure, and for breaching the geology-geophysics mental barrier.

References ALSDORF, D. & NELSON, K. D. 1999. Tibetan satellite magnetic low: Evidence for widespread melt in the Tibetan crust? Geology, 27, 943-946. ALSDORF, D., BROWN, L., NELSON, K. D., MAKOVSKY, Y., KLEMPERER, S. & ZHAO, W. 1998. Crustal deformation of the Lhasa terrane from Project INDEPTH deep seismic reflection profiles. Tectonics, 17, 501-519. ARITA, K. 1983. Origin of the inverted metamoqghism of the lower Himalaya, central Himalaya. Tectonophysics, 95, 43-60. BANKS, C. J. & WARBURTON,J. 1986. 'Passive-roof' duplex geometry in the frontal structures of the Kirthar and Sulaiman mountain belts, Pakistan. Journal of Structural Geology, 8, 229-237. BEAUMONT, C., JAMIESON,R. A., NGUYEN,M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON,R. A., NGUYEN,M. H. & MEOVEI)EV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06406. BERGMAN, S. C., COFFIELD, D. Q., DONELICK, R., CORRIGAN, J., TALBOt, J., CERVENY, P. & KELLEY, S. 1993. Late Cenozoic compressional and extensional cooling and exhumation of the Qomolangma ( M t . Everest) region, Nepal. Geological Society of America Abstracts with Programs, 25(6), A-176 (abstract). BIRD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10,27510,286. BROWN, L. D., ZHAO, W., NELSON, K. D. eTAL. 1996. Bright spots, structure and magmatism in Southern Tibet from INDEPTH seismic reflection profiling. Science, 274, 1688-1690.

EVEREST STRUCTURE, RESTORATION AND EVOLUTION BURCHFIEL, B. C., ZHILIANG, C., HODGES, K. V., YUPING, L., ROYDEN, L., CHANGRONG, D. & JIENE, X. 1992. The South Tibetan Detachment System, Himalayan Orogen: extension contemporaneous with and parallel to shortening in a collisional mountain belt. Geological Society of America, Special Paper, 269. BURG, J.-P. 1983. Carte Gdologique du Sud du Tibet. Ministry of Geology, Peking, and CNRS, Paris. BURG, J. P., BRUNEL,M., GAPAIS,D., CHEN, G. M. & L1U, G. H. 1984. Deformation of leucogranites of the crystalline Main Central thrust sheet in southern Tibet (China). Journal of Structural Geology, 6, 535-542. CAROSI, R., LOMBARDO,B., MOLLI,G., MUSUMECI,G. & PERTUSATI,P. 1998. The South Tibetan detachment system in the Rongbuk valley, Everest region. Deformation features and geological implications. Journal of Asian Earth Sciences, 16, 299-311. CAROSI, R., LOMBARDO, B., MUSUMECI, G. & PERTUSATI, P. 1999a. Geology of the Higher Himalayan crystallines in Khumbu Himal (Eastern Nepal). Journal of Asian Earth Sciences, 17, 785-803. CAROSI, R., MUSUMECI,G. & PERTUSATI,P. C. 1999b. Extensional tectonics in the higher Himalayan crystallines of Khumbu Himal, eastern Nepal. Geological Society of America Special Paper 328, 211 - 223. CATLOS, E. J., HARRISON, T. M., MANNING, C. E., GROVE, M., RAI, S. M., HUBBARD,M. & UPRETI, B. N. 2002. Records of the evolution of the Himalayan orogen from in situ Th-Pb ion microprobe dating of monazite: Eastern Nepal and western Garhwal. Journal of Asian Earth Sciences, 20, 459 -48O. CATTIN, R. ~i; AVOUAC,J.-P. 2000. Modeling mountain building and the seismic cycle in the Himalaya of Nepal. Journal of Geophysical Research, 105, 13,389-13,407. CHEN, W.-P. & YANG, Z. 2004. Earthquakes beneath the Himalaya and Tibet: Evidence for strong lithospheric mantle. Science, 304, 1949-1952. CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. COLCHEN, M., LE FORT, P. & PECHER, A. 1986. Annapurna-Manaslu-Ganesh Himal. CNRS, Paris. CORFIELD, R. I. & SEARLE, M. P. 2000. Crustal shortening estimates across the north Indian continental margin, Ladakh, NW India. In: KHAN, M. A., TRELOAR, P. J., SEARLE, M. P. & JAN, M. Q. (eds) Tectonics of the Nanga Parbat Syntaxis and the Western Himalaya. Geological Society, London, Special Publications, 170, 395410. DECELLES P. G., GEHRELS, G. E., QUADE,J., LAREAU, B. & SVURLIN,M. 2000. Tectonic implications of U-Pb zircon ages of the Himalayan orogenic belt in Nepal. Science, 288, 497-499. DEMPSTER, T. J., HAY,D. & BLUCK,B. J. 2004. Zircon growth in slate. Geology, 32, 221-224.

375

DEZES, P. J., VANNAY,J.-C., STECK, A., BUSSY, F. & COSCA, M. 1999. Synorogenic extension: Quantitative constraints on the age and displacement of the Zanskar shear zone (northwest Himalaya). Geological Society of America Bulletin, 111, 364-374. DUCEA, M. N., LUTKOV, V., MINAEV, V. T. Er AL. 2003. Building the Pamirs: The view from the underside. Geology, 31, 849-852. DUNCAN, C., MASEK, J. & FIELDING, E. 2003. How steep are the Himalaya? Characteristics and implications of along-strike topographic variations. Geology, 31, 75-78. FIELDING, E., ISACKS, B., BARAZANGI, M. & DUNCAN, C. 1994. How flat is Tibet? Geology, 22, 163-167. FRANCHETEAU, J., JAUPART, C., XIAN, J. S. ET AL. 1984. High heat flow in southern Tibet. Nature, 307, 32-36. GAILLARD, F., SCAILLET, B. • PICHAVANT,M. 2004. Evidence of present-day leucogranite pluton growth in Tibet. Geology, 32, 801-804. GANSSER, A. 1964. Geology of the Himalaya. John Wiley & Sons, Chichester. GANSSER, A. 1983. Geology of the Bhutan Himalaya. Birkh~iuser Verlag, Basle. GODIN, L., PARRISH, R. R., BROWN,R. D. & HODGES, K. V. 2001. Crustal thickening leading to exhumation of the Himalayan core of central Nepal: insights from U-Pb geochronology and 40Ar/ 39Ar thermochronology. Tectonics, 20, 729-747. GODIN, L., GRUJIC, D., LAW, R. D. & SEARLE, M. P. 2006a. Channel flow, extrusion and exhumation in continental collision zones: an introduction. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 1-23. GODIN, L., GLEESON, T., SEARLE, M. P., ULLRICH, T. D. & PARRISH, R. g. 2006b. Locking of southward extrusion in favour of rapid crustal-scale buckling of the Greater Himalayan sequence, Nar valley, central Nepal. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society London, Special Publications, 268, 269-292. GRASEMANN, B., FRITZ, H. & VANNAY, J.-C. 1999. Quantitative kinematic flow analysis from the Main Central Thrust zone, (NW Himalaya): implications for a decelerating strain path and the extrusion of orogenic wedges. Journal of Structural Geology, 21, 837-853. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L., KUNDIG, K., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. GRUJIC, D., HOLLISTER, L. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel: insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191.

376

M.P. SEARLE ETAL.

HACKER, B. R., GNOS, E., RATSCHBACHER, L. ETAL. 2000. Hot and dry deep crustal xenoliths from Tibet. Science, 287, 2463-2466. HARRISON, T. M., RYERSON, F. J., LE FORT, P., YIN, A., LOVERA, O. & CATLOS, E. J. 1997. A Late Miocene-Pliocene origin for the central Himalayan inverted metamorphism. Earth and Planetary Science Letters, 146, E l - E 8 . HARRISON, T. M., GROVE, M., LOVERA, O. & CATLOS, E. J. 1998. A model for the origin of Himalayan anatexis and inverted metamorphism. Journal of Geophysical Research, 103, 27,017-27,032. HARRISON, T. M., GROVE, M., MCKEEGAN, K. D., COATH, C. D., LOVERA, 0. 8r LEFORT, P. 1999. Origin and episodic emplacement of the Manaslu Intrusive Complex, Central Himalaya. Journal of Petrology, 40, 3-19. HAUCK,M. L., NELSON, K. D., BROWN, L. D., ZHAO, W. & ROSS, A. R. 1998. Crustal structure of the Himalayan orogen at approximately 90 ~ east longitude from Project INDEPTH deep reflection profiles. Tectonics, 17, 481-500. HODGES, K. V. 2000. Tectonics of the Himalaya and southern Tibet from two perspectives. Geological Society of America Bulletin, 112, 324-350. HODGES, K. V., PARRISH, R. R., HOUSCH, T. B., LUX, D. R., BURCHFIEL, B. C., ROYDEN, L. H. & CHEN, Z. 1992. Simultaneous Miocene extension and shortening in the Himalayan orogen. Science, 258, 1466-1470. HODGES, K. V., BURCHFIEL, B. C., ROYDEN, L. H., CIaEN, Z. & LIU, Y. 1993. The metamorphic signature of contemporaneous extension and shortening in the central Himalayan orogen: data from the Nyalam transect, southern Tibet. Journal of Metamorphic Geology, 11, 721-737. HODGES, K. V., PARRISH, R. R. & SEARLE,M. P. 1996. Tectonic evolution of the central Annapurna Range, Nepalese Himalaya. Tectonics, 15, 1264-1291. HODGES, K. V., BOWRING, S., DAVIDEK, K., HAWKINS, D. & KROL, M. 1998. Evidence for rapid displacement on Himalayan normal faults and the importance of tectonic denudation in the evolution of mountain ranges. Geology, 26, 483 -486. HODGES, K. V., HURTADO, J. M. & WHIPPLE, K. X. 2001. Southward extrusion of Tibetan crust and its effect on Himalayan tectonics. Tectonics, 20, 799-809. HUBBARD, M. S. 1989. Thermobarometric constraints on the thermal history of the Main Central Thrust Zone and Tibetan Slab, eastern Nepal Himalaya. Jaurnal of Metamorphic Geology, 7, 19-30. HUBBARD, M. S. 1996. Ductile shear as a cause of inverted metamorphism: Example from the Nepal Himalaya. Journal of Geology, 194, 493-499. HUBBARD, M. S. & HARRISON, T. M. 1989.4~ age constraints on deformation and metamorphism in the Main Central Thrust zone and Tibetan Slab, eastern Nepal Himalaya. Tectonics, 8, 865-880. HUBBARD, M. S. & HOUSE, M. 2000. Low temperature dating of high mountain rocks: (U-Th)/He ages from Higher Himalayan samples, Eastern Nepal.

Earth Sciences Frontiers, Beijing, 15 th HimalayaKarakoram-Tibet Workshop, Chengdu, 16-17, (abstract). HURTADO, J. M., HODGES, K. V. • WHIPPLE, K. X. 2001. Neotectonics of the Thakkhola graben and implications for recent activity on the South Tibetan Fault system in the central Nepalese Himalaya. Geological Society of America Bulletin, 113, 222-240. JACKSON, J., AUSTRHEIM, H., MCKENZIE, D. & PREISTLEY, K. 2004. Metastability, mechanical strength, and the support of mountain belts. Geology, 32, 625-628. JAMIESON, R. A., BEAUMONT, C., MEDVEDEV 8r NGUYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06407. JESSUP, M. J., TRACY, R., SEARLE,M. P. & LAW, R. D. 2004. Staurolite schist marks fight-way up metamorphic isograds at the top of the High Himalayan Slab: Mount Everest, Tibet/Nepal. In: SEARLE, M. P., LAW, R. D. GODIN, L. (convenors) Channel Flow, Ductile Extrusion and Exhumation of Lower-mid Crust in Continental Collision Zones, Abstract Books, Geological Society, London. JESSUP, M. J., LAW, R. D., SEARLE,M. P. & HUBBARD, M. S. 2006. Structural evolution and vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/Nepal: implications for orogen-scale flow partitioning. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 379-413. JIN, Y., MCNUTT, M. K. & ZHU, Y. 1994, Evidence from gravity and topography data for folding of Tibet. Nature, 371, 669-674. KIND, R., NI., J., ZHAO, W. ETAC. 1996. Evidence from earthquake data for a partially molten crustal layer in southern Tibet. Science, 274, 1692-1694. KOLA-OJO, O. & MEISSNER, R. 2001. Southern Tibet: its deep seismic structure and some tectonic implications. Journal of Asian Earth Sciences, 19, 249-256. KOSAREV, G., KIND, R., SOBOLEV, S. V., YUAN, X., HANKA, W. & ORESHIN, S. 1999. Seismic evidence for a detached Indian lithospheric mantle beneath Tibet. Science, 283, 1306-1309. LAW, R. D., SEARLE, M. P. & SIMPSON, R. L. 2004. Strain, deformation temperatures and vorticity of flow at the top of the Greater Himalayan Slab, Everest massif, Tibet. Journal of the Geological Society, London, 161, 305-320. LE FORT, P. 1981. Manaslu leucogranite: a collision signature of the Himalaya-a model for its genesis and emplacement. Journal Geophysical Research, 86, 10,545-10,568. LOMBARDO, B., PERTUSATI, P. & BORGI, S. 1993. Geology and tectonomagmatic evolution of the eastern Himalaya along the Chomolungma Makalu transect. In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological

EVEREST STRUCTURE, RESTORATION AND EVOLUTION Society, London, Special Publications, 74, 341355. MEANS, W. D. 1989. Stretching faults. Geology, 17, 893-896. MILLER, C., THONI, M., FRANK, W., GRASEMAN, B., KLOTZLI, U., GUNTLI, P. & DRAGANTIS, E. 2001. The early Palaeozoic magmatic event in the Northwest Himalaya, India: source, tectonic setting and age of emplacement. Geological Magazine, 138, 237-251. MURPHY, M. A. & HARRISON, T. M. 1999. Relationship between leucogranites and the Qomolangma Detachment in the Rongbuk valley, south Tibet. Geology, 27, 831-834. MURPHY, M. A. & YIN, A. 2003. Structural evolution and sequence of thrusting in the Tethyan fold-thrust belt and Indus-Yalu suture zone, southwest Tibet. Geological Society of America Bulletin, 115, 21-34. NELSON, K. D., ZHAO, W. J., BROWN, L. D. ET AL. 1996. Partially molten middle crust beneath southern Tibet: synthesis of Project INDEPTH results. Science, 274, 1684-1688. NOBLE, S. R. & SEARLE, M. P. 1995. Age of crustal melting and leucogranite formation from U-Pb zircon and monazite dating in the western Himalaya, Zanskar, India. Geology, 23, 1135-1138. OWENS, T. J. & ZANDT, G. 1997. Implications of crustal property variations for models of Tibetan plateau evolution. Nature, 387, 37-43. PARRISH, R. R. & HODGES, K. V. 1996. Isotopic constraints on the age and provenance of the Lesser and Greater Himalayan sequences, Nepalese Himalaya. Geological Society of America Bulletin, 108, 904-911. POGNANTE, U. & BENNA,P. 1993. Metamorphic zonation, migmatization and leucogranites along the Everest transect of eastern Nepal and Tibet: record of an exhumation history, In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 323-340. PRINCE, C., HARRIS, N. & VANCE, D. 2001. Fluidenhanced melting during prograde metamorphism. Journal of Geological Society, London, 158, 233 -241. RASMUSSEN, B., FLETCHER, I. R. & MACNAUGHTON, N. J. 2001. Dating of low grade metamorphic events by SHRIMP U-Pb analysis of monazite in shale. Geology, 29, 963-966. ROBINSON, D., DECELLES, P. G., GARZIONE, C. N., PEARSON, O. N., HARRISON, T. M. & CATLOS, E. J. 2003. Kinematic model for the Main Central thrust in Nepal. Geology, 31, 359-362. ROBYR, M., VANNAY,J.-C., EPARD, J.-L. & STECK, A. 2002. Thrusting, extension, and doming during the polyphase tectonometamorphic evolution of the high Himalayan crystalline zone in NW India. Journal of Asian Earth Sciences, 21, 221-239. ROYDEN, L. H., BURCHFIEL, B. C., KING, R. W., WANG, E., CHEN, Z., SHEN, F. & LIU, Y. 1997. Surface deformation and lower crustal flow in eastern Tibet. Science, 276, 788-790.

377

RUBIN, J. N., HENRY, C. D. & PRICE, J. G. 1993. The mobility of zirconium and "immobile" elements during hydrothermal alteration. Chemical Geology, 110, 29-47. SCHARER, U. 1984. The effect of initial 23~ disequilibration on U-Pb ages: the Makalu case. Earth and Planetary Science Letters, 67, 191-204. SEARLE, M. P. 1986. Structural evolution and sequence of thrusting in the High Himalaya, Tibetan-Tethys and Indus suture zones of Zanskar and Ladakh, Western Himalaya. Journal of Structural Geology, 8, 923-936. SEARLE, M. P. 1999a. Extensional and compressional faults in the Everest-Lhotse massif, Khumbu Himalaya, Nepal. Journal of Geological Society, London, 156, 227-240. SEARLE, M. P. 1999b. Emplacement of Himalayan leucogranites by magma injection along giant sill complexes: examples from the Cho Oyu, Gyachung Kang and Everest leucogranites (Nepal Himalaya). Journal of Asian Earth Sciences, 17, 773 -783. SEARLE, M. P. 2003. Geological Map of the Everest Massif Nepal and South Tibet. Scale 1:100,000. Department of Earth Sciences, Oxford University. SEARLE, M. P. & GODIN, L. 2003. The South Tibetan Detachment and the Manaslu Leucogranite: A structural re-interpretation and restoration of the Annapurna - Manaslu Himalaya, Nepal. Journal of Geology, 111, 505-523. SEARLE, M. P. & REX, A. J. 1989. Thermal model for the Zanskar Himalaya. Journal of Metamorphic Geology, 7, 127-134. SEARLE, M. P. & SZULC, A. G. 2005. Channel Flow and ductile extrusion of the high Himalayan slabthe Kangchenjunga--Darjeeling profile, Sikkim Himalaya. Journal of Asian Earth Sciences, 25, 173-185. SEARLE, M. P., WATERS, D. J., REx, D. C. & WILSON, R. N. 1992. Pressure, temperature and time constraints on Himalayan metamorphism from eastern Kashmir and western Zanskar. Journal of Geological Society, London, 149, 753-773. SEARLE, M. P., METCALFE,R. P., REX, A. J. • NORRY, M. J. 1993. Field relations, petrogenesis and emplacement of the Bhagirathi leucogranite, Garhwal Himalaya. In: TRELOAR, P. J. & SEARLE, M. P. (eds) Himalayan Tectonics. Geological Society, London, Special Publications, 74, 429-444. SEARLE, M. P., PARRISH, R. R., HODGES, K. V., HURFORD, A., AYRES, M. W. & WHITEHOUSE, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya: Field Relations, geochemistry, age, origin and emplacement, Journal of Geology, 105, 295-317. SEARLE, M. P., NOBLE, S. R., HURFORD, A. J. & REx, D. C. 1999a. Age of crustal melting, emplacement and exhumation history of the Shivling leucogranite, Garhwal Himalaya. Geological Magazine, 136, 513-525. SEARLE, M. P., WATERS, D. J., DRANSFIELD, M. W., STEPHENSON, B. J., WALKER, C. B., WALKER, J. D. & REX, D. C. 1999b. Thermal and mechanical models for the structural evolution of Zanskar High

378

M.P. SEARLE ET AL.

Himalaya. In: MACNIOCAILL, C. & RYAN, P. D. (eds) Continental Tectonics. Geological Society, London, Special Publications, 164, 139-156. SEARLE, M. P., SIMPSON, R. L., LAW, R. D., WATERS, D. J. & PARRISH, R. R. 2002. Quantifying displacement on the South Tibetan detachment normal fault, Everest massif, and the timing of crustal thickening and uplift in the Himalaya and South Tibet. Journal of Nepal Geological Society, 26, 1-6. SEARLE, M. P., SIMPSON, R. L., LAW, R. D., PARRISH, R. R. & WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal-south Tibet. Journal of the Geological Society, London, 160, 345-366. SEEBER, L. & GORNITZ, V. 1983. River profiles along the Himalayan arc as indicators of active tectonics. Tectonophysics, 92, 335-367. SIMPSON, R. L. 2002. Metamorphism, melting and extension at the top of the High Himalayan slab, Mount Everest region, Nepal and Tibet. DPhil. thesis, Oxford University. SIMPSON, R. L., PARRISH, R. R., SEARLE, M. P. & WATERS, D. J. 2000. Two episodes of monazite crystallization during metamorphism and crustal melting in the Everest region of the Nepalese Himalaya. Geology, 28, 403-406. SMITH, H. A. & BARRIERO, B. 1990. Monazite U-Pb dating of staurolite grade metamorphism in pelitic schists. Contributions to Mineralogy and Petrology, 105, 602-615. STEPHENSON, B. J., WATERS, D. J. & SEARLE, M. P. 2000. Inverted metamorphism and the Main Central Thrust: field relations and thermobarometric constraints from the Kishtwar Window, NW Indian Himalaya. Journal of Metamorphic Geology, 18, 571-590. STEPHENSON, B. J., SEARLE, M. P. & WATERS, D. J. 2001. Structure of the Main Central Thrust zone and extrusion of the High Himalayan crustal wedge, Kishtwar-Zanskar Himalaya. Journal of Geological Society, London, 158, 637-652. VANCE, D. & HARRIS, N. 1999. Timing of prograde metamorphism in the Zanskar Himalaya. Geology, 27, 395-398. VANNAY, J.-C. & GRASEMANN, B. 2001. Himalayan inverted metamorphism and syn-convergence

extension as a consequence of a general shear extrusion. Geological Magazine, 138, 253-276. VANNAY,J.-C. GRASEMANN,B., RAHN, M., FRANK,W., CARTER, A., BAUORAZ, V. & COSCA, M. 2004. Miocene to Holocene exhumation of metamorphic crustal wedges in the NW Himalaya: Evidence for tectonic extrusion coupled to fluvial erosion. Tectonics, 23, TC1014. VISKUPIC, K. & HODGES, K. V. 2001. Monazitexenotime thermochronometry: methodology and an example from the Nepalese Himalaya. Contributions to Mineralogy and Petrology, 141, 233 -247. VISONA, D. & LOMBARDO,B. 2002. Two-mica and tourmaline leucogranites from the Everest-Makalu region (Nepal-Tibet); Himalayan leucogranite genesis by isobaric heating? Lithos, 62, 125-150. WALKER, C. B., SEARLE, M. P. & WATERS, D. J. 2001. An integrated tectonothermal model for the evolution of the High Himalaya in western Zanskar with constraints from thermobarometry and metamorphic modeling. Tectonics, 20, 810-833. WALKER, J. D., MARTIN, M. W., BOWRING, S. A., SEARLE, M. P., WATERS, D. J. • HODGES, K. V. 1999. Metamorphism, melting and extension: Age constraints from the High Himalayan slab of Southeastern Zanskar and Northwestern Lahoul. Journal of Geology, 107, 473-495. WANG, Q., ZHANG, P., FREYMUELLER, J. T. ET AL. 2001. Present-day crustal deformation in China constrained by Global Positioning System measurements. Science, 294, 574-577. WEI, W., UNSWORTH, M., JONES, A. ET AL. 2001. Detection of widespread fluids in the Tibetan crust by magnetotelluric studies. Science, 292, 716-718. WEINBERG, R. F. & SEARLE, M. P. 1999. Volatileassisted intrusion and autometasomatism of leucogranites in the Khumbu Himalaya, Nepal. Journal of Geology, 107, 27-48. Wu, C., NELSON, K. D., WORTMAN, G. ETAL. 1998. Yadong cross-structure and South Tibetan Detachment in the east central Himalaya (89~176 Tectonics, 17, 28-45. ZHAO, W., NELSON, K. D. and Team Project INDEPTH. 1993. Deep seismic reflection evidence for continental underthrusting beneath southern Tibet. Nature, 366, 557-559.

Structural evolution and vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/Nepal: implications for orogen-scale flow partitioning M. J. J E S S U P 1, R. D. L A W l, M. P. S E A R L E 2 & M. S. H U B B A R D 3

1Department of Geosciences, Virginia Tech, Blacksburg, Virginia 24061, USA (e-mail: mjessup @vt. edu) 2Department of Earth Sciences, Oxford University, Oaford, OX1 3PR, UK 3Department of Geology, Kansas State University, Manhattan, Kansas 66506, USA Abstract: The Greater Himalayan Slab (GHS) is composed of a north-dipping anatectic core, bounded above by the South Tibetan detachment system (STDS) and below by the Main Central thrust zone (MCTZ). Assuming simultaneous movement on the MCTZ and STDS, the GHS can be modelled as a southward-extruding wedge or channel. New insights into extrusionrelated flow within the GHS emerge from detailed kinematic and vorticity analyses in the Everest region. At the highest structural levels, mean kinematic vorticity number (Wm) estimates of 0.74-0.91 (c. 45-28% pure shear) were obtained from sheared Tethyan limestone and marble from the Yellow Band on Mount Everest. Underlying amphibolite-facies schists and gneisses, exposed in Rongbuk valley, yield Wm estimates of 0.57-0.85 (c. 62-35% pure shear) and associated microstructures indicate that flow occurred at close to peak metamorphic conditions. Vorticity analysis becomes progressively more problematic as deformation temperatures increase towards the anatectic core. Within the MCTZ, rigid elongate garnet grains yield Wm estimates of 0.63-0.77 (c. 58-44% pure shear). We attribute flow partitioning in the GHS to spatial and temporal variations that resulted in the juxtaposition of amphibolite-facies rocks, which record early stages of extrusion, with greenschist to unmetamorphosed samples that record later stages of exhumation.

The > 2 5 0 0 k m length of the Himalayan orogen is cored by a suite of north-dipping metamorphic rocks (the Greater Himalayan Slab; GHS), that are bounded above and below by the normal-sense South Tibetan detachment system (STDS) and reverse-sense Main Central thrust zone (MCTZ), respectively (Figs 1 & 2). Assuming simultaneous movement along these crustal-scale bounding shear zones (see review by Godin et al. 2006b), the GHS is often modelled as a north-dipping wedge or channel of mid-crustal rocks that was extruded southward from beneath the Tibetan plateau (Fig. 2) beginning in early Miocene time (e.g. Burchfiel & Royden 1985; Burchfiel et al. 1992; Hodges et al. 1992). Although consensus on this general concept of extrusion during crustal convergence exists, and a range of orogen-scale kinematic and thermal-mechanical extrusion models have been proposed, surprisingly little research has focused on quantifying the kinematics (vorticity) of flow within the slab and its potential

causal relationship with progressive exhumation of the GHS. The use of vorticity analysis to quantify flow within sheared rocks has proven to be a useful tool for quantifying the nature and distribution of flow regimes within a range of tectonic settings including contractional (e.g. Simpson & De Paor 1997; Xypolias & Doutsos 2000: Xypolias & Koukouvelas 2001, Xypolias & Kokkalas 2006), extensional (Wells 2001; Bailey & Eyster 2003) and transpressional (Wallis 1995; Klepeis et al. 1999; Holcombe & Little 2001; Bailey et al. 2004) regimes. Vorticity analysis enables estimation of the relative contributions of pure and simple shear, yielding important constraints for GHS extrusion models. Identification of a pure shear component is critically important because such flow would result in: (1) thinning and transport-parallel extension of the slab itself, and (2) an increase in both strain rates and extrusion rates relative to strict simple shear. Attempts to quantify

From: LAW, R. D., SEARLE,M. P. & GODIN,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 379-413. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

380

M.J. JESSUP E T A L .

Fig. 1. Simplified geological map of the Himalaya including the distribution of the main lithotectonic elements of the orogen. The location of the Everest transect and significant geographical locations and areas referred to in the text are indicated. Inset (a) is a digital elevation model displaying the range of elevation for SE Asia (dark grey = low; light grey = high; white = ocean). The Himalaya marks the transition from high elevations of the Tibetan plateau to the lowlands of the Indian plate. MBT, Main Boundary thrust; MCTZ, Main Central thrust zone; MFT, Main Frontal thrust; MMT, Main Mantle thrust; STDS, South Tibetan detachment system; ZSZ, Zanskar shear zone.

flow within the GHS that accommodated this southward extrusion are limited to: (1) a single transect through the lowermost 900 m of the GHS exposed in the Sutlej valley of NW India (Fig. 1; Grasemann et al. 1999; Vannay & Grasemann 2001); (2) preliminary results from the top of the GHS exposed in the Rongbuk valley on the north side of the Everest massif, Tibet (Law et al. 2004); and (3) preliminary results from the middle of the GHS in the Bhutan Himalaya (Carosi et al. 2006). Quantifying and characterizing flow within the GHS is important for development of more realistic models for evolution of the Himalaya, particularly those that propose a synergistic interplay between extrusion, erosion and exhumation (e.g. Beaumont et al. 2001, 2004, 2006; Hodges et al. 2001; Grujic et al. 2002; Jamieson et al. 2004, 2006; Hodges 2006).

The topographic relief of the Everest massif, Tibet/Nepal (Fig. 1), provides a window into mid-crustal processes responsible for extrusion of the GHS, and a particularly appropriate field area to test the various components of extrusion models. In this paper, we combine field-based structural analysis with detailed vorticity analyses of samples from a north-south transect through the GHS in the Everest region using the rigid grain technique of Wallis et al. (1993) and Wallis (1995). Samples collected for vorticity analyses are from a variety of structural and lithologic settings, including high-altitude and summit samples collected during two pioneering climbing expeditions (1933 and 1953) on the north and south sides of Mount Everest, respectively. Rigid grain analysis using elongate garnet porphyroclasts quantify flow along the MCTZ and, by integration

FLOW PARTITIONING IN THE HIMALAYA

Fig. 2. (a-d) Extrusion and channel flow models proposed for the evolution of the Greater Himalayan Slab (GHS). See text for detailed review of each model. GHS, Greater Himalayan Slab; LHS, Lesser Himalayan Sequence; MBT, Main Boundary thrust; MCTZ, Main Central thrust zone; MFT, Main Frontal thrust; STDS, South Tibetan detachment system. with microstructural analysis, constrain the timing of mylonite formation in relation to peak metamorphism (Hubbard 1988, 1989). Field-based structural analysis characterizes deformation in the core of the GHS where deformation temperatures exceed the upper limit for robust rigid grain vorticity analysis. As the first attempt to quantify flow in a transect across the entire GHS, many new insights into vorticity of flow emerge, including an understanding of how flow was partitioned during extrusion and exhumation of the GHS.

Tectonic setting The Himalaya-Tibet orogenic system has accommodated crustal convergence since initiation of collision between India and Asia at c. 55-50 Ma

381

(Searle et al. 1987; Hodges 2000; Yin & Harrison 2000; Figs 1 & 2). The Tibetan plateau encompasses an area of > 5 x 106 km 2 of subdued topography (Fig. 1, inset a), with an average elevation of c. 5000 m (Fielding et al. 1994). To the south stretch the flat, low elevations characteristic of the undeformed internal margin of the Indian plate. Between lies the crest of the Himalaya, which extends for c. 2500 km along-strike, contains the highest elevations in the world (8850 m), and provides exposure of mid-crustal rocks belonging to the GHS (Figs 1 & 2). The GHS forms a 5 30 km thick section of metasedimentary rocks that are intruded by leucogranite dykes and migmatized to varying degrees (Hodges 2000). The age of leucogranite crystallization (c. 23-13 Ma) suggests they were part of a protracted event that marks the culmination of peak metamorphism (Searle 1996; Hodges 2000). The metamorphic evolution of the GHS is often split into two tectonothermal events that may mark distinct thermal pulses or a thermal continuum. Kyanite-grade assemblages are interpreted as relicts of an early event (M1) that U - T h - P b monazite geochronology suggests occurred at c. 3 5 - 3 0 M a (Walker et al. 1999; Simpson et al. 2000). M1 is often overprinted by the pervasive high temperature-low pressure tectonothermal event (M2; c. 23-17 Ma) associated with decompression, migmatization, and emplacement of leucogranite sills (Hodges 2000; Simpson et al. 2000). 4~ thermochronology of muscovite and biotite from leucogranites yields cooling ages that are usually only slightly younger than the U - P b crystallization age of the leucogranites, suggesting rapid decompression following their emplacement (e.g. Hodges et al. (1998) for the Everest transect). Two shear zones bound the GHS: the STDS above and the MCTZ below (Figs 1 & 2). The STDS juxtaposes Tethyan sedimentary rocks of the Tibetan Zone against the GHS, while the MCTZ separates the GHS above from rocks of the Lesser Himalaya Zone (LHZ) below (for detailed review see Godin et al. 2006b).

Extrusion models Despite on-going controversy regarding evidence for simultaneous movement on the MCTZ and STDS (see Godin et aI. 2006b), many researchers continue to view the tectonic evolution of the GHS in the context of models involving southward extrusion of mid-crustal rocks from beneath the Tibetan plateau towards the topographic surface at the plateau margin. All of these models assume simultaneous motion on the upper and lower

382

M.J. JESSUP ETAL.

surfaces of the extruding unit. Two types of models may be distinguished: (1) kinematic models for wedge extrusion (Fig. 2a-c) based on the assumption that the STDS and MCTZ join at depth as, for example, suggested by early interpretations of INDEPTH seismic data (Nelson et al. 1996); (2) more complex coupled t h e r m a l - m e c h a n i c a l finite-element models involving lateral flow of relatively low-viscosity material within a tabular mid-crustal channel in response to a horizontal gradient in lithostatic pressure between the Tibetan plateau and Himalayan foreland (Fig. 2d; Beaumont et al. 2001, 2004, 2006). Kinematic models

In the original wedge extrusion model (Burchfiel & Royden 1985; Royden & Burchfiel 1987; Ktindig 1988, 1989; Burchfiel et al. 1992; Hodges et al. 1993), the extruding wedge developed by gravity-driven collapse in response to the extreme topographic gradient developed along the southern margin of the Himalayan orogen. In this basic conceptual framework, only one main phase of north-south extension, triggered by a fundamental, non-reversible change in the stress state of the orogenic system as a whole (e.g. England & Molnar 1993), occurred in the Himalaya. Synconvergence extrusion, rather than gravitydriven collapse, has been emphasized in more recent models. In the original wedge extrusion model the nature of deformation/flow within the interior portions of the wedge was not explicitly addressed (Fig. 2a), although a broad zone of reverse-sense shearing along the base of the wedge was suggested (e.g. Brunel & Kienast 1986; Hubbard 1988, 1989, 1996; but see also Harrison et al. 1999) as a potential explanation for the long-known inversion of metamorphic isograds adjacent to the MCT (Heim & Gansser 1939). Based on the mapping of antiformally folded isograds in the Zanskar section of the GHS, this was expanded upon in an alternative model by Searle & Rex (1989) who proposed that the entire sequence of rocks contained between the MCT and STDS (Zanskar shear zone) was isoclinally folded during extrusion. More recently proposed kinematic extrusion models are largely based on local evidence for spatial strain path (or vorticity) partitioning within the GHS, and the relative importance of pure shear and simple shear flow components. Model 1

Quantitative evidence for a significant pure shear deformation component associated with flow along the base of the GHS was reported by Grasemann et al. (1999) from the Sutlej Valley

Model 2

(NW India) section of the MCTZ (Fig. 1). Based on correlation between: (1) a downward increase in estimated pure shear component, (2) a downward decrease in deformation temperatures within this zone of inverted isograds, and (3) a high pure shear component indicated by late vein sets, Grasemann et al. (1999) proposed that their vorticity data were most readily interpreted as indicating a temporal (rather than spatial) change in flow regime associated with a decelerating strain path (Simpson & De Paor 1997; Fossen & Tikoff 1997), where simple shear flow at higher temperatures is replaced by pure-shear-dominated flow at lower temperatures. Grasemann et al. (1999) proposed a model for wedge extrusion in which deformation is concentrated towards the boundaries of the wedge and, due to strain compatibility, the centre of the wedge extrudes mainly by pure shearing (Fig. 2b). An important aspect of this model is that the wedge is detached from the footwall and hanging wall. Microstructural and quartz petrofabric data were employed by Grujic et al. (1996) to qualitatively investigate deformation/flow within the lower-central portion of the GHS exposed in Bhutan. These data indicated that at least the lower-central part of the slab had undergone plane strain to weakly constrictional deformation, with flow involving components of both simple shear (reverse or top-to-the-south shear sense) and pure shear. Based on these data, Grujic et al. (1996) proposed a wedge extrusion model in which pervasive shearing occurred throughout the evolving wedge, with opposite shear senses on the top and bottom halves of the wedge (and highest extrusion velocities in the centre) leading, as previously proposed by Searle & Rex (1989) for the Zanskar Himalaya, to antiformal folding of isograds (Fig. 2c; see Godin et al. 2006b). The pure shear component indicated by petrofabric data was not explicitly addressed in this channel flow model, although Grujic et al. (1996) did note that pure shear would lead to thinning and transport-parallel stretching of the wedge/channel during extrusion. Subsequent fieldwork in Bhutan led Grujic et al. (2002) to abandon their wedge-shaped model for the GHS and to regard the GHS as a 10-15 km thick mid-crustal layer, or channel, extending for at least 200 km northward beneath Tibet. In this revised channel flow model (incorporating combined Couette and Poiseuille flow; see review by Grujic 2006) the influence of changing thermal conditions (viscosity) on flow patterns was considered, and qualitative predictions made on the likely influence of changes in boundary conditions and viscosities on domainal variation in flow vorticities Model 3

_.J

LU Z E3 Z

I

rn

LU I-Z

O LU -rILL

O Z I

O I(2 u.i

O _J

O

I

< O (..9

O _! O

LU (.9

...==

F-i

Q.

Z

t-. 0

z

0

o~

Oll:l

z~.

t--!IIl O nr

l-"

~ i-"tli

_

~0> i::

f

0 v--

---+

.

s -r

i

Y

U3

0

elevation (km) above sea level

:,,

E

.~

-8

~/

"J/

a/

/~

I

r o

I---

g

~1 ~ ml,~

~."~!~, +S~200~ (Burkhard 1993; Ferrill et al. 2004). Observed microstructures, and well-defined Rc values in all three of these subgreenschist-facies Tethyan 'limestone' samples indicate that the detrital quartz grains acted as at least semi-rigid clasts rotating in a plastically flowing and dynamically recrystallizing calcite matrix. Wm estimates (Fig. l lb) in these samples range from 0.84 to 0.89 (35-30% pure shear). The pristine grain boundaries of the white mica laths in the limestone suggest that they may have recrystallized during deformation, rather than being of detrital origin (G. Oliver pers comm. 2006). We attribute the lack of a well-developed grain-shape foliation within the calcite matrix, together with the lack of any sedimentary structures, to the operation of grain boundary migration recrystallization, as indicated by the observed microstructures in this sample. Samples of Everest summit 'limestone' have previously been described by Gansser (1964, pp. 164-171) and Sakai et al. (2005). The microstructures described by Gansser (including samples originally described by Gysin & Lombard 1959, 1960) are very similar to those recorded in our samples, except for the presence of crinoid fragments (see also Odell 1965). In contrast, the sample described by Sakai et al. (2005), and collected at c. 6 m beneath the summit (8850 m), contains crinoid, brachiopod and trilobite fragments, and seems to be much less extensively sheared and recrystallized. Microstructurally, the most obvious difference between the summit limestone and the underlying Yellow Band marble is the change in size of recrystallized matrix calcite grains, which abruptly increases from 20-50 Ixm in the Tethyan limestone above the QD to 150-200 Ixm in the Yellow Band marble beneath the detachment. The calcite grains are equant to slightly elongate and define a weak foliation in thin section that is parallel to the strong macroscopic foliation. Larger single calcite grains (400-800 Ixm), together with randomly orientated white mica laths (up to 100 ixm in length) and equant-elongate detrital quartz grains (generally 25-80 Ixm long), are scattered

Everest Series, Yellow Band marble

397

throughout the matrix. Calcite twins are thicker and more closely spaced than in the Tethyan limestone, and both multiple twin sets and tight chevron-style buckling of twin lamellae are commonly developed in the larger calcite grains (particularly in sample GB-25/1 + 2). The presence of thick twins and microstructural evidence for widespread calcite recrystallization involving grain boundary migration indicates deformation temperatures > 250~ (Ferrill et al. 2004). However, the detrital quartz grains exhibit very little undulose extinction, and appear to have acted as semi-rigid porphyroclasts in the flowing calcite matrix, indicating deformation temperatures , 0.4 r . . . . . . . . . . . . c0.3

.I~.HI., . . . . . . ~- Malashan marginR_ MSA29 n = 458 _

6

=

Cuobu margin MSL4 n = 945 a = 37.9

(g)

17.1

0.2

0.1

.......

HI. ,. ~El[

m_. ,--H,H , ] . . . . . . .

.,., ~

n, H ] I _ H , ~ .

,-g~,H,m,=..,.,..

0.5

(d)

Malashan marginMSW14 n = 905 L ~ = 23.2

0.4 0.3

0.0 ......... =~90~

_id ,,.H,H.H3[ _45 o

(h)

Paiku margin MSW22 n = 738 = 30.1

.

0.2 0.1

-

.

.

.

.

.

.

.

.

.

.

m-

.

.

.

.

.

.

.

.

it. ,][H,H,n,-, 0o

45 ~

=~90 ~

.45 ~

0~ 45 ~ (fol) geographic orientation not fixed

A n g l e b i n s (9 ~) Fig. 12. Histograms showing distribution of angles between biotite grains and foliation in granite samples from the Malashan area. Data were taken in planes cut perpendicular to the foliation and parallel to the stretching lineation. The horizontal axes represent 20 angle bins each of 9 ~ width; orientation of foliation is taken as zero degrees. For mylonitic samples of the Malashan granite (a-d), the signs of the angles are defined so that the top-N (NE) shear bands are recorded as minus and the top-S (SW) as plus (see Fig. 1lc). For other less deformed samples without development of shear bands (e-h) orientation was not marked at the time of sampling. The vertical axis represents the lengthnormalized frequency (ratio of measured total biotite length in each bin against that in the whole measurement). For samples (a-d) excesses of bin height against those of oppositely signed bins are shown by black shading to highlight the asymmetry of the angular-data distribution. ( a - d ) Mylonitic samples from the outermost part of the Malashan granite; see Figure 1 la for localities. (e) Weakly foliated sample from central part of the Malashan granite; see Figure l l a for locality. (f, g) Weakly foliated sample from the marginal part of the Cuobu granite; see Figure 3 for localities. (h) Weakly foliated sample from the marginal part of the Paiku granite; see Figure 3 for locality.

.~ >

~ t"-I

~o

.a ~ ~-'

f~ c

N"

d d ~ d d ~ d d d ~

Nm~

M d M ~ d d ~ d ~ ~9

NNe rmN

M d ~ s

II

0

z

i

mN

N N N =~

'NNe

.o

N d ~ d d ~ d d d ~

~ ~ ~o~o~

OoO..

~

488

M. AOYA E T A L .

0.4[[ ~

o

0~

~o

r+

TiO 2 o

+

[

+

~

~, ++

[

+ +-1~ C ~ d -

[

9 +/#,

AI203

31

[] []

Fe203

o + + [] ~

+ ~Ju Tff_~i ~ " -

,4t+

+

+ , - ~ o

_~

L+o " ~ + . . /

++~,+ .~

o.1[ +

16 [ ~ 9

o o+ o , o

++ o

.+.~. ' ~ . ; ~~ ++ c ~

+

++

.r

1

[]

0.4

o

2

"~ _.

o

+

MnO

o

0.8

[]

MgO

CaO

§ oo

0.3

0.6

+

,-,

Oo ~ [] I::l:] r.~ 9

~

+

o

0~

0, 9

++

+4-

01

r

~- o% 0 +.

+

~.# +

[]

[]

--.4.-~.~'_.

+ Na20

++

+

9

+-+r

oo'&'I~r +

+

,

§

9

9 9

5 +

I +

10.3

+++~--~a~.co

P205

+ + +

~

. o.~ + + 0 . +,[]+~o,~ : 1 ~ . ~ , ~ ' 1~o+

~" ++

0.1 + + - - ~ y + + 9

K20

2

0

++

oOo

4 3

++

[]

i

++

6

-o~+

+t~

i

+

+

i~o3;~lC ~

+

+ +o ++

+

. ++ 7

+

+,~.~~~ +

#43

i

+

i

i

+ i

oo 9

o

i

+

+

+ ++

o ~ c o '~ 9

[]

[]

+~

C

i

600

+

. R b (ppm)

:

+

+

++

+

+o

e~ ~+

+13 20(:

200

9

10(: §

+

[]

72

++

9

Ba

9

+

+

+4- 200

7;4

7~5

M a l a s h a n area 9 Malashan granite 9 Cuobu granite 9 Paiku granite

76

0 71

' 72

73

74"

' +o 75

76

(ppm)

ql.[i~ +

+

+

0 71

Others [] Tethys Himalayan granite 1 (Lagoi Kangri,Mabja, Kangmarand Karl La regions)[ + Greater Himalayan leucogranite ] (around Manaslu,Shishapangma& Kula Kangri) /

"72

o

+

0 ['.+ , . ~ t ~

+

o,W~ 73

600

400

~- ~_ ,.,~-++

[] 71

(ppm)

*

+ +o o . &q .~t+. . * + +

[]

Sr

+

30(

+ ~+

73"- 74" . 759

~ ~cEach 1o0m2p o n e n t

wt.%)

Fig. 13. Plots of bulk-chemical data for granites from the Malashan area on SiO2 variation diagrams for 12 elements. Data for the Tethys Himalayan granites from Lagoi Kangri (Gyaco La), Mabja, Kangmar and Kari La regions (Debon et al. 1986; Zhang et al. 2004), and those of the Greater Himalayan leucogranites around Manaslu, Shishapangma and Kula Kangri regions (Dietrich & Gansser 1981; Le Fort 1981; Debon et al. 1986; Guillot & Le Fort 1995; Inger & Harris 1993; Searle et al. 1997; Zhang et al. 2004) are also plotted.

76

MALASHAN GNEISS DOME, SOUTH TIBET granite, we will focus on the Cuobu and Paiku granites in this section. The first indicator of the nature of the Cuobu granite is given by the observation that it contains dykes originating from the main body (Fig. 14a). A second important observation is that the pelitic schists adjacent to the Cuobu granite contain andalusite (Fig. 3). The andalusite generally replaces biotite (Fig. 14b) and locally staurolite, indicating

489

that andalusite grew significantly later than the main phase of Barrovian-type metamorphism. Thirdly, formation of skarn is observed in one area adjacent to the Cuobu granite (Fig. 14c & d). The andalusite and skarn are best explained as the products of contact metamorphism. The above three features, therefore, indicate that the Cuobu granite is an intrusive body. An important question for our study is the timing of its emplacement

Fig. 14. Geological features of the Cuobu and Paiku granite. And, andalusite; Bt, biotite; Cpx, clinopyroxene; Grt, garnet; P1, plagioclase; Qtz, quartz. (a) Field photo showing dykes (arrowed) originating from the Cuobu granite. (b) Photomicrograph showing occurrence of andalusite in a pelitic schist adjacent to the Cuobu granite (crossed nicols); andalusite replaces biotite developed along $2. (c, d) Outcrop photo (C) and slab photo (d) of skarn found in the region adjacent to the Cuobu granite. (e) Photomicrograph showing occurrence of slightly boudinaged andalusite in pelitic schist adjacent to the Paiku granite (crossed nicols). (f) Slab photo of sample from marginal part of the Paiku granite (MSW22); this is the only Paiku sample that can be used for angle measurement of biotite grains (Fig. 12h) because of the general low biotite content. The white vein in the left side cross-cuts Sa but is pulled apart in the direction of $2, as suggested by the muscovite grains on both edges of the vein. We consider the syn-D2 nature of the vein to justify plotting this vein sample (MSW22V; Table 1) as part of the Paiku granite (Fig. 13).

490

M. AOYA E T AL.

relative to deformation. The observations that the dykes cross-cut S 2 (Fig. 14a) and that andalusite overgrows $2 (Fig. 14b) indicate that emplacement post-dated at least a part of D2 deformation. In addition, garnet-rich layers in the skam are slightly boudinaged in a north-south direction (Fig. 14d) indicating that skarn formation and granite emplacement pre-dated at least a part of D2. These two observations lead to the conclusion that emplacement of the Cuobu granite occurred during D2 deformation. Although large-scale dykes and skarn were not observed around the Paiku granite, it is also associated with formation of andalusite that post-dates the garnet-staurolite-biotite assemblage in the adjacent pelitic schists (Figs 3 & 14e). This suggests that the Paiku granite also intrudes the surrounding schists. In this case the andalusite is locally slightly pulled apart within $2 (Fig. 14e) suggesting that intrusion of the Paiku granite also took place during the D2 phase. To compare the strength of the D 2 deformation experienced by the three granite bodies, the orientation distribution of biotite was also measured for samples from the Cuobu and Paiku granites (Fig. 12f-h). The three samples are taken from the marginal parts of the two granites (see Fig. 3 for localities) to enable comparison with the marginal samples of the Malashan granite (Fig. 12a-d). For the Paiku granite, only one sample (MSW22; Fig. 14f) was measurable because of the generally low biotite content. Comparison of Figure 12a-d with Figure 12f-h shows that the Paiku and Cuobu granites are significantly less deformed than the Malashan granite. The strength of deformation, or deviation from random distribution

0.5

I,-:".."'~.."'.::'.:.".." 9. . . . . . . . . . . 1 --

~i

Summary:

in the Malashan

of granites area

As already mentioned, the Paiku granite is chemically equivalent to the Greater Himalayan

Period affected by D 1

9

-9v: 5 .~-" : -"-'" : -""5: -:''~l ] h ~ ~ .I:..-'." "..:..'.:.'..:.'..:..'.:..'.:

+' 04 o o r.o

origin

~

~9.". -.'."-""."':-.".'.'1 ".."...'.:v'::..'.".'l 9" . . . . . . ' . . . . . . .

(thick lines in Fig. 12), is reflected in two indicators: (i) the concentration of data in the central bin (0 ___4.5~ and (ii) the standard deviation (o-) for a particular set of data. As the deformation becomes stronger the concentration in the central bin becomes greater and the standard deviation becomes smaller. This systematic trend is represented in Figure 15a with the Malashan granite showing stronger deformation than the Paiku and Cuobu granites. The Malashan and Cuobu granites, and even the measured Paiku sample (MSW22), all have similar bulk-chemical compositions (Table 1; Fig. 13). In addition, all three granite bodies are commonly surrounded by calc-schist (Fig. 3). Therefore, the strain contrast shown in Figure 15a cannot be explained by differences in rheological properties at the pluton contacts. The most likely explanation is that there were differences in the timing of emplacement of the granite bodies with respect to the D2 deformation. The Cuobu granite was intruded during the later stage of De and experienced only a short period of D2 deformation (Fig. 15b). In contrast, formation of the Malashan granite was earlier, and this body was more strongly affected by D 2 (Fig. 15b). The intermediate strain obtained from the Paiku sample suggests that intrusion of the Paiku granite took place during D2 but at a time earlier than intrusion of the Cuobu granite (Fig. 15b).

~

_

Period affected by D2

Malashan

-

n ! ~ ~,~,2 "~iaiasha (margin) "..".'.1 \ ~ ~ . . . . '"" ..... "" " "i:'."i ~~ . -,~_ ~a,asnan ..',. :',':..'.'..':'..-':'..-7,'."~ ",',s~ (centre)

or"

(margin) P ~

~ ' ~

"t L a!h

IHg'xfU Cuobu ~. ..: ,-.,.:.~

0.2

(margin) ~ ql) 0

(a)

I

I

15

20

I

25

I

30

standard deviation (~

~

Mal an ntrusion

Paiku

Cuobu

t Paiku intrusion

time

Cuobu intrusion

I

35

40

(b)

Interpretation

Fig. 15. (a) Summary of the angle measurements for biotite grains shown in Figure 12. (b) Interpretation of the strain contrast between the Malashan, Cuobu and Paiku granites as shown in (a). The strain contrast is interpreted to reflect their different timings of emplacement with reference to D2 deformation.

MALASHAN GNEISS DOME, SOUTH TIBET leucogranites (Fig. 13), which have crystallization ages of 32-12 Ma. Assuming that the Paiku granite is also Himalayan, syn-D2 intrusion of the Paiku granite (Fig. 15b) leads to the interpretation that D2 is a Himalayan deformation. The Cuobu granite can also be regarded as Himalayan granite because its intrusion post-dates that of the Paiku granite (Fig. 15b). In addition, the striking bulkchemical similarity between the Cuobu and Malashan granites (Fig. 13) suggests the Malashan granite is also a Himalayan intrusion. This consideration leads us in turn to the conclusion that the D1 event is also a Himalayan deformation, because D1 affects the Malashan granite, which we conclude to be a Himalayan intrusion (Fig. l lc & d). This conclusion is supported by SHRIMP spot dating of zircons from the Malashan and Cuobu granites (Aoya et al. 2005). We also suggest that intrusion of the Malashan granite occurred during the D1 phase (Fig. 15b). This is based on the syn-Di growth of asymmetric K-feldspar (Fig. 1 ld), associated high-temperature deformation (>500~ and magmatic growth of plagioclase inclusions (Fig. l le), which are suggestive of a partially molten stage: the Malashan granite experienced the D1 deformation when it was crystal mush. Our main conclusion is that the Malashan granite, which has many similarities with the Kangmar granite, can be interpreted as an intrusive body formed during the Himalayan orogeny.

Discussion General features o f the North Himalayan gneiss dome With the main exception of the association of two weakly deformed granites (Cuobu and Paiku granites), the Malashan metamorphic complex shows a number of striking similarities with the Kangmar dome: (i) presence of Barrovian-type metamorphism whose grade increases towards the granite bodies located in the central part of the metamorphic regions (referred to as granite cores hereafter; Figs 2 & 3); (ii) presence of two major deformation stages, DI and D2, in metasedimentary schists and increasing strength of the D2 deformation towards the granite cores (Figs 2 & 5); (iii) development of strong D2 foliation (gneissosity) in the outermost parts of the granite cores (Fig. 12a-e); and (iv) a roughly north-south flow direction associated with D2 deformation, as suggested by distribution of L2 (Fig. 6). In the case of the Malashan area, we consider the Malashan granite to be a granite core similar to the Kangmar granite, because it contains D2 shear zones in its outermost part (Fig. 1 l a - c ) , forming not only a metamorphic but also a

491

deformational core to the surrounding metasedimentary units. In other words, the Malashan complex is a North Himalayan gneiss dome cored by the Malashan granite. We propose that the association of features (i)-(iii) can be regarded as a definition for the North Himalayan gneiss domes. We do not interpret the Cuobu and Paiku granites as cores to the Malashan dome, because: (i) their emplacement is documented to be significantly later than the main Barrovian-type metamorphism (Fig. 14b) whose distribution defines the granite core; and (ii) they show only minor effects of the D2 deformation (Fig. 12f-h), forming a significant strain gap at their contacts. That is, they form neither a metamorphic nor a deformational core to the surrounding metasedimentary schists. In the case of the two examples discussed in this study, and also the recently reported example of the Mabja dome (Lee et al. 2004; Fig. 1), the flow direction of D2 was similarly orientated north-south (Fig. 6). There is, however, considerable local variation in the orientation of this lineation and it may differ from one region to another. It remains to be seen to what extent this is a general feature of the North Himalayan gneiss domes.

Contradictory interpretations o f the North Himalayan gneiss domes The combination of structural and geochemical studies presented in this contribution leads to the conclusion that D2 deformation in the Malashan dome took place during the Himalayan orogeny. This is also the case for the Kangmar and Mabja domes where geochronologic and thermochronologic studies indicate a late Oligocene to Miocene age for D2 deformation (Lee et al. 2000, 2006). However, as mentioned above, there is disagreement concerning the kinematics of the D2 deformation in Kangmar; i.e. dominantly top-tothe-north shear (Chen et al. 1990) or bulk pure shear with a combination of top-to-the-north and top-to-the-south on the north and south flanks of the dome respectively (Lee et al. 2000). Our kinematic study in Malashan indicates a top-tothe-north sense of shear, directly comparable with the results of Chen et al. (1990) from Kangmar. In our study we have determined the map-scale distribution of D2 shear senses and then assessed the reliability of each shear-sense determination by documenting whether it was made by two or more independent methods (Fig. 9). In the case of the Kangmar dome it is at present difficult to conclude which of the present views (Chen et al. 1990; Lee et al. 2000) is correct because the map-surface distribution of the determined shear senses has yet to be published. It is therefore necessary to examine the kinematic nature of the D2 deformation in the

492

M. AOYA ETAL.

Kangmar dome, and in other still enigmatic gneiss domes to assess whether the D2 deformation can kinematically be correlated with activity on the STD at the scale of the whole Tethys Himalaya. Another contradiction revealed in the present study concerns the origin of the granite cores. We interpret the Malashan granite, the core of the Malashan dome, as an intrusive body formed during the Himalayan orogeny. Our conclusion is based on the presence of three granite bodies that show variations in their bulk composition and in development of D2 deformational fabrics, and is supported by radiometric dating (Aoya et al. 2005). It is the variation in granitic bodies of the Malashan area that allows us to draw this conclusion. In this sense, for some types of study, the Malashan area has clear advantages in its geological setting over the Kangmar area. We suggest, therefore, that the Kangmar granite may also be a Himalayan intrusive based on its striking similarities with the Malashan granite shown in this study. Even if we exclude from our discussion the inference that the Paiku granite is Himalayan, the evidence demonstrating that the Cuobu and Paiku granites are syn-D2 intrusions (Fig. 15b) remains unchanged. Because D2 is well constrained to be a Himalayan deformation in the Kangmar and Mabja areas, and the Malashan granite is chemically equivalent to the Cuobu granite (Fig. 13), it is difficult to escape the conclusion that the Malashan granite is a Himalayan intrusion. This leaves a major geochronological problem: how to reconcile our proposal with the results of conventional U - P b zircon dating of the Kangmar granite, which yields ages of 562 +_ 4 Ma (Sch~irer et al. 1986) and 508 ___ 1 Ma (Lee et al. 2000). The major difference of c. 50 Ma between the two studies greatly exceeds the estimated errors. One possibility is that the U - P b ages obtained using mineral separates of zircon from the Kangmar granite do not represent a single magmatic event but complexly mixed results reflecting most common ages of xenocrysts. This explanation can account for the large range in estimated ages for the Kangmar granite and is compatible with our present study. If the North Himalayan gneiss domes are confirmed to be dominantly cored by intrusions formed during the Himalayan orogeny, at least a part of the concentrically distributed metamorphic isograds surrounding the granite cores may be explained as the result of deep-seated contact metamorphism. Implications for channel flow models

The North Himalayan gneiss domes have recently been viewed as windows into the Greater Himalayan sequence exposed on hinges of a large-scale antiform by several researchers (e.g. Beaumont et al.

2001, 2004; Searle et al. 2003; Zhang et al. 2004), and the strongest basis for this idea is the above-mentioned Cambrian to Ordovician U - P b ages for the Kangmar granite (Scharer et al. 1986; Lee et al. 2000). In this interpretation the ductile D2 shear zones developed around the marginal part of the granite cores themselves are regarded as a direct continuation of the STD. The stratigraphic position of the Malashan area in the Tethys Himalaya (Fig. 10e), however, suggests that the Malashan granite, the core of the Malashan dome, is unlikely to be a part of the Greater Himalaya. The Greater Himalayan sequence has ages up to Ordovician (c. 480 Ma; DeCelles et al. 2000), while the Tethys Himalaya preserves a sedimentary succession from Ordovician to Eocene (Liu & Einsele 1994). To propose that the Malashan granite is part of the Greater Himalayan sequence would therefore imply that an unreasonably large amount of the succession, everything between the Jurassic and Ordovician, had been locally tectonically excised from within the structural pile. In this study we suggest an alternative interpretation where the Malashan granite is an intrusive body that formed during the Himalayan orogeny. This idea is compatible with the stratigraphic position of the Malashan area (Fig. 10e) and is supported by SHRIMP spot dating of zircons from the Malashan and Cuobu granites (Aoya et al. 2005). Himalayan formation ages of 28-8 Ma have also been obtained for several nearly undeformed granites in the Tethys Himalaya (Sch~irer et al. 1986; Harrison et al. 1997; Zhang et al. 2004), and the age range partly overlaps that of the Greater Himalayan granites (32-12 Ma). Zhang et al. (2004) further show that largely undeformed granites from the Mabja and Lhagoi Kangri regions and even the gneissose Kangmar granite have N d - S r isotopic compositions similar to the Greater Himalayan leucogranites and metasediments. These studies suggest that the variously deformed Tethys Himalayan granites were derived from melting of the Greater Himalayan sequence. The original magma of the Malashan granite, therefore, probably originated in the Greater Himalaya, migrated upward through the lower Tethys Himalaya, and was then emplaced into the Jurassic sediments located in the middle- to upper-stratigraphic section of the Tethys Himalaya (Fig. 10e). This implies that the D2 shear zone observed around the Malashan granite is not directly comparable to the STD, but represents a deformation that occurred within the Tethys Himalaya, at structurally higher levels than the STD. The two main conclusions of this study, (i) the origin of the Malashan granite as Himalayan intrusive and (ii) correlation of the D2 deformation with activity of the STD, are strongly supported by

MALASHAN GNEISS DOME, SOUTH TIBET radiometric age determinations reported by Aoya et al. (2005). These age determinations for formation of granite bodies and for the D2 deformation have important implications for channel flow models of the region (e.g. Beaumont et al. 2001, 2004; Jamieson et al. 2004), which are strongly dependent on the strength of the upper crust. As already discussed, the D2 deformation in Malashan occurred within the Tethys Himalaya (Fig. 10), which in the channel flow models behaves as the rigid hanging wall to a channel of middle- to lower-crustal flow. If, as we suggest, the D2 deformation of the Malashan area occurred simultaneously with the activity of the STD, it indicates that the top-to-the-north deformation represented by the STD also affected the hanging wall and, therefore, that the hanging wall was at least locally weak during extrusive flow of the Greater Himalaya.

Conclusions Considering the Malashan granite to be a granite core, the Malashan metamorphic complex shows the following similarities with the Kangmar dome: (i) presence of Barrovian-type metamorphism with grade increasing towards a granite core (Figs 2 & 3); (ii) development of two major ductile deformation stages, D1 and D2, and increasing strength of the D 2 deformation towards the granite core (Figs 2 & 5); (iii) development of a D2 shear zone in the outermost part of the granite core (Fig. 12a-e); and (iv) roughly north-south flow direction during D2 deformation as suggested by the distribution of L2 (Fig. 6). The Malashan complex can therefore be regarded as a North Himalayan gneiss dome similar to the Kangmar dome. We also propose that features (i)(iii) can be used to define North Himalayan gneiss domes. Other significant features of the Malashan dome are: (i) D2 deformation is associated with a dominantly top-to-the-north sense of shear, which contradicts the most recent view on kinematics of D2 in the Kangmar dome (Lee et al. 2000); (ii) deposition age of the associated metasediments is upper Jurassic suggesting that the Malashan dome is located not on the base, but in the middle of the Tethys Himalaya; and (iii) none of the associated granites is a representative of the Indian basement as proposed for the Kangmar granite (Chert et al. 1990; Lee et al. 2000); instead all formed as intrusive bodies during the Himalayan orogeny. These results suggest that the formation mechanism of the North Himalayan gneiss domes needs to be re-evaluated, including tests of the rigidity of the hanging wall proposed in the channel flow models (e.g. Beaumont et al. 2001, 2004; Jamieson et al. 2004).

493

We would like to thank T. Argles, J.-P. Burg and R. Law, whose comments helped improve this study. R. Law is also acknowledged for his editorial assistance. This research was supported by a JSPS grant-in-aid awarded to S. Wallis and a grant from Central Washington University awarded to J. Lee.

References ALSDORF, D. & NELSON, D. 1999. Tibetan satellite magnetic low; evidence for widespread melt in the Tibetan crest? Geology, 27, 943-946. AOYA, M., WALLIS, S. R., TERADA, K., LEE, J., KAWAKAMI, Z., WANG, Y. & HEIZLER, M. 2005. North-south extension in the Tibetan crust triggered by granite emplacement. Geology, 33, 853- 856. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON, R. A., NGUYEN,M. H. & MEDVEDEV, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06406. BRODIE, K. H. & RUTTER, E. H. 1985. On the relationship between deformation and metamorphism, with special reference to the behavior of basaltic rocks. In: THOMPSON, A. B. & RUBIE, D. C. (eds) Metamorphic Reactions: Kinetic Textures and Deformation. Springer-Verlag, New York, 138-179. BURCHFIEL, B. C., CHEN, Z., HODGES, K. V., LIU, Y., ROYDEN, L. H., DENG, C. & Xu, J. 1992. The South Tibetan detachment system, Himalayan orogen: Extension contemporaneous with and parallel to shortening in a collisional mountain belt. Geological Society of America Special Paper, 269. BURG, J.-P. & CHEN, G. M. 1984. Tectonics and structural zonation of southern Tibet, China. Nature, 311,219-223. BURG, J.-P., BRUNEL,M., GAPAIS, D., CHEN, G. M. & LIu, G. H. 1984a. Deformation of leucogranites of the crystalline Main Central Sheet in southern Tibet (China). Journal of Structural Geology, 6, 535-542. BURG, J.-P., GUIRAUD,M., CHEN, G. M. & LI, G. C. 1984b. Himalayan metamorphism and deformations in the North Himalayan Belt (southern Tibet, China). Earth and Planetary Science Letters, 69, 391-400. BURG, J.-P., LEYRELOUP,A., GIRARDEAU,J. & CHEN, G. M. 1987. Structure and metamorphism of a tectonically thickened continental crust: the Yalu Tsangpo suture zone (Tibet). Philosophical Transactions of the Royal Society of London, A321, 67-86. CHEN, L., BOOKER, J. R., JONES, A. G., NONG, W., UNSWORTH, M. J., WENBO, W. & HANDONG, T. 1996. Electrically conductive crust in southern Tibet from INDEPTH magnetotelluric surveying. Science, 274, 1694-1696. CHEN, Z., LIU, Y., HODGES,K. V., BURCHFIEL,B. C., ROYDEN, L. H. & DENG, C. 1990. The Kangmar

494

M. AOYA ETAL.

dome: a metamorphic core complex in southern Xizang (Tibet). Science, 250, 1552-1556. DEBON, F., LEFORT, P., SHEPPARD, S. M. F. & SONET, J. 1986. The four plutonic belts of the TranshimalayaHimalaya: a chemical, mineralogical, isotopic, and chronological synthesis along a Tibet-Nepal section. Journal of Petrology, 27, 219-250. DECELLES, P. G., GEHRELS,G. E., QUADE,J., LAREAU, B. • SPURLIN, M. 2000. Tectonic implication of U-Pb zircon ages of the Himalayan orogenic belt in Nepal. Science, 288, 497-499. DIETRICH, V. & GANSSER, A. 1981. The leucogranites of the Bhutan Himalaya (crustal anatexis versus mantle melting). Schweizerische Mineralogische und Petrographische Mitteilungen, 61, 177-202. DUGUET, M. & FAURE, M. 2004. Granitoid emplacement during a thrusting event: structural analysis, microstructure and quartz c-axis patterns. An example from Hercynian plutons in the French Massif Central. Journal of Structural Geology, 26, 927-945. EDWARDS, M. A. & HARRISON, T. M. 1997. When did the roof collapse? Late Miocene north-south extension in the high Himalaya revealed by Th-Pb monazite dating of the Khula Kangri granite. Geology, 25, 543-546. FITZ GERALD, J. D. & STI2NITZ, H. 1993. Deformation of granitoids at low metamorphic grade I: reactions and grain size reduction. Tectonophysics, 221, 299-324. GUILLOT, S. & LE FORT, P. 1995. Geochemical constraints on the bimodal origin of High Himalayan leucogranites. Lithos, 35, 221-234. HARRISON, T. M., LOVERA, O. M. & GROVE, M. 1997. New insights into the origin of two contrasting Himalayan granite belts. Geology, 25, 899-902. HARRISON, T. M., GROVE, M., MCKEEGAN, K. D., COATH, C. D., LOVERA, O. M. 8z LE FORT, P. 1999. Origin and episodic emplacement of the Manaslu intrusive complex, central Himalaya. Journal of Petrology, 40, 3-19. HIROI, Y., MOTOYOSHI,Y., SHIRAISHI,K. & ELLIS, D. J. 1995. The significance of euhedral calcic plagioclase inclusions in garnet from the Liitzow-Holm Complex, East Antarctica: A textural indicator of partial melting in pelitic gneisses. Proceedings of the NIPR Symposium on Antarctic Geosciences, 8, 107-120. HODGES, K. V. 2000. Tectonics of the Himalaya and southern Tibet from two perspectives. Geological Society of America Bulletin, 112, 324-350. HODGES, K. V., PARRISH, R. R., HOUSH, T. B., LUX, D. R., BURCHEIEL,B. C., ROYDEN, L. H. & CHEN, Z. 1992. Simultaneous Miocene extension and shortening in the Himalayan orogen. Science, 258, 1466-1470. HODGES, K. V., PARRISH, R. R. & SEARLE, M. P. 1996. Tectonic evolution of the central Annapurna Range, Nepalese Himalayas. Tectonics, 15, 1264-1291. INGER, S. 8z HARRIS, N. 1993. Geochemical constraints on leucogranite magmatism in the Langtang Valley, Nepal Himalaya. Journal of Petrology, 34, 345-368.

JAMIESON, R. A., BEAUMONT, C., MEDVEDEV, S. & NGUYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06407. JIN, Y., MCNUTT, M. K. & ZHU, Y. 1994. Evidence from gravity and topography data for folding of Tibet. Nature, 371, 669-674. LAW, R. D., SEARLE, M. P. & SIMPSON, R. L. 2004. Strain, deformation temperatures and vorticity of flow at the top of the Greater Himalayan Slab, Everest Massif, Tibet. Journal of Geological Society, London, 161, 305-320. LE FORT, P. 1981. Manaslu leucogranite: a collision signature of the Himalaya, a model for its genesis and emplacement. Journal of Geophysical Research, 86, 10545-10568. LEE, J., HACKER, B. R., DINKLAGE,W. S. ETAL. 2000. Evolution of the Kangmar Dome, southern Tibet: Structural, petrologic, and thermochronologic constraints. Tectonics, 19, 872-895. LEE, J., DINKLAGE, W. S., WANG, Y. & WAN, J.-L. 2002. Geologic map of Kangmar dome, southern Tibet. Map and Chart Series, Geological Society of America, Report MCH090. LEE, J., HACKER,B. R. & WANG, Y. 2004. Evolution of north Himalayan gneiss domes: Structural and metamorphic studies in Mabja dome, southern Tibet. Journal of Structural Geology, 26, 2297-2316. LEE, J., MCCLELLAND, W., WANG, Y., BLYTHE, A. & MCWILLIAMS, M. 2006. Oligocene-Miocene middle crustal flow in southern Tibet: Geochronology of Mabja dome. In: LAW, R. O., SEARLE,M. P. & GODIN, L. (eds) Channel Flow, Ductile Extru-

sion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 445-469. LIu, G. & EINSELE, G. 1994. Sedimentary history of the Tethyan basin in the Tibetan Himalayas. Geologische Rundschau, 83, 32-61. MAKOVSKY, Y., KLEMPERER, S. L., HUANG, L. & LU, D. 1996. Structural elements of the southern Tethyan Himalaya crust from wide-angle seismic data. Tectonics, 15, 997-1005. MEANS, W. D. 1981. The concept of steady-state foliation. Tectonophysics, 78, 179-199. NELSON, K. D., ZHAO, W., BROWN, L. D. ETAL. 1996. Partially molten middle crust beneath southern Tibet: Synthesis of project INDEPTH results. Nature, 274, 1684-1688. PASSCHIER, C. W. 1990. Reconstruction of deformation and flow parameters from deformed vein sets. Tectonophysics, 180, 185-199. PASSCHIER, C. W. & TROUW, R. A. J. 1996. Microtectonics. Springer-Verlag, Berlin. PLATT, J. P. 1980. Extensional structures in anisotoropic rocks. Journal of Structural Geology, 2, 397-410. SCHARER, U. 1984. The effect of initial 23~ disequilibrium on young U-Pb ages: the Makalu case, Himalaya. Earth and Planetary Science Letters, 67, 191-204. SCHARER, U., Xu, R.-H. & ALLI~GRE, C. J. 1986. U-(Th)-Pb systematics and ages of Himalayan

MALASHAN GNEISS DOME, SOUTH TIBET leucogranites, South Tibet. Earth and Planetary Science Letters, 77, 35-48. SEARLE, M. P., PARRISH, R. R., HODGES, K. V., HURFORD, A., AYRES, M. W. & WHITEHOUSE, M. J. 1997. Shisha Pangma leucogranite, South Tibetan Himalaya: Field relations, geochemistry, age, origin and emplacement. Journal of Geology, 105, 295-317. SEARLE, M. P., SIMPSON, R. L., LAW, R. D., PARRISH, R. R. • WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal-south Tibet. Journal of Geological Society, London, 160, 345-366. SIMPSON, R. L., PARRISH, R. R., SEARLE, M. P. 8,= WATERS, D. J. 2000. Two episodes of monazite crystallization during metamorphism and crustal melting in the Everest region of the Nepalese Himalaya. Geology, 28, 403-406.

495

SPATH, L. F. 1923. The ammonites of the Shales-withBeef. Quarterly Journal of the Geological Society, London, 70, 66-88. SPRY, A. 1969. Metamorphic Textures. Pergamon, London. UHLIG, V. 1903-1910. The fauna of the Spiti shales. Palaeontologia Indica, Ser. XV, 4, fasc. 1-3. WALLIS, S. R. 1992. Vorticity analysis in a metachert from the Sanbagawa Belt, SW Japan. Journal of Structural Geology, 14, 271-280. WESTERMANN, G. E. G. & WAN6, Y.-G. 1988. Middle Jurassic ammonites of Tibet and the age of the Lower Spiti Shales. Palaeontology, 31(2), 295 -339. ZHANG, H., HARRIS, N., PARRISH, R. e r AL. 2004. Causes and consequences of protracted melting of the mid-crust exposed in the North Himalayan antiform. Earth and Planetary Science Letters, 228, 195-212.

Heterogeneous ductile deformation along a mid-crustal extruding shear zone: an example from the External Hellenides (Greece) P. X Y P O L I A S & S. K O K K A L A S

D e p a r t m e n t o f Geology, University o f Patras, GR-26500, Patras, Greece (e-mail: p.xypolias @ upatras, gr) Abstract: Petrofabric, finite strain and kinematic vorticity data were used to investigate the het-

erogeneous nature of ductile deformation along a 1.5-2 km thick extruding shear zone in the south Peloponnese, that formed under blueschist-facies conditions. Asymmetric quartz c-axis fabrics confirm westward thrust movements on an east-dipping shear zone and provide evidence for localized top-down-to-the-east shear sense at the front of the zone. Strain ratio (Rxz) is nearly constant (c. 3.0-4.0) along the upper structural levels of the zone but increases systematically from the middle to the bottom, approaching a value of c. 9.5 in the frontal parts, close to the basal thrust, and a value of c. 7.0 in the inner parts. The distribution of kinematic vorticity number depicts a simple-shear-dominated domain in the lower half of the shear zone and shows that the pure shear component always increases upwards in the zone, becoming dominant at the top of the inner parts of the zone. Integration of strain and vorticity data yields a shear-parallel elongation of c. 60-90% at the top and c. 40-60% at the bottom of the zone, revealing that both upper and lower surfaces of the extruding slices were 'stretching faults'. Minimum total displacements of 25 and 41 km and slip rates of 6.5 and 10 mm/year were estimated for the basal and roof faults, respectively.

Recent studies in collisional orogenic belts such as the Himalaya (e.g. Hodges et al. 1993; Grujic et aL 1996; Grasemann et al. 1999; Searle et al. 2003), the Alps (Escher & Beaumont 1997; Bucher et al. 2003), the Canadian Cordillera (Johnston et al. 2000) the Neoproterozoic Belt of Brazil (Campos Neto & Caby 2000) and the Hellenides (e.g. Xypolias et al. 2003) suggest that upward solid-state extrusion is an effective mechanism for bringing rocks which have undergone deep subduction and high-pressure (HP) metamorphism, near to the ground surface. In all these natural examples, the extruding slices of HP rocks occupy low-angle zones, which are locally sandwiched between rigid crustal blocks. These slices are commonly bounded by a subduction-related thrust fault at the base and a fault with a normal shear sense at the top, that operated contemporaneously with a compressional tectonic regime. The deformation of an extruding slice is likely to be heterogeneous and is characterized by a complex distribution of strain and vorticity of flow. Previous studies have indicated either localization of high strain along the two walls of the shear zone with an intervening low-strain and pure-shear-dominated domain (e.g. Grasemann et al. 1999), or a pervasive deformation of the material throughout the shear zone which is characterized by a down-section increase in both strain magnitude and simple shear component of deformation (e.g. Escher & Beaumont 1997; Xypolias & Koukouvelas 2001).

Although buoyancy forces assist the upward extrusion of HP rocks along crustal-scale shear zones (Ernst 2001), the pure shear component of deformation seems to play a more crucial role in this mechanism (Vannay & Grasemann 2001). For isochoric plane strain, the pure shear component of ductile flow induces a thinning normal to the boundaries of a low-angle shear zone and therefore requires that some elongation of the material also occurs (Wallis 1995). Given that the shear zone has non-parallel downward-closing sides, the elongation of the material is directed towards the Earth's surface, contributing to the exhumation/ extrusion of rocks. In this study, we present new petrofabric, finite strain and kinematic vorticity data from a thin crustal-scale extruding zone in the External Hellenides orogenic belt (Fig. la), which formed under blueschist-facies conditions. Previous quantitative studies at the northern lateral edge of the zone, where less deep-seated rocks are exposed, have shown that the nature of deformation is strongly heterogeneous (e.g. Xypolias & Koukouvelas 2001). The present study aims to quantify the ductile deformation along this zone using data from two areas occupying the inner and the frontal parts of the extruding zone. Based on these results, we have tried to estimate the contribution of internal deformation to the total displacement field.

From: LAW, R. D., SEARLE,M. P. & GODIN,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 497-516. 0305-8719/06/$15.00 9 The Geological Society of London 2006.

498

P. XYPOLIAS & S. KOKKALAS

Fig. 1. (a) Simplified geological map showing the position of the External Hellenides in relation to the Dinaric-Tauric arc. (1-3) Evolutionary profiles from the Peloponnese indicating the orgenic evolution of the SW Hellenides and the exhumation mechanism of HP rocks (after Xypolias & Doutsos 2000; Doutsos et al. 2000). Inset: generalized map of the Alpine chain in southwestern Europe (modified after Coward & Dietrich ! 989). (b) Geological map of the southern Peloponnese showing the Taygetos and Parnon tectonic windows. The locations of the maps in Figure 2 and 3 are indicated by boxes.

Geotectonic framework The External Hellenides are part of the Alpine Orogenic belt (Fig. la, inset) and form an orocline connecting the Dinarides to the N W with the Taurides to the SE (Fig. I a). They mainly consist of Mesozoic and Cenozoic sedimentary rocks that were deposited on the rifted eastern margin of the Apulia microcontinent. The present disposition of the thrust sheets of the External Hellenides resulted from the eastward subduction and collision between Apulia and Pelagonian (Internal Hellenides)

microcontinents (Robertson e t al. 1991; Doutsos e t al. 1993).

The metamorphic rocks of Phyllite-Quartzite (PQ) and Plattenkalk (PLK) units belong to the External Hellenides and constitute a HP-LT belt, which extends over a distance of 600 km from the northern Peloponnese, through Kithira and Crete, to the Dodecanese Islands (Fig. l a). The PQ unit, which is the body of the studied extruding shear zone, is tectonically intercalated between the non-metamorphic cover thrust sheets (Pindos and Tripolitza units) and the lowermost structural

DEFORMATION AND DUCTILE EXTRUSION unit of the External Hellenides, the PLK unit (Fig. 1). The Pindos and Tripolitza units are mainly composed of Triassic to Eocene carbonate rocks and an upper Eocene flysch. The Pindos unit tectonically rests on the Tripolitza unit and both units have a combined structural thickness no greater than 6 km. At the base of Tripolitza carbonate rocks, a thin Permo-Triassic sequence, referred to as the Tyros Beds, which is in tectonic contact with the underlying PQ unit, have suffered very low-grade metamorphism (200-350~ 3 - 6 kbar; e.g. Thi~boult & Triboulet 1984). The PQ unit is predominantly composed of phyllites, quartzites and metaconglomerates, with local marble intercalations and small lenses of basic metavolcanic rocks, especially on top of the sequence. The protolith of the PQ unit has been considered as a mid-Carboniferous to Triassic rift sequence (Krahl et al. 1983), originally lying west of the depocentre of the stratigraphically younger Tripolitza unit. Isolated relict slices of the preAlpine basement of the External Hellenides have also been identified within the PQ unit (Romano et al. 2004, and references therein). Structural and geochronological data from Kithira Island (Fig. la) document that the pre-Alpine rocks mainly occur in the contact zone between the PQ unit and the thrust sheets of non-metamorphic cover rocks (Xypolias et al. 2006). The PLK unit consists of upper Triassic to upper Eocene carbonate rocks, which rest on top of a thin Permo-Triassic sequence of dark metapelites referred to as the Kastania Phyllites (Fig. lb). The latter has been considered as the westward lateral equivalent of the PQ protolith (Doutsos et al. 2000). The tectonothermal evolution of the HP rocks of the External Hellenides began in Oligocene times and involved underthrusting of PQ protolith and its basement beneath the Tripolitza basement (e.g. Xypolias & Doutsos 2000; Kokkalas & Doutsos 2004; Fig. la (1)). In the course of this intracontinental subduction the PQ unit suffered HP-LT metamorphism, dated at c. 24 Ma (Panagos et al. 1979; Seidel et al. 1982). The degree of metamorphism appears to be higher in the southern Peloponnese and decreases both northward and southeastward. Quartz, phengite, paragonite, glancophane, chroritoid and locally garnet form the primary mineral assemblage in the study area (e.g. Katagas 1980; Skarpelis 1982; Theye 8: Seidel 1991). Variant peak P-T conditions in the PQ unit have been estimated at 400 __ 50~ and 11-13 kbar (Theye 1988), 450-t-30~ and 17___4 kbar (Theye & Seidel 1991) or 400-500~ and up to 10 kbar (Bassias & Triboulet 1993). Recent studies in the southern Peloponnese suggest that peak metamorphic conditions are characterized by

499

an upward temperature increase from 400-450~ at the base to 450-550~ at the top of the sequence, under nearly constant pressures (Blum6r 1998; Trotet 2000). The PLK unit has also been metamorphosed reaching peak P-T conditions at c. 7 - 8 kbar and 310-360~ in the Taygetos area (Blum6r et al. 1994) and c. 5 kbar and 450-480~ in the Paruon area (Bassias 1989). After peak metamorphism (Fig. la(2)), the PQ unit was detached from its basement and extruded upward to the west between a thrust fault at the base (referred to as 'basal thrust') and the Tripolitza basement at the top. The effect of this extrusion process was the emplacement of the PQ unit over the PLK unit bringing it into contact with the overlying cover thrust sheets. Several proposed P-T paths (e.g. Bassias & Triboulet 1993; Blum6r 1998) for the PQ unit in the Peloponnese imply near-isothermal decompression, as indicated by the relatively low pressures (c. 5 kbar) associated with relatively high temperatures (c. 350~ Zircon fission track ages from Crete also indicate that exhumation of the PQ unit to a depth of 15 km should have been completed before 19 Ma (Thomson et al. 1998). The last stage of exhumation of the HP rocks, which are now exposed in several tectonic windows (i.e. Taygetos, Parnon windows; Fig. lb), was characterized by regional backthrusfing, folding of the major thrust contacts around northsouth trending axes and gravity sliding of the cover thrust sheets (Fig. la(3)). In this study, we chose to analyse ductile deformation in two key areas located in the Parnon and Taygetos windows (Fig. lb). The PQ unit in the Parnon area occupies more internal and deeper parts of the shear zone compared to the Taygetos area. It is emphasized that in both areas we chose the gently to moderately west-dipping flanks of the tectonic windows, because these flanks are less affected by late stage deformation than the eastern flanks.

Major fabric elements in the PQ unit The PQ unit is affected, at the microscopic scale, by at least two ductile deformation events (see also Seidel et al. 1982; Doutsos et al. 2000; Zulauf et al. 2002). However, in many cases only the second penetrative event can be distinguished. Evidence for the first event (D1) is limited to an internal foliation ($1) in porphyroclastic HP-related minerals. The second deformation event (D2) postdates the main growth of HP-related minerals and is associated with the west-directed ductile extrusion of the PQ unit. D2-related structures and fabrics are pervasively developed throughout the PQ unit, and are by far the most common structural

500

P. XYPOLIAS & S. KOKKALAS

features in the study area. The main foliation ($2) is chiefly defined by alignment of micaceous films and flattened quartz aggregates. It wraps around clasts (i.e. chloritoid grains) and is associated with formation of a o-shaped clast matrix system and S/C fabrics indicating top-to-the-west sense of shearing. On $2 foliation planes an east-west to E N E - W S W trending stretching lineation is developed (Figs 2 & 3). The latter orientation is well expressed by quartz-mica alignment, strain shadows around garnet and pyrite crystals, as well as by elongated pebbles of metaconglomerates.

Sampling For the purpose of this study we collected samples from 54 locations within the PQ unit; 29 locations in Taygetos and 25 locations in the Parnon area. Our aim was the systematic collection of samples throughout the PQ unit in order to obtain estimates of the strain ratio and vorticity of flow with varying distances above the basal thrust. The observable structural thickness of the PQ unit is c. 1400 m in Taygetos (Fig. 2) and c. 1950 m in the Parnon area (Fig. 3). Therefore, the mean sample spacing, measured normal to the main foliation, is c. 50 m for the Taygetos area and c. 75 m for the Parnon

area. Details of sample locations and estimated sampling distances measured perpendicular to the basal thrust are given in Figures 2, 3 and Table 1. All analysed samples are characterized by a single, homogeneously developed, penetrative $2 foliation. In each sampling location, where possible, more than one sample was collected. Our aim for each locality was to collect samples that were best suited for each analytical technique employed, i.e. pure quartzite for quartz c-axis fabrics, quartzose metaconglomerate for strain analysis, and chloritoid-bearing schist for vorticity estimates. All microtectonic analyses were carried out on XZ thin-section planes, orientated parallel to stretching lineation and perpendicular to the main foliation.

Quartz c-axis fabrics Twenty-nine orientated quartz-rich (>90% quartz) samples from the PQ unit were collected for quartz c-axis fabric analysis. Fourteen of these samples were taken from the Taygetos area (Figs 2 & 4) and the remaining 15 from the Parnon area (Figs 3 & 5). All analysed samples are strongly foliated and show evidence of extensive dynamic recrystallization of quartz associated with both

Fig. 2. Structural map of the western flank of the Taygetos window, showing location of samples. The cross-section A1 -A2 illustrates the structural position of the samples. Note that the observed west-dipping geometry of the flank, and folding of the tectonic contact between PQ and PLK units were formed during the late exhumation stage which followed west-directed extrusion of the PQ unit. For location of map see Figure lb.

DEFORMATION Table

1.

AND DUCTILE

EXTRUSION

Petrofabric, strain and vorticity data for samples from the PQ unit in Taygetos and Parnon areas

Sampling location

Structural Distance (m)

/3-angle* (degree)

Opening angle* (degree)

Strain analysis

Rxz +_ ( l o - ) Taygetos AA2 AA10 AA3

Error (%)

Vorticity Method

analysis 1'

(Wm)

Method

area

AA4 AAll AA6 AA12 AA7 AA5 AA9 AA8 ABI7 ABI6 AB9 AB15 AB14 AB8 AB13 AB6 AB1 AB4 AB5 AB2 AB7 AB3 A B 18 AB12 AB10 ABll Parnon PA26 PAl7 PA46 PA21 PA23 PA24 PA25 PA30 PA27 PA29 PA22 PA28 PA31 PA33 PA34 PA32 PA42 PA35 PA41 PA40 PA39 PA36 PAl2 PAl3 PA37

501

1380 1230 1120 910 820 770 740 720 690 590 560 500 490 480 470 390 360 350 300 260 250 240 220 200 190 180 80 70 60

10-15 8-13 7-11 9-13 7-11 10-15 13-17 11-16 18-22 13-18 10-14 11-15

54 56 45 41 57 56 59 61 56 54 55 57 55 53

3.6 3.5 4.0 4.4 3.8 4.2 4.0 4.8 4.6 4.0 4.6 5.8 4.8 4.4 4.8 5.6 5.0 5.7 5.5 5.8 5.4 6.7 6.2 7.1 6.6 5.7 7.6 6.8 8.2

_ 0.3 __+ 0 . 4 +__ 0 . 3 ___ 0 . 3 _ 0.4 __+ 0 . 4 + 0.4 ___ 0 . 4 __+ 0 . 2 + 0.3 + 0.3 _ 0.6 ___ 0 . 4 __+ 0 . 3 ___ 0 . 5 + 0.4 ___ 0 . 3 +_ 0 . 6 _ 0.5 __+ 0 . 6 _ 0.6 • 0.4 _ 0.5 + 0.5 + 0.5 +__ 0 . 7 __+ 0 . 4 _ 0.6 _ 0.6

17 22 15 14 21 20 18 16 10 15 13 20 17 14 20 16 14 22 18 21 21 13 17 15 16 25 11 18 15

0.67-0.86 0.56-0.79 0.53-0.74 0.68-0.84 0.61-0.81 0.74-0.90 0.87-0.96 0.85-0.96 0.98-0.99 0.90-0.98 0.89-0.97 0.93-0.98

0.60-0.68 0.72-0.77 0.74-0.79 0.73-0.78 0.79-0.82 0.84-0.86 0.82-0.85 -

56 62 57 60 61 61 61 58 56 -

3.0 3.4 2.9 3.9 3.2 2.8 3.1 3.3 3.3 3.4 3.3 3.7 4.0 3.8 4.2 3.9 3.8 4.3 4.6 4.3 5.2 4.7 5.9 6.0 6.3

+ + + + + + + + + + + + + + + + + + + + + + + + +

24 16 27 14 18 23 20 22 13 18 20 19 17 23 17 15 17 27 16 14 19 20 19 13 18

0

0.35-0.51 0.57-0.66 0.72-0.77 0.66-0.72 0.77-0.81 -

area 1950 1930 1900 1820 1810 1800 1780 1660 1580 1340 1310 1150 960 930 870 800 700 600 560 440 300 240 120 100 60

0 0 0 0 4-8 4-9 2-7 13-19 11-15 10-14 12-17 10-15 11-15 8-13 -

*For definition see inset in Figure 4. *Win estimated integrating quartz c-axis fabric and strain data. *Win estimated using rigid rotated porhyroclasts.

0.4 0.3 0.4 0.2 0.3 0.3 0.3 0.4 0.2 0.3 0.3 0.3 0.3 0.4 0.4 0.3 0.3 0.6 0.4 0.4 0.5 0.5 0.6 0.4 0.6

0 0 0 0.30-0.55 0.32-0.63 0.17-0.53 0.82-0.95 0.78-0.90 0.75-0.89 0.81-0.94 0.76-0.92 0.86-0.96 0.74-0.92 -

2*

502

P. XYPOLIAS & S. KOKKALAS

Fig. 3. Structural map of the western flank of the Parnon window, showing location of samples. The cross-section A1-A2 illustrates the structural position of the samples. Note that the observed west-dipping geometry of the flank, and folding of the tectonic contact between PQ and PLK units were produced during the late exhumation stage which followed west-directed extrusion of the PQ unit. For the location of map see Figure I b. Map legend is the same as in Figure 2.

low temperature grain boundary migration and subgrain rotation (regime 2 of Hirth & Tullis 1992; Fig. 6a). In each sample, the orientation of 250 or more recrystallized quartz grains was measured on XZ thin-sections using a Leitz universal stage mounted on a Nikon microscope. The stereographic projection of the c-axis data and the contouring of these data were made using the PC software package StereoNett (by J. Duyster; Ruhn-Univerit~it Bochum). In Figures 4 and 5 the c-axis fabric diagrams of all the analysed samples are presented.

Patterns and shear sense All 14 quartzite samples from the Taygetos area are characterized by well-developed quartz c-axis fabrics (Fig. 4). In terms of density distribution all diagrams may be described as Type-I cross-girdle fabrics (Lister 1977) with c-axis point maxima mainly at high angles to the foliation trace. Twelve out of the 14 diagrams, from varying distances above the basal thrust, display an obliquity

of the central girdle segment (external parameter, ~p; Law 1987), with respect to the main foliation and lineation, indicating a component of noncoaxial top-to-the-west shear. In contrast, the remaining two samples (Fig. 4, diagrams AA3, AA4) yielded clear asymmetric fabrics indicating an opposite sense of shearing (top-down-tothe-east). These samples are located at 1120 and 910 m above the basal thrust, respectively. The quartz c-axis fabric diagrams obtained from the PQ unit in the Parnon window can be grouped into three different patterns: cleft girdles, and Type-I and Type-II crossed girdles (Fig. 5; Lister 1977). Symmetrical cleft girdle patterns are observed in four samples (Fig. 5; diagrams PA24, 26, 30, 46) collected from the uppermost structural levels of the PQ unit at distances of 1650-1950 m above the basal thrust (Fig. 5, Table 1). One sample (PAl7), which is located at the contact with the overlying cover thrust sheets, yielded a slightly asymmetric cleft girdle fabric indicating top-to-the-west shearing. Type-I crossed girdle

DEFORMATION AND DUCTILE EXTRUSION

503

Fig. 4. Optically measured quartz c-axis fabrics from 14 samples collected from the Taygetos area; lower hemisphere equal-area projections; foliation is orientated west-east and is vertical; stretching lineation is horizontal. Contours are generally 0.8, 1.6, 3, 5 times uniform distribution. The location of samples is projected onto a restored cross-section (lower left). This section was constructed by restoring the cross-section A1-A2 in Figure 2, assuming that westdirected ductile extrusion of the PQ unit occurred along a low-angle east-dipping shear zone. For location of samples see also Figure 2.

fabrics were recorded in eight samples (Fig. 5; diagrams PAl2, 13, 22, 28, 35, 36, 40, 41), which were collected from the lower structural levels of the unit. A slightly asymmetric Type-I crossed girdle pattern consistent with west shear sense was detected in two (PA22, 28) of these samples, which are located at distances of 1310 and 1150 m above the basal thrust, respectively (Fig. 5). The contoured fabric diagrams from the remaining six samples display clear asymmetry of the central girdle segment (Fig. 5, Table 1), with respect to the main foliation and lineation, indicating a component of non-coaxial top-tothe-west shear. These six samples were collected from the lowermost structural levels of the PQ

unit (distances of 100-800 m; Table 1). One sample (PA32) exhibits an asymmetric fabric pattern consistent with west shear sense, which can be classified between cleft girdle and Type-I crossed girdle pattem. A slightly asymmetric Type H crossed girdle pattern was observed in one sample (PA33), located at a distance of 930 m above the basal thrust (Fig. 5, Table 1). Summarizing, it seems that cleft girdle quartz c-axis patterns characterize the upper structural levels of the PQ unit in Parnon area, while Type-I crossed girdle pattems characterize the lower levels. Moreover, there is an increase in the asymmetry of pattems, consistent with west shear sense, from the top to the base of the unit.

504

P. XYPOLIAS & S. KOKKALAS

Fig. 5. Optically measured quartz c-axis fabrics from 15 samples collected from the Parnon area; lower hemisphere equal-area projections; foliation is orientated west-east and is vertical; stretching lineation is horizontal. Contours are generally 0.8, 1.6, 3, 5 times uniform distribution. The location of samples is projected onto a restored cross-section (lower left). This section was constructed by restoring the cross-section A1-A2 in Figure 3, assuming that west-directed ductile extrusion of the PQ unit occurred along a low-angle east-dipping shear zone. For location of samples see also Figure 3. For explanation of quartz c-axis diagrams see Figure 4.

Finite strain symmetry Although quantitative analysis for three-dimensional strain is not presented in this study (see below), information about strain symmetry can be retrieved from the obtained quartz c-axis fabrics (e.g. Schmid & Casey 1986; Law 1990). The crossed-girdle pattern of quartz c-axis fabrics in 20 samples (Figs 4 & 5) is interpreted to indicate approximately plane-strain (k = 1) conditions during top-to-the-west ductile shearing in the PQ unit. The cleft girdle pattern in five samples (Fig. 5), collected from the uppermost structural levels of PQ unit in the Pamon area, imply deviation from plane-strain indicating ductile deformation under constrictional conditions (k > 1) for this part of the unit.

Deformation temperatures In order to estimate temperatures during ductile deformation and dynamic recrystallization of the

analysed quartzites, the geothermometer of Kruhl (1998) modified by Law et al. (2004) and Morgan & Law (2004) was used. The geothermometer correlates graphically the opening angle of the quartz c-axis fabric with increasing deformation temperature, with an uncertainty of __+50~ (for details see Morgan & Law 2004, p. 220, fig. 5). For planestrain conditions, the opening angle (OA; Fig. 4, inset) is defined as the angle between the two c-axis girdles measured in the XZ plane. Ten out of 12 Type-I crossed-girdle fabrics from Taygetos area, consistent with top-to-the-west sense of shear, are characterized by c-axis point maxima that formed by basal (a) and [c] slip and display opening angles ranging from 53 ~ to 57 ~ For the Parnon area, the opening angles in seven samples are slightly greater and range between 56 ~ and 62 ~ Using the geothermometer, these angles indicate deformation temperatures of 400-430___ 50~ and 420-470 _ 50~ for the Taygetos and Parnon areas, respectively.

DEFORMATION AND DUCTILE EXTRUSION

505

Fig. 6. Micrographs of samples from the PQ unit in Taygetos and Parnon areas; all micrographs are from thin sections cut perpendicular to foliation and parallel to stretching lineation. All sections are viewed towards the north. Scale bar in each micrograph is 0.2 mm. (a) Sample AB 14: typical quartzite. The elongate and deformed quartz grains form a well-developed shape-preferred orientation which defines a continuous foliation. The quartz grains also show evidence of dynamic recrystallization (crossed nicols). (b) Sample PA34: plastically deformed quartz (Qtz) grains used for strain analysis (crossed nicols). (c) Sample AB4: rigid &type chloritoid (Ct) porphyroclast rotating in a homogeneous deforming matrix (top-to-the-west sense of shear); matrix consists mainly of white mica (Mc); secondary chlorite (Ch) growth in strain shadow (crossed nicols). (d) Sample ABS: plastically deformed quartz (Qtz) grains used for strain analysis (crossed nicols). (e, f, g, h) Samples PA42, AA8, AB13, AA5: rigid o-type chloritoid (Ct) porphyroclasts rotating in a homogeneous deforming matrix; all samples indicate top-to-the-west sense of shear (e, f, h: crossed nicols; g: plane polarized light). See Figures 2 and 3 for sample localities.

506

P. XYPOLIAS & S. KOKKALAS

Examining these results in combination with the structural position of the studied samples on two restored cross-sections (Figs 4 & 5), it seems that temperature during deformation of the PQ unit was nearly constant along the two hinterland-dipping sampling lines. These data could be interpreted as indications that isotherms were hinterland-dipping during ductile extrusion of the PQ unit. The two fabric diagrams from Taygetos area, which indicate a top-to-the-east sense of shear (Fig. 4, diagrams AA3, AA4), are characterized by c-axis point maxima that formed mainly by basal (a) slip and have opening angles ranging from 41~ to 45 ~ suggesting deformation temperatures of 325-350 _ 50~ This relatively low deformation temperature enables us to assume that backward-directed ductile shearing occurred on the upper part of the sequence during the late stage of exhumation.

Finite strain analysis Finite strain ratio in the XZ plane (Rxz) was estimated at 54 locations within the PQ unit (29 in Taygetos and 25 in Pamon area) using elliptical strain markers (Ramsay & Huber 1983, p. 73), such as plastically deformed quartz grains (Fig. 6b, d) in fine-grained metaconglomerates, metapsammites and metasiltstones (in 50 samples), and elliptical pebbles in deformed metaconglomerate layers (from four roadcut sections). It is emphasized that strain analysis was not attempted on pure quartzites due to the fact that in these rocks the aspect ratio of quartz grains was significantly modified by dynamic recrystallization. At least 60 objects were digitized from enlarged photographs of thin-sections in each sample and then Rf/cp diagrams were constructed using the software package INSTRAIN v. 3.02 (Erslev & Ge 1990). This program calculates the mean object ellipticity (Rxz strain ratio) based on a least-square algorithm. It also gives the average radial error of this calculation, which can be expressed as standard deviation at _ 1o(Table 1). The matrix strain in ten samples of varying lithologies was also estimated using the normalized Fry method (Erslev & Ge 1990) in order to compare the results with those obtained by the gf/~p shape analysis (particle strain). The comparison revealed small differences between the methods, in the order of 4-8%, indicating that the competence contrast between matrix and particles was not significant during strain accumulation. Hence it seems reasonable that the Rxz values obtained by the Rf/cp method are meaningful and reflect bulk strain in the samples.

Finite strain profiles The results of strain analysis are listed in Table 1 and plotted in a graph (Fig. 7), which displays

Fig. 7. Graph illustrating the variation in Rxz (strain ratio in XZ section) values versus distances above the basal thrust; + lo- (standard deviation) error bars are shown. For further details see Table 1.

strain ratio (Rxz) versus distance above the basal thrust. The results from both areas indicate a nonlinear (possibly negative exponential) increase in strain magnitude towards the lower structural levels of the PQ unit. In more detail, the Rxz values increase slightly from c. 3.0-3.5 near the upper structural levels of PQ unit to about 4.0 at a projected distance of 800-900 m above the basal thrust, for both areas. A rapid increase in the Rxz value is observed from the middle (c. 800900 m) to the bottom of the unit, where the strain ratio reaches values of 9.5 and 7.0 in the Taygetos and Parnon areas, respectively.

Vorticity analysis Mean kinematic vorticity number (Wm) is the most commonly used numerical measure to specify the shear-induced or internal vorticity (e.g. Jiang 1999) caused by the non-coaxial component of deformation. W,,, is considered as a non-linear relation between the pure shear and simple shear components of deformation, with Wm = 1 implying simple and Wm= 0 implying pure shear flow. Pure and simple shearing components contribute equally to the instantaneous flow at Wm = 0.71 (Law et al. 2004). Several methods have been proposed to estimate two-dimensions (2D) kinematic vorticity in natural deformed rocks (Wallis 1995, and references therein). Vorticity analysis in three-dimensional has

DEFORMATION AND DUCTILE EXTRUSION

507

been attempted but is complicated by a number of unsolved problems (see also Tikoff & Fossen 1995). All 2D methods require orthorhombic or monoclinic ductile flow with the vorticity vector (w) parallel to the Y-axis of finite strain. The nearly plane-strain deformation conditions which are suggested by the crossed girdle pattern of the quartz c-axis fabrics in some of our samples, as well as the abundance of asymmetric structures on XZ planes, are compatible with the required conditions, implying that 2D vorticity analysis is valid in our samples. In order to estimate W,n from our samples, we used two methods which are based on completely different vorticity criteria: Method 1 combines quartz c-axis fabric and strain data while Method 2 examines the behaviour of rotated porphyroclasts. Whether these two methods provide time-averaged measurements of rock flow or record different parts of deformation history remains an open question (Tikoff & Fossen 1995; Wallis 1995; Bailey et al. 2004; Law et al. 2004). Therefore, wherever possible both methods were applied to samples collected from a given locality. Both vorticity methods are well established and have been described in detail by others (see Wallis 1995, and references therein). Below we briefly summarize our application of these analyses. Method 1 was proposed by WaUis (1995), who demonstrated that the angle /3 (Fig. 4, inset) between the perpendicular to the central girdle of a crossed girdle quartz c-axis diagram and foliation is a function of Rxz (strain ratio) and Win:

sin/tolEE

l

1_______~) (Rxz+ ] • (Rxz - 1)

(1)

(Rxz § 1)/(Rxz - 1)] - cos (2/3)

This method was applied to 14 samples collected from the Taygetos area and ten samples from the Parnon area, where asymmetric quartz c-axis fabrics are observed. Because the central girdle segments of quartz c-axis fabrics are not perfectly straight, the determination of angle/3 involves an error (about 4-5~ Table 1). This error, as well as uncertainty in strain ratio, were taken into account in the calculation of Wm. The range in Wm values for each sample is presented by error bars in Figure 8, which represents the graphical expression of Equation 1. The results of this analysis are also listed in Table 1. The analysis of samples from both areas yielded values of Wm in the range of 0.60-0.99 for Taygetos and 0.30-0.90 for the Parnon area.

Fig. 8. Vorticity numbers (Win) calculated by Method 1 in PQ unit rocks at Taygetos and Parnon areas taking into account uncertainties in determination of Rxz strain ratio and/3-angle. Contours of equal vorticity number indicated. Method 2 was proposed by Wallis et al. (1993) and is based on recording graphically the relation between the orientation and aspect ratio of rigid porphyroclasts rotating in a homogeneous matrix. On such graphs (Fig. 9), there is a critical aspect ratio (Rc) below which the measurements scatter in a wide range of orientations, reflecting continuously rotated porphyroclasts, and above which clasts lie close to the foliation plane implying a stable end orientation of porphyroclasts. The Rc is

508

P. XYPOLIAS & S. KOKKALAS

Fig. 9. Stable-orientation analysis (vorticity Method 2) for chloritoid porhyroclasts in 12 samples. Angle ~p denotes the angle between the long axis of chloritoid and mesoscopic foliation, R the aspect ratio of chloritoid porphyroclasts, and Rc the value of the critical aspect ratio. The number (n) of objects measured in each sample is also indicated. The shaded parts of the diagrams indicate the range of uncertainty in the R,: estimation.

a function of Win, where Wm = (R 2 - 1)/(R 2 + 1) (Passchier 1987; Wallis et al. 1993). This method may tend to underestimate the vorticity number if porphyroclasts of large aspect ratio are not present (Law et al. 2004). Alternative interpretations suggest that at low strains rigid porphyroclasts may not have rotated far enough to reach their stable orientation and therefore such a method can overestimate the vorticity number (Bailey et al. 2004). Stable orientation analysis (Fig. 9), in 12 of our samples containing chloritoid porphyroclasts (Fig. 6c, e, f, g, h), yielded well defined cut-off aspect ratios (Rc) implying that vorticity estimates are meaningful and close to the true Wm value. Analysis of eight samples from Taygetos and four samples from the Parnon area using this method yielded Wm values in the range of 0.68-0.86 and 0.51-0.81, respectively (Fig. 9, Table 1).

Kinematic vorticity profiles To examine the spatial variation of kinematic vorticity within the PQ unit we constructed graphs showing the distribution of Wm in vertical profiles for the Taygetos and Parnon areas, respectively (Fig. 10a, b). From both vorticity profiles it seems that in any level the Wm values estimated by Method 2 are, with few exceptions, slightly lower than the values estimated by Method 1. However, it is also obvious that Wm estimates from both methods record the same pattern in each profile. Therefore, we propose that the observed Wm variation reflects a true flow-path partitioning in the PQ unit. Similar differences in the estimates of these two methods have been recorded in the High Himalayan zone by Law et al. (2004).

DEFORMATION AND DUCTILE EXTRUSION

509

Fig. 10. Graph illustrating the variation in W m values versus structural distance above the basal thrust in the Taygetos (a) and Parnon (b) areas, respectively. Length of bars reflects the uncertainty in the calculation of Win.For further details see Table 1.

The vorticity profile of the Taygetos area (Fig. 10a) shows that the W m value is close to 1 at the base of the PQ unit and decreases progressively upward reaching a value of 0.71 (equal contribution of pure and simple shear) at a projected distance of c. 700-800 m above the basal thrust. At progressively higher structural levels the Wm tends to be constant around the value 0.71. In the Parnon area (Fig. 10b), the W,n decreases progressively from 0.9 close to the basal thrust to 0.71 in the middle (c. 900 m) of the unit, showing a similar distribution to the Taygetos profile within the lower structural levels. The Wm decreases drastically within a distance interval of c. 300 m, reaching an extrapolated value of 0.2 at a distance of 1200 m above the basal thrust. Based on the symmetric cleft girdle quartz c-axis patterns, which are observed on the uppermost structural levels (1650-1950 m) of the PQ unit in the Parnon area, we assume that W,,, approaches zero for this part (Fig. 10b).

Interpretation and synthesis The PQ unit defines an east-dipping intracontinental shear zone, which is characterized by a high aspect

ratio (transport-parallel length/thickness normal to the shear zone). Kinematic indicators within the unit indicate a clear top-to-the-west shear sense, similar to that proposed by others for the northern Peloponnese (Xypolias & Koukouvelas 2001). Evidence for top-to-the-east shearing (backward motion) is mainly found in samples collected from the frontal and upper structural levels of the extruding zone. The west-directed ductile shearing in the PQ unit possibly occurred at deformation temperatures of 400-470 __. 50~ while the east-directed movements occurred at lower temperatures (c. 350~ These data enable us to assume that backward shearing occurred during the late stages of the ductile extrusion process. Strain ratio is nearly constant (Rxz ~ 3.0-4.0) within the upper structural levels of the zone but increases systematically from the middle to the bottom, approaching a value of c. 9.5 close to the basal thrust in the frontal part (Taygetos area) and a value of c. 7.0 in the inner part (Parnon area). Based on the observed strong correlation of strain magnitude with distance above the basal thrust, we suggest that tectonic emplacement of the PQ unit over the PLK unit occurred during accumulation of ductile strain

510

P. XYPOLIAS & S. KOKKALAS

within the former unit. A similar tendency for down-section strain increase has been observed in previous studies from the northern Peloponnese (Xypolias & Doutsos 2000), the eastern flank of the Taygetos window (Doutsos et al. 2000), as well as from eastern Crete (Kokkalas & Doutsos 2004). However, compared to other regions in Greece, higher Rxz strain values are recognized in the current study area (southern Peloponnese). These data imply both a northward and southeastward decrease in strain along the orogen. Quantitative data on the vorticity of the flow show that a simple-shear-dominated domain (0.71 < Wm < 1.0) occupies the lower half of the shear zone. The pure shear component usually increases upwards in the zone, and specifically appears to dominate the upper structural levels ( W m < 0.4) of the inner part of the zone (Parnon area).

Thinning and dip-parallel elongation Deviation from ideal simple shear deformation implies thinning of the PQ unit perpendicular to the shear zone boundaries (flow plane) and resultant dip-parallel elongation. For isochoric plane strain deformation, the stretch magnitude both normal and parallel to the flow plane can be calculated, combining strain and vorticity data by using the mathematical expression (Fig. l ld) suggested by Wallis et al. (1993). However, the application of this equation in non-plane-strain situations gives underestimated values of dip-parallel elongation for constrictional strain or overestimated values for flattening strain (R. D. Law, pers. comm. 2005). In this study, petrofabric analysis revealed that PQ unit rocks in the Taygetos and the lower structural levels of the unit in the Pamon area were deformed under approximately plane strain (k = 1) conditions. For these parts, where possible, strain and vorticity data from a given sample location were integrated using the equation in Figure l ld. Uncertainties in strain ratio and vorticity number were taken into account in the calculation of stretch magnitude. The range in dip-parallel elongation values for each sample is presented by error bars in the graphs of Figure 1 la and b(2), which represent the graphical expression of the equation suggested by Wallis et al. (1993). The estimated values of dip-parallel elongation generally range from 40 to 90%. The cleft girdle pattern in five samples (Fig. 5), collected from the uppermost structural levels of the PQ unit in the Pamon area, implies ductile deformation under constrictional conditions (k > 1). Three-dimensional strain analyses in samples from this part confirm this assumption, showing that these samples are characterized by a Lode's parameter smaller than zero ( - 0 . 5 < v < - 0 . 2 ) and

a strain intensity (es) ranging from 0.7 to 0.9. Plotting these values on a Hsu diagram (Fig. llb(1)), a dip-parallel elongation on the order of 70-100% is calculated for this part of the PQ unit in the Parnon area. To investigate the spatial variation of the estimated stretch values in the Taygetos and Parnon areas, two vertical profiles were constructed (Fig. 1 lc). In both profiles, the average stretching parallel to the shear zone appears to be higher in the middle and upper structural levels of the PQ unit, ranging between 60 and 90%, and lower in the deeper parts of the zone, ranging from 40% to 60%. Restoring a detailed cross-section across the southem Peloponnese, similar to that illustrated in Figure la(3), we estimate that the minimum length of the PQ unit at the base of the shear zone is c. 80 km. Therefore, a mean stretch value of 1.5 parallel to the base of the shear zone, implies an original length of c. 53 km. Both vertical profiles indicate no significant variation in ductile thinning normal to the shear zone, which is on the order of 30-45%. Considering that the PQ unit in the study area is a 1.5-2 km thick zone, the calculated values imply an original thickness (pre-ductile extrusion) of c. 2.5-3.5 km.

Displacement and extrusion rate The PQ unit is considered to lie between two rigid and downward-converging basement sheets. Therefore, the obtained elongation parallel to the shear zone boundaries, which is the effect of the pure shear component of ductile flow, confirms an upward extrusion of the PQ unit and implies the formation of 'stretching faults' (in the sense of Means 1989) along both the lower and the upper surface of the unit as a consequence of extrusion. In this case, slip of the material along the basal stretching thrust fault should be induced by the pure shear component of flow, while along the roof stretching fault slip is induced by both pure and simple shear components (Fig. 12a). In order to measure minimum displacement along the basal thrust using a common PC graphic program, we created a 53 mm long slice (scale 1:100 000) and pinned it at the right edge, which was then 'numerically stretched' by 50%. This resulted in the basal slip surface recording an increasing displacement towards the left (foreland or west) end of the stretched slice along the basal slip surface (Fig. 12a, graph). The displacement of the most distal reference point (between F and F') along the basal thrust fault is c. 25 km (Fig. 12a). A similar procedure was followed to measure displacement due to the pure shear component along the roof fault. In this case, a slice of the same original length stretched by 75%. Along the roof slip

DEFORMATION AND DUCTILE EXTRUSION

511

Fig. 11. (a) Percentage elongation (E) parallel to flow plane in transport direction for samples from Taygetos area calculated by integrating strain and vorticity data and assuming plane strain deformation. Wm and Rxz values for this analysis are listed in Table 1. Contours of equal elongation are calculated applying equation in (d). (b) Percentage dip-parallel elongation for uppermost structural levels of PQ unit in Parnon area (2) calculated plotting threedimensional strain data on a Hsu diagram (1). Percent elongation for the middle and lower levels of the PQ unit calculated by integrating strain and vorticity data and assuming plane strain deformation. (e) Diagram showing the range of thinning and elongation values calculated perpendicular and parallel to the shear zone versus structural distance above the basal thrust in the Taygetos and Parnon areas. The results are plotted for 100 m distance interval (vertical error bar). Length of horizontal error bars indicates the possible range of stretch values. Each rectangle indicates the average stretch value for samples in a specific structural level with an interval of 100 m. The individual sample numbers used for this analysis are also shown. (d) Equation used for calculation of shortening value (S), perpendicular to the shear zone, from Wm and Rxz values; elongation parallel to the shear zone, is given by S-1 (after Wallis et al. 1993). Equation applied to samples deformed under plane strain conditions.

512

P. XYPOLIAS & S. KOKKALAS

Fig. 12. (a) Diagrammatic explanation of procedure followed in order to calculate the displacement of material due to pure and simple shear components of deformation. Two 53 mm (scale 1:100 000) long slices, representing the uppermost and the lowermost parts of the PQ unit, were horizontally stretched by 75% and 50%, respectively. Results show that both the upper and the lower surface are stretching faults. A to F, A' to FI and A" to Fr~ are reference points. The graph (lower left) shows displacement values versus reference points. Displacement due to simple shear component of deformation was measured using Wm and Rx: values obtained for Parnon area. Knowing the rotational and stretching components of deformation, and assuming homogeneous deformation in a thickness interval of 200 m, ten quadrangle boxes were deformed and arranged one next to the other. (b) On a planar shear zone, the obtained difference in displacement along the upper and lower surfaces induces a space problem in front of the extruding zone. A curvilinear shear zone in combination with discrete slip along a thrust-parallel zone in the middle could overcome this problem.

DEFORMATION AND DUCTILE EXTRUSION surface, the material movement also accelerates towards the left (west) reaching at the reference point F" a value of c. 38 km (Fig. 12a). Given that the variation of both Wm and Rxz has been specified throughout the entire thickness of the PQ unit, the offset due to the simple shear component of flow can be estimated using the method presented in Fossen & Tikoff (1993). As schematically displayed in Figure 12a, this displacement is on the order of 3 km. Therefore, the total displacement of reference point F" (F-F") along the roof fault is approximately 41 krn (Fig. 12a, graph). Assuming ductile extrusion of the PQ unit along a low-angle shear zone and adopting proposed P-T paths and structural geometries (i.e. mean dip value of 30 ~ for the area, then it seems that a minimum displacement of c. 40 km is required for bringing to a depth of 15-20 km a material point originally buried at c. 4 0 - 4 5 km. This implies that the estimated minimum displacements (25-41 kin), which are mainly induced by the pure shear component, make an important contribution to the net displacement of material during ductile extrusion. Moreover, based on these displacement values, and taking into account that the exhumation of the PQ unit took place from 23 to 19 Ma (Thomson et al. 1998), we estimated slip rates of c. 6.5 m m / year and c. 10.5 ram/year for the frontal parts of the basal and roof stretching faults, respectively. The above analysis also indicates that displacement of the material along the roof stretching fault (41 km) is higher than that of the basal thrust (25 km). Therefore, within the frontal domain of the extruding zone a difference in displacement of c. 16 km is estimated. This difference in displacement, on a planar 1.5-2.0 km thick shear zone, could be explained with extremely high strain values (Rxz > 15-20) at the base to maintain strain compatibility (Fig. 12b). However, an Rxz value of 15-20 seems unrealistic for the studied rocks. Therefore, discrete slip along a thrust-parallel zone in the middle of the extruding slice, which separated differently stretched layers above and below, could occur during ductile extrusion (Fig. 12b). Alternatively, ductile upward extrusion of the PQ rocks along a curvilinear zone would partly overcome the observed displacement differences (Fig. 12b). In a curved shear zone the upper arc is obviously longer than the lower arc and therefore difference in displacement can be partly balanced. There is no strong field evidence to support significant slip along a distinct shear zone in the middle, and a curvilinear geometry is not sufficient to independently overcome the space problem, therefore we assume that discrete slip and curvature of the shear zone have operated contemporaneously during the upward extrusion of the PQ unit.

513

Comparison with other zones of ductile extrusion The observed increase in Rxz strain ratios from 3.04.0 at the top to 7.0-9.5 at the base of the PQ unit is comparable with patterns recorded within other zones of ductile extrusion. For instance, low Rxz strain ratios of about 2.5 to 3.5 have been recorded at the uppermost structural levels of the extruding High Himalayan zone (Law et al. 2004), and high strain ratios, possibly up to 20, close to the Main Central Thrust at the base (Grasemann et aL 1999). Moreover, an overall downward strain increase is at least locally independently documented within this zone by the intensity of lattice preferred orientation of quartz c-axis fabrics (e.g. Bhattacharya & Weber 2004). The strain magnitude within the (U)HP nappes of the Penninic Alps is also considered to increase downward from a value of 2.0-6.0 at the top to an extremely high value (c. 30) near the base (Escher & Beaumont 1997). The observed variation in strain symmetry along the PQ unit, which is characterized by localized constrictional strain in the innermost part (Parnon area) and approximately plane strain geometry in the remaining parts, also resembles strain patterns recorded within the Penninic nappes. For instance, extrusion/exhumation of the Dora Maira Massif and the Adula Nappe is accompanied by constrictional strain deformation (dip-parallel elongation), especially at the deepest and innermost parts, of the nappes and to a lesser extent by plane strain (e.g. Kurz 2005). The steepening at the rear parts of the zone, the increase of horizontal compressional stresses due to progressive continental underthrusting, and the upward material escape at high rates (Kurz 2005) are possible factors controlling constriction and axial elongation within the deepest parts of extruding zones. The observed spatial variation in flow type from this study show similarities with the pattern of vorticity distribution observed in the extruding High Himalayan zone, which is characterized by a strong reduction of simple shear component of deformation from the base to the centre of the zone and also a general shearing with opposite shear sense at the top (e.g. Grasemann et al. 1999; Law et al. 2004; Bhattacharya & Weber 2004; but cf. Jessup et al. 2006). However, opposite downdip shearing on the top of the studied shear zone is much less pronounced than in the Himalaya case. The estimated values of ductile thinning (c. 3 0 45%) normal to the studied shear zone, and the resultant dip-parallel elongation (c. 40-90%), are comparable with the ductile thinning of c. 43% (or c. 75% elongation) which has been estimated by Vannay & Grasemann (2001) for the whole

514

P. XYPOLIAS & S. KOKKALAS

High Himalayan zone along the Sutlej Valley of N W India and with the average thinning values of c. 20% for the uppermost structural levels of this zone in the Everest area (Law et al. 2004).

Conclusions Quantitative strain and vorticity analyses, along an extruding shear zone in the south Peloponnese, demonstrate that both strain ratio and the simple shear component of deformation increase downwards, towards the basal thrust. The observed pattern seems to be comparable with those observed in adjacent along-strike areas of this orogen, as well as in zones of ductile extrusion of other orogens. Even for a high simple shear component, the shear parallel stretching can be significant if strain magnitude is sufficient. Quantitative analyses of this work showed that the simple-shear-dominated domain along the base is associated with shearparallel elongation of 4 0 - 6 0 % , implying the presence of a stretching basal thrust. Shear-parallel elongation of 6 0 - 9 0 % identified along the upper surface of the zone also implies the presence of a roof stretching fault. The pure shear component of deformation makes an important contribution to the total displacement field of the studied zone of upward ductile extrusion. The simple shear component of deformation can partly contribute to the upward movement of rocks, but it seems to be more effective in thick shear zones. However, future work is needed to estimate to what degree the pure and simple shear components of deformation are sufficient on their own to explain the total displacement along a shear zone, and to determine the role of the other factors, such as buoyancy forces, in material extrusion. This work is dedicated to the memory of Theodor Doutsos, for his invaluable contribution to the understanding of geology of the Hellenides. We are grateful to S. Wallis and C. Bailey for helpful reviews of the manuscript; and to R.D. Law for thoughtful editorial comments as well as for sending us a copy of his unpublished strain and vorticity data from the Moine thrust zone. All contributions significantly improved the manuscript.

References BAILEY, C. M., FRANCIS, B. E. & FAHRNEY, E. E. 2004. Strain and vorticity analysis of transpressional high-strain zones from the Virginia Piedmont, USA. In: ALSOP, G. I., HOLDSWORTH, R. E., MCCAFFREY K. J. W. & HAND, M. (eds) Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224, 249-264.

BASSIAS, Y. 1989. Pal6og6ographie jurassique des 'Plattenkalk' ioniens dans le P61oponnbse oriental (Parnon), Grbce. Comptes Rendus de l'Acad~mie des Sciences s~ries II, 309, 275-281. BASSIAS, Y. & TRIBOULET, C. 1993. Tectonometamorphic evolution of blueschist formations in the Peloponnesus (Parnon and Taygetos Massifs, Greece) - A model of nappe stacking during Tertiary orogenesis. Journal of Geology, 102, 697-708. BHATTACHARYA, A. R. & WEBER, K. 2004. Fabric development during shear deformation in the Main Central Thrust Zone, NW-Himalaya, India. Tectonophysics, 387, 23-46. BLUMOR, T. 1998. Die Phyllit-Quartzit-Serie SE-Lakoniens (Peloponnes, Griechenland): Hochdruc kmetamorphite in einem orogenen Keil. Frankfurter Geowissenschaftiche Arbeiten, 17, 1-187. BLUMOR, T., DOLLINGER, J., KNOBEL, M., MUTTER, A., ZARDA, S. & KOWALCZYK,G. 1994. Plattenlalk series and Kastania phyIlites of the Taygetos Mts.: new results on structure and succesion. Geological Society of Greece Bulletin, 15, 83-92. BUCHER, S., SCHMID, S. M., BOUSQUET, R. & FOGENSCHUH, B. 2003. Late-stage deformation in a collisional orogen (Western Alps): nappe refolding, back-thrusting or normal faulting? Terra Nova, 15, 109-117. CAMEOS NETO, M. C. & CABY, R. 2000. Terrane accretion and upward extrusion of high-pressure granulites in the Neoproterozoic nappes of Brazil: Petrologic and structural constrains. Tectonics, 19, 669-687. COWARD, M. & DIETRICH,D. 1989. Alpine tectonics an overview. In: COWARD, M., DIETRICH, D. & PARK, R. G. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 1-29. DOUTSOS, T., PIPER, G., BORONKAY, K. & KOUKOUVELAS, I. 1993. Kinematics of the Central Hellenides. Tectonics, 12, 936-953. DOUTSOS, T., KOUKOUVELAS, I., POULIMENOS, G., KOKKALAS, S., XYPOLIAS, P. & SKOURLIS, K. 2000. An exhumation model of the south Peloponnesus, Greece. International Journal of Earth Science, 89, 350-365. ERNST, W. G. 2001. Subduction, ultrahigh-pressure metamorphism, and regurgitation of buoyant crustal slices - implications for arcs and continental growth. Physics of the Earth and Planetary Interiors, 127, 253-275. ERSLEV,E. A. & GE, H. 1990. Least-squares center-tocenter and mean object ellipse fabric analysis. Journal of Structural Geology, 12, 1047-1059. ESCHER, A. & BEAUMONT, C. 1997. Formation, burial and exhumation of basement nappes at crustal scale, a geometric model based on the Western Swiss-Italian Alps. Journal of Structural Geology, 19, 955-974. FOSSEN, H. & TIKOFF, B. 1993. The deformation matrix for simultaneous simple shearing, pure shearing and volume change, and its application to transpression-transtension tectonics. Journal of Structural Geology, 15, 413-422.

DEFORMATION AND DUCTILE EXTRUSION GRASEMANN, B., FRITZ, H. & VANNAY, J. C. 1999. Quantitative kinematic flow analysis from the Main Central Thrust Zone (NW-Himalaya, India): implications for a decelerating strain path and the extrusion of orogenic wedges. Journal of Structural Geology, 21, 837-853. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KUNDIG, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan Crystalline in Bhutan: evidence from the quartz microfabrics. Tectonophysics, 260, 21-43. HIRTH, G. & TULLIS, J. 1992. Dislocation creep regimes in quartz aggregates. Journal of Structural Geology, 14, 145-159. HODGES, K. V., BURCHFIEL, B. C., ROYDEN, L. H., CHEN, Z. & LIU, Y. 1993. The metamorphic signature of contemporaneous extension and shortening in the central Himalayan orogen: data from Nyalam transect, southern Tibet. Journal of Metamorphic Geology, 11, 721-737. JESSUP, M.J., LAW, R. D., SEARLE, M. P. & HUBBARD, M. S. 2006. Structural evolution and vorticity of flow during extrusion and exhumation of the Greater Himalayan Slab, Mount Everest Massif, Tibet/Nepal: implications for orogenscale flow partitioning. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 379-413. JIANG, D. 1999. Vorticity decomposition and its application to sectional flow characterization. Tectonophysics, 301, 243-259. JOHNSTON, D. H., WILLIAMS, P. F., BROWN, R. L., CROWLEY, J. L. & CARR, S. D. 2000. Northeastward extrusion and extensional exhumation of crystalline rocks of the Monashee complex, southeastern Canadian Cordillera. Journal of Structural Geology, 22, 603-625. KATAGAS, C. 1980. Ferroglaucophane and chloritoidbearing metapelites from the phyllite series, southern Peloponnese Greece. Mineralogical Magazine, 43, 975-978. KOKKALAS, S. & DOUTSOS, T. 2004. Kinematics and strain partitioning in the southeast Hellenides (Greece). Geological Journal, 39, 121 - 140. KRAHL, J., KAUFFMANN,G., KOZUR, H., RICHTER, D., FORSTER, O. & HEINRITZI, F. 1983. Neue Daten zur Biostratigraphie and zur tektonischen Lagerung der Phyllit-Gruppe und der Trypali-Gruppe auf der Insel Kreta (Griechenland). GeologischeRundschau, 72, 1147-1166. KRUHL, J. H. 1998. Prism- and basal-plane parallel subgrain boundaries in quartz: a microstructural geothermobarometer. Journal of Metamorphic Geology, 16, 142-146. KURZ, W. 2005. Constriction during exhumation: Evidence from eclogite microstructures. Geology, 33, 37 -40. LAW, R. D. 1987. Heterogeneous deformation and quartz crystallographic fabric transitions: natural examples from the Stack of Glencoul, northern Assynt. Journal of Structural Geology, 9, 819-833.

515

LAW, R. D. 1990. Crystallographic fabrics, a selective review of their applications to research in structural geology. In: I~IPE, R. J. & RUTTER, E. H. (eds) Deformation Mechanisms, Rheology and Tectonics. Geological Society, London, Special Publications, 54, 335-352. LAW, R. D., SEARLE, M. P. & SIMPSON, R. L. 2004. Strain, deformation temperatures and vorticity of flow at the top of the Greater Himalayan Slab, Everest Massif, Tibet. Journal of the Geological Society, London, 161, 305-320. LISTER, G. S. 1977. Discussion: Crossed girdle c-axis fabrics in quartzites plastically deformed by plane strain and in progressive simple shear. Tectonophysics, 39, 51-54. MEANS, W. D. 1989. Stretching faults. Geology, 17, 893-896. MORGAN, S. S. & LAW, R. D. 2004. Unusual transition in quartzite dislocation creep regimes and crystal slip systems in the aureole of the Eureka ValleyJoshua Flat-Beer Creek pluton, California: a case for anhydrous conditions created by decarbonation reactions. Tectonophysics, 384, 209-231. PANAGOS, A. G., PE-PIPER, G. G., PIPER, D. J. W. & KOTOPOULI, C. N. 1979. Age and stratigraphic subdivision of the Phyllite series Krokee region Peloponnese, Greece. Neues Jahrbuch far Geologie und Paliiontologie, Monatshefte, 1979, 181-190. PASSCHIER, C. W. 1987. Stable positions of rigid objects in non-coaxial flow-a study in vorticity analysis. Journal of Structural Geology, 9, 679-690. RAMSAY, J. G. & HUBER, M. I. 1983. The Techniques of Modern Structural Geology, vol. 1. Academic Press, New York. ROBERTSON, A. H. F., CLIFT, P. D., DEGNAN, P. & JONES, G. 1991. Palaeogeographic and palaeotectonic evolution of the Eastern Mediterranean Neotethys. Palaeogeography Palaeoclimatology Palaeoecology, 87, 289-344. ROMANO, S. S., DORR, W. & ZULAUF, G. 2004. Cambrian granitoids in the pre-Alpine basement of Crete (Greece): Evidence from U-Pb dating of zircon. International Journal of Earth Science, 93, 844-859. SCHMID, S. M. & CASEY, M. 1986. Complete fabric analysis of some commonly observed quartz c-axis pattems. In: HOBBS, B. E. & HEARD, H. C. (eds) Mineral and Rock Deformation Laboratory Studies: The Paterson Volume. American Geophysical Union, Geophysical Monograph, 36, 263 -286. SEARLE, M. P., SIMPSON, R. R., LAW, R. D., PARRISH, R. R. & WATERS, D. J. 2003. The structural geology, metamorphic and magmatic evolution of the Everest massif, High Himalaya of NepalSouth Tibet. Journal of the Geological Society, London, 160, 344-366. SEIDEL, E., KREUZER, H. & HARRE, W. 1982. A Late Oligocene/Early Miocene High Pressure Belt in the External Hellenides. Geologishes Jahrbuch, E23, 165-206. SKARPELIS, N. 1982. Metallogenesis of massive sulfides and petrology of the External metamorphic

516

P. XYPOLIAS & S. KOKKALAS

belt of the Hellenides. PhD thesis, University of Athens. THERE, T. 1988. Aufsteigende Hochdruckmetamorphose in Sedimenten der Phyllit-Quarzit-Einheit Kretas und des Peloponnes. PhD thesis, Universit~it Braunschweig. THEYE, T. & SEIDEL, E., 1991. Petrology of low-grade high pressure metapelites from the External Hellenides (Crete, Peloponnese) A case study with attention to sodic minerals. European Journal of Mineralogy, 3, 343-366. THIEBAULT, F. & TRmOULET, T. 1984. Alpine metamorphism and deformation in Phyllite nappes (external Hellenides southern Peloponnesus Greece): Geodynamic implication. Journal of Geology, 92, 185-199. THOMSON, S. N., STOCKHERT, B., RAUCHE, H. & BRIX, M. R. 1998. Thermochronology of the high-pressure metamorphic rocks of Crete, Greece: implications for the speed of tectonic processes. Geology, 26, 259-262. TIKOFF, B. & FOSSEN, H. 1995. The limitations of three-dimensional kinematic vorticity analysis. Journal of Structural Geology, 12, 1771 - 1784. TROTET, F. 2000. Exhumation des roches de Haute Pression-Basse Temperature le long d' un transect des Cyclades au P~loponnkse (GrOce). Implications Ggodynamiques. PhD thesis, Universit6 Paris XI. VANNAY, J. C. & GRASEMANN, B. 2001. Himalayan inverted metamorphism and syn-convergence extension as a consequence of a general shear extrusion. Geological Magazine, 138, 253-276. WALLIS, 8. R. 1995. Vorticity analysis and recognition of ductile extension in the Sanbagawa belt, SW

Japan. Journal of Structural Geology, 17, 10771093. WALLIS, S. R., PLATT, J. P. & KNOTT, S. D. 1993. Recognition of syn-convergence extension in accretionary wedges with examples from the Calabrian arc and the eastern Alps. American Journal of Science, 293, 463-495. XYPOLIAS, P. & DOUTSOS, T. 2000. Kinematics of rock flow in a crustal-scale shear zone: implication for the orogenic evolution of the southwestern Hellenides. Geological Magazine, 137, 81-96. XYPOLIAS, P. t~z KOUKOUVELAS, I. 2001. Kinematic vorticity and strain rate patterns associated with ductile extrusion in the Chelmos Shear zone (External Hellenides, Greece). Tectonophysics, 338, 59-77. XYPOLIAS, P., KOKKALAS,S. • SKOURLIS, K. 2003. Upward extrusion and subsequent transpression as a possible mechanism for the exhumation of HP/ LT rocks in Evia Island (Aegean Sea, Greece). Journal of Geodynamics, 35, 303-332. XYPOLIAS, P., DORR, W. t~ ZULAUF, G. 2006. Late Carboniferous plutonism within the pre-Alpine basement of the External Hellenides (Kithira, Greece): evidence from U-Pb zircon dating. Journal of the Geological Society, London, 163 539-547. ZULAUF, G., KOWALCZYK,G., KRAHL, J., PETSCHICK, R. ~z SCHWANZ, S. 2002. The tectonometamorphic evolution of high-pressure low-temperature metamorphic rocks of eastern Crete, Greece: constraints from microfabrics, strain, illite crystallinity and paleodifferential stress. Journal of Structural Geology, 24, 1805-1828.

The Appalachian Inner Piedmont: an exhumed strike-parallel, tectonically forced orogenic channel R O B E R T D. H A T C H E R , JR. & A R T H U R J. M E R S C H A T

Department o f Earth and Planetary Sciences & Science Alliance Center o f Excellence, University o f Tennessee-Knoxville, Knoxville, TN 37996-1410 USA (e-mail: bobmap @ utk. edu) Abstract: The Appalachian Inner Piedmont (IP) extends along orogenic strike some 700 km from North Carolina to Alabama. Its physical attributes contrast with those of other Appalachian tectonic elements: gentle dip of dominant foliation; imbricate stack of fold nappes; dominant sillimanite-grade metamorphism and near ubiquitous migmatization; heterogeneous, non-plane deformation; and earlier S-foliations transposed to C-foliations southeast of the mid-Palaeozoic Brevard fault zone forming a 10-20 km wide amphibolite-facies shear zone along the western flank of the IP. The IP contains west- and SW-directed thrust sheets and mineral stretching lineation, sheath folds on all scales, and other indicators that define a curved crustal flow pattern throughout the belt. Field and modern geochronologic data confirm that the IP is not exotic. It contains a Laurentian component (eastern Tugaloo terrane) and an internal terrane (Cat Square) that contains both Laurentian and Gondwanan detrital zircons, separated by the Brindle Creek fault. Cat Square terrane rocks likely accumulated in a Devonian remnant ocean that closed beginning c. 400 Ma. The complex but consistently asymmetric, NW- to west- to SW-directed flow pattern throughout the IP reflects confinement beneath a > 15 km thick overburden produced during subduction of Cat Square and Laurentian components beneath the approaching Carolina superterrane along the Central Piedmont suture. Oblique NE-to-SW transpressive subduction to > 15 km depth initiated partial melting, forcing escape from the collision zone in an along-strike orogenic channel. The IP detached from rocks to the west of the mid-Palaeozoic Brevard fault zone as the collision zone tightened and the IP mass flowed c. 200 km southwestward in the channel. The top of the channel is preserved at the NE end of the IP, and the base (Brevard fault zone) is preserved to the west and SW. As an exhumed orogenic channel, the curved IP flow paths may provide insight for middle to lower crustal deformation and flow in modern orogens.

Application of the channel flow concept to the Himalayas originated with Nelson et al. (1996) and was further developed by Grujic et al. (i996, 2002), Clark & Royden (2000), and numerically modelled by Beaumont et al. (2001, 2004). INDEPTH seismic imaging of the top of Indian crust beneath southern Tibet (Nelson et al. 1996; Hauck et al. 1998) better defined the Indiabeneath-Asia crustal geometry that had been hypothesized for many years. It also revealed several 'bright spots' beneath the Tibetan Plateau that were interpreted as low velocity melt zones, suggesting magma exists today beneath the Tibetan Plateau. Clark & Royden (2000) and Burchfiel (2004) presented the case for complex flow beneath the eastern Tibetan Plateau based on real-time GPS data, which indicate a complex movement pattern at the present topographic surface. Gravity-driven orogenic channel flow explains four key elements of Himalayan tectonics: (1) synchronous but opposite senses of slip on the Main Central Thrust and South Tibetan Detachment; (2) recognition of a low-velocity middle

crust beneath Tibet; (3) down-dip projection of the Greater Himalayan sequence to low-velocity (molten) middle crust beneath Tibet; and (4) strain patterns (Hodges 2006). Yet another, perhaps more important, test for the channel flow hypothesis is whether or not it can be exported to other orogens. If the model is only applicable to the Himalaya, its importance for advancing our understanding of collisional tectonics becomes minimal. If it can be recognized in other orogens, however, a major step forward will have been made towards a better understanding of processes that affect crustal evolution. Beaumont et al. (2004, 2006) produced alternative models analogous to channel flow that might develop in other tectonic settings, recognizing that gravity-driven (Himalayan-type), tectonically driven and mixed types of channels are possible. Beaumont et al. (2004) suggested that natural channels are likely to be more complex than predicted by numerical models using homogeneous crust, and therefore are difficult to recognize. Application of the channel flow concept, which combines simple

From: LAW, R. D., SEARLE,M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 517-541. 0305-8719/06/$15.00 9 The Geological Society of London 2006.

518

R.D. HATCHER & A. J. MERSCHAT

shear and Poiseuille flow to orogens (Brunel 1986; England & Houseman 1989; Bird 1991; Mancktelow 1995; Nelson et al. 1996; Clark & Royden 2000; Beaumont et al. 2001, 2004) immediately raises two questions. (1) What would the hypothesized channel beneath the Himalaya and Tibetan Plateau look like if the upper crust were removed? (2) Are there places in the world where we can view an exhumed channel to better understand its internal components? The central gneiss belt in the southern Grenville province (Culshaw et al. 1983; Hanmer 1988; Culshaw et al. 1997) may be one case where a mid-crustal channel formed that involved strike-normal flow (Jamieson et al. 2004). Azcfirraga et al. (2002), however, have described strike-parallel sheath folds and other fault-related structures near the base of the Cabo Ortegal complex in Spain (Galicia). In contrast with other examples citing strike-normal flow, this example from Spain indicated strike-parallel flow. Our purpose here is to suggest that the Appalachian Inner Piedmont (IP) may represent a tectonically forced, orogenic strike-parallel channel. Recent tectonic models have hypothesized curved ductile flow in the IP at high metamorphic grade during the Devonian and early Mississippian (Merschat et al. 2005). Flow in the IP, instead of being across orogenic strike, was initially directed towards the north and NW, then deflected west and SW parallel to strike. This paper compares observations and data related to mid-Palaeozoic high-temperature ductile flow in the southern Appalachian IP to channel flow models, to determine if the IP may represent a type of exhumed orogenic channel. If a channel existed during high temperature deformation of the IP, it would have been a net along-strike channel wherein crustal material was extruded at the SW end of the Appalachian orogen. This may provide insight into the nature of ductile flow beneath the eastern part of the Tibetan Plateau where curved flow trajectories may be present in the middle crust. There are several alternative interpretations to the IP being an orogenic channel. (1) The IP is a stack of deep crustal, SW-vergent, fold-nappe thrust sheets. (2) Stacks of SW-vergent thrust sheets formed under an orogenic 'lid' (Laubscher 1988), or an infrastructure beneath a suprastructure (Wegmann 1935; Hailer 1956; Griffin 1971a). (3) The IP is a crustal-scale ductile shear zone (i.e. a shear zone that encompasses the entire IP thickness) that formed in the mid-crust (Davis et al. 1991; Hatcher 1993, 2002; Vauchez et al. 1993). Crustal-scale shear zones (that may have relatively small thickness, but deformed over a large region) have been identified in many places (e.g. Platt & Behrmann 1986; Azc~irraga et al. 2002), but these are not comparable in magnitude to the IP structures. (4) The IP may be a 'metamorphic core complex' (Coney 1980), like part of

the Miocene deformation in Alboran zone in the core of the Betic Cordillera, southeastern Spain (Mart/nez-Martinez et al. 1997), or a mid-crustal, subhorizontal, extensional shear zone comparable to those in the Archean Abitibi-Wawa orogen in the Superior province in Ontario (Moser et aL 1996). The latter formed at some 30 km beneath an undeformed suprastructure. This would require that the high temperature IP fabrics be extensional; no mesoscopic extensional shear-sense indicators or map-scale structures have been recognized to date to support this alternative hypothesis. The IP fits alternatives 1, 2 and 3, except the strike-parallel vergence cannot be easily explained with these models. We also consider the IP a shear zone that occupies the entire thickness of the IP (Davis et al. 1991; Hatcher 2001). A second purpose is to suggest a unified model for IP structure and other IP attributes.

Tectonic setting The IP extends some 700 km along-strike from Winston-Salem, North Carolina, southwestward to the Coastal Plain in Alabama (Fig. 1). It has been long recognized for its high metamorphic grade and contrasting structural style with adjacent terranes (King 1955; Bentley & Neathery 1970). The IP is composite and consists of the eastern Tugaloo (western IP) and Cat Square terranes, separated by the Brindle Creek fault (Fig. 1). Eastern Blue Ridge (western Tugaloo terrane) rocks west of the Brevard fault cannot be separated stratigraphically from the rocks in the western IP (eastern Tugaloo terrane), however, because the stratigraphic sequences on either side of this fault are the same (Fig. 2). The IP is multiply deformed with shallow-dipping meso- and macroscale structures that contrast with those in adjacent terranes. A gently dipping stack of large, crystalline, Type F thrust sheets (formed by plastic excision of the common limb between a recumbent or reclined antiform and synform-fold nappes) (Hatcher & Hooper 1992; Hatcher 2004), shallow-dipping foliations, map-scale sheath folds, curved mineral stretching lineation pattern, dominance of sillimanite-grade rocks and migmatite, and a long, hot thermal history, form the attributes of mid-Palaeozoic mid-crustal flow in the IP. Northwest of the Tugaloo terrane are three terranes in the central Blue Ridge: the Dahlonega gold belt, and Cowrock and Cartoogechaye terranes composed of medium to high-grade metasedimentary, metavolcanic and ultramafic rocks (Fig. 1). Farther west is the Laurentian margin consisting of Grenvillian 1.1 Ga and older basement with a cover of rifted-margin sediments. The IP is bordered in the Carolinas, Georgia and Alabama to the NW by the Brevard fault zone and to the

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

519

~~ ,~d o o

9

~o o

o~ o,-~ o ~ [--, ~

o '~

o

N.a 9

~ ~.~

2~ ~o

Z-~o~ ~.~

~.~

o ~ ~_

~

.~_~,~

o o =.~ =

~.~.~.

a=

~=~~

520

R.D. HATCHER & A. J. MERSCHAT

Fig. 2. Tugaloo and Cat Square terranes lithologic components. (a) Tugaloo terrane sequence showing several Ordovician plutons (purple), possible oceanic and fragmented continental crust upon which the sequences were deposited, Middle Ordovician unconformity, and trajectories of faults that would later propagate through the sequences. (b) Cat Square terrane sequence showing the anatectic Toluca (light blue) and Walker Top (dark blue) granites, and the future trajectory of the Brindle Creek fault. A, away; T, towards. Fig. 3. Pattern of dominant (52) foliation and lineation in the northern Inner Piedmont. (a) Histogram of 4587 dip measurements of dominant foliation in the northern Inner Piedmont. (b) Form-line map of $2 foliation in the northern Inner Piedmont. Form lines are parallel to strike; teeth on form lines indicate dip direction. Density of form lines indicates density of data coverage used in map compilation. Based on 4587 measurements from Hadley & Nelson (1971), Rankin et al. (1972), Espenshade et al. (1975), Heyn (1984), Goldsmith et al. (1988), Nelson et al. (1998) and R. D. Hatcher (unpublished data). New data collected after this map was compiled reinforce the broad patterns depicted in this map. (c) Histogram of 764 mostly mineral elongation lineations in the northern IP. Note the dominance of gentle plunges. (d) Distribution of measured lineations (filtered to create cartographic spacing) illustrating the vortex-like pattern in the northern IP. Arrowhead indicates direction of plunge; arrowhead on both ends of fine indicates horizontal lineation. Line on each measurement indicates trend. Sources of lineation data are Grant (1958), Hadley & Nelson (1971), Lemmon & Dunn (1973a, b), Griffin (1974a), Whisonant (1979), Conley & Drummond (1981), Goldsmith (1981), Hatcher & Acker (1984), Heyn (1984), Hopson (1984), Willis (1984), Dennis (1989), McConnell (1990), Liu (1991), Davis (1993a), Maybin (1995, 1997), Niewendorp & Maybin (1994a, b), Yanagihara (1994), Niewendorp (1995a, b, 1996, 1997), West (1996, 1997, unpublished data), Curl (1998), Nelson et al. (1998), J. M. Garihan (unpublished data), R.D. Hatcher, Jr (unpublished data), and W.M. Schwerdtner (unpublished data). From Merschat et al. (2005) Additional measurements added to this data set since the original compilation was made reinforce the curved pattern illustrated by the data (see Figs 7 & 12).

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

521

522

R.D. HATCHER & A. J. MERSCHAT

SE by the Central Piedmont suture (Fig. 1). East of the IP is the Peri-Gondwanan Carolina superterrane, a 500-650 Ma volcanic arc that was metamorphosed c. 550 Ma, intruded by 535 Ma plutons after it was amalgamated with another arc (Dennis & Wright 1997a; Hibbard et al. 2003), and collided with Laurentia during the middle Palaeozoic.

Attributes of the Inner Piedmont: facilitating conditions of mid-Palaeozoic ductile flow Rocks west of the Brevard fault zone in the eastern Blue Ridge (western Tugaloo terrane) contain the same stratigraphic sequence as western IP rocks, with the exception of Middle Ordovician metavolcanic and metasedimentary rocks in the western IP and Siluro-Devonian metasedimentary rocks in the Cat Square terrane (eastern IP; Fig. 2). The dominant assemblage consists of Neoproterozoic-Cambrian(?) deep-water, oceanic, sedimentary and volcanic rocks (now biotite paragneiss, pelitic and aluminous schist, and amphibolite), metamorphosed to middle- and upperamphibolite-facies assemblages. These rocks are overlain by Cambrian(?) metasiltstone, quartzite, graphitic schist and impure marble, then unconformably by Middle Ordovician metavolcanic and metasedimentary rocks (Fig. 2). The Middle Ordovician sequence consists of mafic volcanic rocks containing felsic laminae overlain by quartzitefelsic metatuff and relatively pure marble. The felsic metatuff from this sequence has been dated (sensitive high resolution ion microprobe, SHRIMP, U - P b ) at 4 5 9 + 4 and 445 + 4 Ma (Bream 2003) (Fig. 2). The Tugaloo terrane was intruded by Ordovician, Silurian and Devonian plutons, with some Carboniferous plutons near Atlanta, Georgia. A window of Tugaloo terrane rocks (NW in Fig. 1) occurs in the eastern Cat Square terrane in North Carolina, with rocks composed of biotite paragneiss, mafic and ultramafic rocks. These rocks may be relict early Palaeozoic ocean crust and mantle. The eastern IP (Cat Square terrane) was initially separated from the Tugaloo terrane on the basis

of SHRIMP detrital zircon age dates (Bream 2003); otherwise, the two assemblages can only be separated by relative abundances of different lithologies, and some unique sequences to the west. It is bounded to the west by the Brindle Creek fault and to the SE by the Central Piedmont suture (Fig. 1). Cat Square terrane rocks consist of mixed LaurentianGondwanan affinity Siluro-Devonian metasedimentary rocks dominated by biotite paragneiss and aluminous schist intruded by Late Devonian and younger anatectic granites, with minor quantities of metabasalt. The Cat Square terrane detrital zircon suite is dominated by Grenvillian 1.1 Ga and older zircons, with additional 600, 500 and 430 Ma zircons (Bream et al. 2001, 2004; Bream 2002, 2003). The 500-600 Ma zircons were probably shed from the approaching Carolina superterrane, while the older zircons were probably derived from Laurentia (Bream 2003; Merschat et aI. 2005). The 430 Ma zircons delimit the maximum age of Cat Square terrane rocks (Bream et al. 2004), but otherwise have an unknown source. Cat Square plutons are all late Devonian and younger and appear to be dominantly anatectic (Mapes 2002). The Brevard fault separates the eastern from the western Tugaloo terrane, but the Brevard played a key mechanical and kinematic role in the deformational history of the IP (Hatcher 2001; Merschat et al. 2005). A zone of strongly aligned foliations and lineations along the western flank of the IP defines the mid-Palaeozoic Brevard fault zone, which consists of a 15 to 20 km wide amphibolitegrade dextral shear zone that dips 10 to 45 ~ SE (Hatcher 2001). It is characterized by strongly aligned mylonitic foliation ($2) striking NE, dipping SE, and a subhorizontal N E - S W trending mineral lineation (L2), with a top-to-the-SW shear sense, and was active from at least 360 to 350 Ma (Davis et al. 1991; Davis 1993a, b; Vauchez et al. 1993; Hatcher 2001; Merschat et al. 2005). The primary characteristics of the IP that separate it from Blue Ridge rocks to the west (west of the Brevard fault) are the gentle dip of foliation throughout the IP (Fig. 3a & b), and the curved subhorizontal mineral lineation (Fig. 3c & d). Several studies have demonstrated that the IP is multiply deformed and the dominant structural elements

Fig. 4. (a) Three-dimensional (3D) block diagram depicting the structure of part of the northern IP from near Hendersonville, NC, to the Sauratown Mountains window. Red lines are map-scale sheath folds. Vertical exaggeration in (a) and (b) 1.3:1:1 (X:Y:Z). Towns: Hk, Hickory; Hv, Hendersonville; Ln, Lenoir; Mg, Morganton; Sh, Shelby; Wk, Wilkesboro; WS, Winston-Salem. (b) More detailed 3D block diagram of area in (a) showing major tectonic units. Trends on block surface drawn from lineations. BCF, Brindle Creek fault; BoCF, Bowens Creek fault; cps, Central Piedmont suture; ct, Carolina superterrane; MF, Marion fault; MSF, Mill Spring fault; Oh, Henderson Gneiss; RF, Ridgeway fault; RsF, Rosman fault; SMW, Sauratown Mountains window; SRA, Smith River allochthon. Town abbreviationsas in (a). (c) Flow model for the northern Inner Piedmont based on detailed geological mapping and lineation data of the kind displayed in Figure 3. Abbreviations as in (a) and (b).

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

523

524

R.D. HATCHER & A. J. MERSCHAT

are $2, L2 and F 2 (e.g. Hopson & Hatcher 1988; Hatcher 2001; Merschat et al. 2005). The dominant $2 foliation throughout the IP dips gently and becomes strongly orientated N E - S W along the western flank of the IP (Fig. 3a). $2 is defined by parallel alignment of phyllosilicates, quartz ribbons, and other non-equant minerals. Transposed compositional layering and new migmatitic layering parallel to $2 is axial planar to recumbent to reclined, isoclinal to tight, gently plunging F2 folds. The high-temperature mineral stretching lineation, L2 (see Merschat et al. 2005, fig. 8b) varies in degree of development with rock type and is defined by elongate sillimanite, hornblende, quartz rods, streaked muscovite or mantled feldspars. Throughout most of the IP, L2 is subhorizontal and coaxial with F2. Goldsmith (1981) was the first to recognize this arcuate pattern defined by L2 mineral lineations and F2 fold hinges. From SE to NW the pattern curves from north-south to N W - S E to east-west to N E - S W (Fig. 3c). As stated by Merschat et al. (2005), this mineral lineation, which yields the flow pattern critical to our interpretation, formed during peak metamorphism and has been dated (zircon rims) at 350-360 Ma (Bream et al. 2001; Bream 2002, 2003). Field and geochronologic relationships clearly indicate that this lineation formed during one event, and the curved pattern is not the product of overprinting during separate events. Griffin (1969, 1971a, b, 1974b) was the first to recognize the fold-nappe (Type F of Hatcher & Hooper 1992; Hatcher 2004) structural style in the IP, and concluded that these nappes are NWvergent. The westemmost of these thrust sheets in NE Georgia, South Carolina, and adjacent North Carolina - the Six Mile thrust sheet and Walhalla nappe (Griffin 1969, 1974b, fig. 13) - place migmatitic rocks at upper kyanite or sillimanite grade over lower-grade garnet- to staurolite-grade rocks of the Chauga belt (Hatcher 1972, 1978) (Fig. 1). Griffin (1974b) depicted the imbricate structure of the westem IP in South Carolina in a three-dimensional block model similar to our modem block model for the northem IP. We similarly interpret the thrust sheets as an infrastructure and the Carolina superterrane a suprastructure above the IP, as the Cat Square and Tugaloo terranes were subducted beneath the Carolina superterrane during midPalaeozoic time (Fig. 4; Merschat et al. 2005). Our block model differs from Griffin's principally in recognizing the SW vergence of thrust sheets and macroscale sheath folds. The Brevard fault zone constitutes the lower boundary to the IP, while the central Piedmont suture and obducted Carolina superterrane constitute the upper boundary. We infer that the Carolina superterrane collided obliquely and thus diachronously from north to

south with the IP so that flow in the migmatitic mass could escape towards the SW as it was buttressed against the underlying Brevard fault zone. Oblique, zippered NE-to-SW collision of Carolina terrane would imply that progressive younging of deformation and metamorphism should also occur. Modem age dates do not exist in western Georgia and Alabama to test this implication, but faults framing the east end of the Pine Mountain window have yielded 330 and 303 Ma ages (Student & Sinha 1992), and these faults truncate IP structures and mesofabrics. Another fundamental structural characteristic recently documented in the IP is both map- and meso-scale sheath folds (see Merschat et al. 2005, figs 8c, d & 11). Because of the small size of exposures throughout most of the IP, we have observed only a few large mesoscale sheath folds (e.g. Fig. 5a; Merschat et al. 2005, figs 8, 9 & 11), but smaller mesoscale sheath folds are fairly common. Sheath folds that have been observed mostly have SW-directed transport confirmed independently by numerous shear-sense indicators and map patterns (Merschat et al. 2005). In addition, map-scale SW-directed sheath folds exist in the western IP in South Carolina, and were mapped as SW-directed thrusts that repeat the Chauga belt sequences (see Merschat et al. 2005, fig. 9). Ideally, if the IP is an orogenic channel, the base of the channel should record top-to-the-SW shear sense and the top should record top-to-the-NE shear sense. Detailed geological mapping by Heyn (1984) along the SW margin of the Sauratown Mountains window in North Carolina (Fig. 1), where it meets the lower boundary of the IP, reveals top-to-the-SW shear sense. Similarly, shear-sense indicators at the SW end of the IP reveal top-to-the-SW transport (Steltenpohl et al. 1990; Steltenpohl 2005). Detailed geological mapping and collection of mesofabric data (R. D. Hatcher unpublished) along the NW flank of the Sauratown Mountains window in North Carolina, however, reveals top-to-the-NE shear sense along the fault at the base of the overlying Smith River allochthon. The combination of the gentle dip of foliation and shallow plunge of the mineral stretching lineation in the IP led us to construct a model for flow with vortex-like trajectories that range from NW to north along the Central Piedmont suture, to west-directed in the central IP, then strongly SWdirected from several kilometres SE of the Brevard fault zone buttress (Fig. 6). New mineral lineation data from the Atlanta area (Higgins et al. 2003) (Fig. 7) reproduce the same pattern that we have recognized farther north in NE Georgia (Fig. 8), but data from the IP in Alabama available only in fabric diagrams (Bentley & Neathery 1970; Neilson 1988; Fig. 9) reveal a

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

525

Fig. 5. Folded stromatic migmatite in saprolite exposures in northwestern South Carolina. (a) NW-vergent F3 folds (or NW limb of a larger sheath fold) on U.S. 762 km west of Westminster, SC. Cut is c. 6 m high; view is toward the NE. Pegmatite related to decompression melting(?) along a small fault (note drag folds) cuts folds in the westernmost part of the exposure. Insets, with hammer for scale, show details of structures. Lenticular amphibolite layers probably are modified elliptical or anvil sections through mesoscopic sheath folds. (b) Isoclinal recumbent F2 folds near Coneross Creek c. 4 km north of Townville. Exposure is c. 3 m high; view is towards the NE. Inset shows complete transposition in hinge of one of the folds. Plunge is < 5 ~ south. Dark-coloured layers are weathered amphibolite; light-coloured layers are weathered granitoid melt. Fold hinges in both (a) and (b) are nearly normal to the road-cuts. strongly N E - S W alignment. SW-directed flow also propagated into the eastern Blue Ridge N W o f the Brevard fault zone, because the lineation in this area also trends N E - S W and plunges

gently, with shear sense indicators again indicating SW-directed transport (Figs 7 & 8). This overprint into the eastern Blue Ridge, however, gives w a y to contrasting Blue Ridge structural regimes a few

526

R.D. HATCHER & A. J. MERSCHAT

Fig. 6. (a) Flow model for the southern Appalachian Inner Piedmont based on available data. (b) Same base map as in (a) with shear-sense data, where they are known. Arrowheads indicate top direction. Red arrows are shear-sense indicators formed during the Neoacadian event (350-360 Ma). Purple arrows in South Carolina are high temperature shear-sense indicators that could have formed during either the Neoacadian or early Alleghanian (from West 1997). Blue arrows in western Georgia represent high temperature Alleghanian shear-sense indicators (from Hooper et al. 1997, and references therein). Each arrow represents at least ten measurements, most represent several tens of measurements. Measurements in Alabama are from Steltenpohl et al. (1990), Steltenpohl (2005) and Hatcher (unpublished data). Measurements in northern Georgia are from Higgins et al. (2003). Measurements in northeastern Georgia and the Carolinas are from data in, or cited in, Hopson & Hatcher (1988), Merschat et al. (2005), W. M. Schwerdtner (unpublished) and R. D. Hatcher (unpublished).

kilometres west of the Brevard fault zone (e.g. Hatcher et al. 2004; fig. 8). Metamorphic isograd maps of the IP (Fig. 10) reveal a metamorphic core of pervasively migmatitic sillimanite and higher-grade rocks throughout the central IP, decreasing to kyanite grade in the southeastern IP in North Carolina, and to kyanite, staurolite or garnet grade in the western IP in N W South Carolina and NE Georgia (Fig. 10). Recent estimates of P-T conditions during peak metamorphism by Mirante & Patino-Douce (2000), Bier (2001), Bier e t al. (2002) and Merschat (2003) indicate the core of the Inner Piedmont consistently reached temperatures of 750-850~ and pressures of 5 0 0 - 8 0 0 MPa. The IP thus reached sillimanite I, sillimanite II, and possibly hornblende-granulite-facies conditions. Classic migmatite structures described in other orogens (e.g. Mehnert 1968) are abundant here: stromatic migmatite is most common (Fig. 5), but agmatite (Fig. 11), diatexite, and other varieties are

also present (Williams 2000). In addition, Cat Square terrane granitoids mostly have an anatectic origin, including the 378 Ma Toluca and 366 Ma Walker Top Granite bodies (Giorgis et al. 2002; Mapes 2002). A phenomenon first noted by Griffin (1969) immediately beneath the Six Mile thrust sheet in northwestern South Carolina is a zone of excessive melting producing what might be called 'super migmatite'. A zone of much more extensive melting and super migmatite occurs beneath the Brindle Creek thrust sheet in North and South Carolina (Giorgis 1999; Williams 2000; Bier et al. 2002; Merschat e t al. 2005). Metamorphism of IP rocks was a protracted event that spanned 30 million years, bracketed by intrusion of the Toluca Granite (378 Ma) and reported metamorphic ages (365, 345 and 330 Ma; Fig. 12). Mapes (2002) concluded from geochemical-tectonic discriminant analysis that both the Toluca and Walker Top Granites are anatectic

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

527

Fig. 7. Mineral stretching lineation and tight fold pattern in the eastem Blue Ridge and western Inner Piedmont near Atlanta, Georgia (after Higgins et al. 2003). (a) Frequency distribution of plunges of 489 mineral stretching lineations and hinges of tight folds in the Inner Piedmont. (b) Frequency distribution of plunges of 153 mineral stretching lineations and hinges of tight folds in the eastern Blue Ridge (western Tugaloo terrane). (c) Map of eastern Blue Ridge (blue) and Inner Piedmont (red) mineral stretching lineations and tight folds. Arrows point in direction of plunge. Double-ended arrows indicate horizontal lineations. Late Carboniferous plutons (little internal deformation):bh, Ben Hill; sm, Stone Mountain; pal, Palisades; uc, Union City. Circle is approximate location of Atlanta.

melts of Cat Square terrane metasedimentary rocks. Calculated zircon saturation temperatures (Mapes 2002) indicate Cat Square terrane rocks reached 800~ or greater by c. 378 Ma. Bream (2003) reported three metamorphic peaks at 365, 345 and 330 Ma from U - P b ion microprobe ages of zircon rims in the northern IP and eastern Blue Ridge in the Carolinas and Georgia. Dennis & Wright (1997a) reported 360 Ma and 320 Ma U - P b monazite ages from five samples across the South Carolina IP. Kish (1997) reported a U - P b age of 357-348 Ma for the Cat Square charnockite. Biotite and hornblende 40Ar]39Arplateau ages from the Georgia IP suggest that high-grade metamorphism and a thermal overprint in the Carolina superterrane occurred at c. 350-360 Ma (Dallmeyer et al. 1986; Dallmeyer 1989). Overlapping metamorphic mineral ages and P-T estimates suggest the IP

remained hot until at least 345 Ma, cooled, and was reheated to high grade again 330-320 Ma (Fig. 12). Late Silurian (c. 415 Ma) magmatism in the Carolina superterrane (Samson & Secor 2000) not only indicates that the superterrane had been assembled, but it may mark the initiation of B subduction of Rheic ocean crust prior to A subduction of the Tugaloo terrane. Mechanical components of the IP thus include the Mid-Palaeozoic Brevard fault and Central Piedmont suture (upper and lower boundaries), plastic rheology of the lithologic assemblage of quartzofeldspathic, pelitic, Ordovician granitoids (Tugaloo terrane), volcanic rocks, and minor quartzite and ultramafic rocks; free water from relict and metamorphic reactions; migmatite; and anatectic melt-generated plutons. Metamorphic mineral assemblages require burial depths of 15 to 18 km

528

R.D. HATCHER & A. J. MERSCHAT

Fig. 8. (a) Mineral stretching lineations and hinges of tight folds in the western Tugaloo terrane (eastern Blue Ridge) of northeastern Georgia and northwestern South Carolina. Note that lineations are strongly aligned NE-SW for a short distance NW of the Brevard fault zone, but become aligned with Blue Ridge structures farther west. Arrows point in plunge direction, with number indicating plunge value; double-ended arrows indicate horizontal lineations. (b) Frequency distribution of plunge of 200 mineral stretching lineations and hinges of tight folds. Note that >50% plunge 7%; Rosenberg & Handy 2005), numerous intriguing possibilities arise for understanding the behaviour of migmatite terranes. The Appalachian IP doubtlessly contained sufficient melt to have become mobilized in a viscous state during the mid-Palaeozoic (Figs 5 & 11), producing the complex flow pattem we have identified (Hatcher 2001; Merschat et al. 2005). It was,

530

R . D . H A T C H E R & A. J. M E R S C H A T

r,~ ,d

~ o ~

o~

o.~

~o~

o

9~

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

531

Fig. 11. Agmatite (blocky migmatite) in saprolite 1 km east of Old Pickens, South Carolina. Exposure is c. 3 m high; view is towards the north. Reddish-brown material (palaeosome) is weathered amphibolite and biotite gneiss (palaeosome); light-coloured material is granitic leucosome (neosome).

Fig. 12. Tectonothermal time line for the Inner Piedmont. Data sources indicated by superscript numbers: 1, Bream (2002); 2, Bream (2003); 3, Bream et aL (2001); 4, Bream et al. (2004); 5, McSween et al. (1984, 1991); 6, Carrigan et al. (2001); 7, Davis (1993a); 8, Dennis & Wright (1997a, b); 9, Giorgis et aL (2002); 10, Luthet al. (1964); 11, Mapes (2002); 12, Mapes et al. (2001); 13, Mirante & Patino-Douce (2000).

532

R . D . HATCHER & A. J. MERSCHAT

,..-,.,~

~

~"

~

I

~

0

I:'~ ~

,~

~ '''"

~ .= ,.~ "~ ~ .~ ,-~ ~,.~

0

=

~ . ~ ,.-.' ~

~

"

~

~ .~-~ ~

~: ~ ,-~

,~

=

~

,=~u

=,.~,~

~

~

c~

o

~ ' =

~I 0

. ~ ~o

.

=

~

~

o~.~

~-~ ~

~ .

~.,~

s

-~

.~.~

~

0

~

I~ I ~ , , -

b _ o ='-~~

~ - ~

0 ~

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT

533

Fig. 14. Relationships in three dimensions between a stack of four hot imbricate Type F thrust sheets that are deflected SW by confinement beneath the overriding Carolina superterrane (removed) and the mid-Palaeozoic Brevard fault zone lower boundary. The lowest sheet (in gold) may not have been deflected, but is forced SW by confinement between the overlying thrust sheets and the underlying Brevard fault zone. however, tectonically driven rather than gravity driven, involving a viscosity > 1019 Pa s, as in one of Beaumont et al.'s (2006) models of large hot orogens. The complex flow pattern observed in the Appalachian IP (Fig. 6; Merschat et al. 2005) raises the possibility that flow may have been initially confined from the NE and less confined towards the SW because of oblique subduction, producing a channel composed of stacks of SWdirected imbricate thrusts that were directed along-strike into what is now Alabama. If so, this provides an opportunity to observe and better understand an exhumed along-strike channel at various depths of erosion. Portions near the upper boundary of the proposed channel are exposed at the NE end of the IP immediately NW of the Sauratown Mountains window (Fig. 1). The SW end in Alabama represents the lowest part of the channel, as suggested by the strong N E - S W alignment of the mineral stretching lineation across the IP (Fig. 3). With that possibility, we may be able to better understand the process of generation of a SW-directed, tectonically driven mid-Palaeozoic channel. If the Appalachian IP is an exhumed mid-Palaeozoic channel consisting of numerous Type F imbricates, it compares favourably

with the 2D Beaumont et al. (2004, fig. 14) model H T - H E T in which they introduced initial anisotropies and produced a series of imbricates as it deformed over a simulated 20 million year time span (Fig. 13). Deformation in the IP, however, occurred in a state of 3D heterogeneous strain, and the migmatitic IP became an along-strike channel with SW-directed flow, a product of oblique collision between Carolina superterrane to the east and various components of the remnant Theic ocean and the Laurentian margin to the west. These were subducted beneath Carolina to depths sufficient to melt and begin detachment and southwestward transport in a channel (Fig. 14). Carolina superterrane likely began to subduct the remnant Theic(?) ocean, then Cat Square terrane, followed by the Tugaloo terrane, during the Late Devonian (c. 408 Ma; Fig. 12). By c. 380 Ma anatectic melting produced the Toluca granitoids, and by 365 to 355Ma Carolina had subducted the Tugaloo and Cat Square terranes to depths of 1 5 - 2 0 k m , indicated by metamorphic mineral assemblages (Figs 10 & 12). Anatectic melting produced the Walker Top granitoid suite at 366 ___4 Ma (Giorgis et al. 2002; Mapes 2002). Water was

534

R.D. HATCHER & A. J. MERSCHAT

Fig. 15. Composite cross-section cartoon (a) showing the relationships between different components of the Inner Piedmont channel (b) now exposed at different erosion levels in different parts of the Inner Piedmont and adjacent terranes. Purple, western Tugaloo terrane (Chattahoochee thrust sheet) west of and below the mid-Palaeozoic Brevard fault zone; lavender and blue, eastern Tugaloo and Cat Square terranes, the components of the Inner Piedmont along-strike channel; light green, Carolina superterrane and Smith River allochthon (SRA), Peri-Gondwanan overriding plate. A, away; T, towards.

released from hydrous minerals via dehydration reactions (e.g. muscovite breakdown), as the threshold of wholesale migmatization was reached at upper-kyanite- to sillimanite-I-zone conditions from 360 to 355 Ma (Dennis & Wright 1997a; Bream 2003). The flow regime developed producing the Type F imbricate thrust stack at or near peak metamorphic conditions, confined above by the Carolina superterrane suprastructure and detaching below on the mid-Palaeozoic Brevard fault zone. Initial melting and southwestward escape of IP nappes occurred over a 25 to 30 million year time span. Cat Square and Tugaloo terrane rocks cooled to temperatures below the blocking temperatures of monazite and zircon and remained cool for the next 30 million years; then they were heated again between 325 and 320 Ma (Dennis & Wright 1997b; Bream 2003) during the early Alleghanian. The oldest direct thermal connection between the Carolina superterrane suprastructure and the Cat Square

and Tugaloo terrane infrastructure consists of the 420 to 380 Ma Silurian-Devonian plutonic suite that occurs in all of these terranes (McSween et al. 1984, 1991; Miller et al. 2000; Samson & Secor 2000). In addition to the plutons, a 360 Ma thermal overprint is also recorded in 4~ plateau ages in the western Carolina superterrane (Dallmeyer et al. 1986). Monazite ages from the Smith River allochthon (Fig. 1) link it to the periGondwanan Carolina superterrane (Hibbard et al. 2003) and support the hypothesis that the Carolina superterrane overrode the IP. Significant differences exist between our proposed IP channel and the Tertiary to Holocene channel hypothesized beneath Tibet and extruding between the Himalayan Main Central thrust and the South Tibetan Detachment. Flow in the IP is curved, instead of involving up-dip and acrossstrike extrusion, as in the Himalayas, so it was directed southwestward subparallel to strike, consistent with oblique collision. Abundant shear

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT Table 1. Hypothetical subduction zone parameters required f o r different dip*

Dip (degree) 20 10 5

tsubduction initiation

tmelting (~

Xsubduction (km)

Subduction direction

385 387 393

383 382 383

52.6 104 206.5

Head-on Oblique Oblique

* Subduction rate = 2 on/a; required (estimated) depth from metamorphic assemblages = 18 km; required depth for minimum melting = c. 7 kin; Toluca Granite crystallized c. 380 Ma. t = time. x = distance.

sense data throughout the central and western IP document the observed flow pattern (Merschat et al. 2005). Ideally, the Central Piedmont suture, by analogy with the South Tibetan Detachment, should be a normal fault if the IP were gravitydriven. Evidence of down-to-the-SE transport along the Central Piedmont suture has not been observed (e.g. West 1998), and is clearly dextral, not sinistral, in central Georgia east of the Pine Mountain window (Fig. 1; Hooper et al. 1997). There are, however, important similarities with the Himalayan-Tibetan channel. A prolonged thermal history occurred in the IP within c. 30 million years following burial and subduction of IP constituents beneath the Carolina superterrane beginning c. 400 Ma, with first melting at c. 385 Ma culminating with peak metamorphism, a meltweakened middle crest, and flow at 360 to 355 Ma (Table 1). Despite the differences cited above, the IP channel as defined herein compares favourably with the Beaumont et al. (2004, fig. 14) tectonically driven ~/> 1019 Pa s model H-HET, and the large, hot orogens model in Beaumont et al. (2006). It also fits most of the other criteria in terms of position in the crust, availability of water for extensive melting in the subducted mass, forming a series of thrust nappes that are west-, then SW-directed. The rheological state of melting and SW-directed flow in a tectonically driven channel had to have been induced by oblique subduction, burial and radioactive decay yielding the high temperatures and moderate pressures, coupled with available free water (both primary and from metamorphic reactions). The major conclusion from this discussion is that the Appalachian IP may be an exhumed tectonically forced orogenic channel, in the context of the Beaumont et al. (2006) large hot orogens concept. Tests of our hypothesis might consist of the following. (1) 3D finite-element or numerical models should be developed incorporating the physical parameters discussed above for the IP. Difficulties with this approach lie in the limited capabilities for 3D modelling with these techniques. (2) Additional modem geochronologic data should be

535

collected throughout the IP, but particularly in Georgia and Alabama, to additionally corroborate (or revise) the 350-360 Ma age of peak metamorphism and coeval formation of fabric elements at all scales. This would augment the data of Dennis & Wright (1997a, b), Bream et al. (2001), Mapes (2002) and Bream (2003). (3) Additional detailed field studies should be conducted along the Central Piedmont suture to collect new shear-sense data. Herein the difficulty lies in the reactivation of suitably orientated segments of the suture as a thrust (in South Carolina; West 1998), and in dextral strike-slip (in central Georgia; Hooper & Hatcher 1988), where detailed studies already exist. Even so, large areas of the IP and its boundaries remain mapped only by reconnaissance wherein little shear-sense data were collected. An average heat production of c. 2 ixW/m 3, like that assumed by Beaumont et al. (2004, 2006) in their models, would probably be appropriate for the IP. If we assume initial subduction was head-on, the subduction zone had a dip of 20 ~ and a subduction rate of 2 cm/a, subduction could have begun as recently as 385 Ma under hydrous conditions for the Toluca Granite to crystallize at 380 Ma. For reasons stated above, subduction more likely was oblique, so if the dip was 10 ~ subduction could have begun as late as 387 Ma, or if the dip was 5 ~ subduction could have begun as early as 393 Ma (Table 1). The latter possibility seems more reasonable in light of the post-430 Ma time of deposition of Cat Square terrane sediments, likely oblique subduction closing the Rheic ocean and the time needed for burial to reach melting temperature. The resulting tectonically driven IP channel would have been SW-directed parallel to the strike of the orogen, because of confinement beneath the overriding Carolina superterrane and buttressing by the underlying mid-Palaeozoic Brevard fault and western Tugaloo terrane to the west (Fig. 15). These formed a closed system above and below, which apparently remained open to the SW. Conclusions

(1) The IP is an exhumed, tectonically forced, strikeparallel orogenic channel that formed as a product of Late Devonian to early Mississippian (390-350 Ma) oblique subduction of previously assembled, dominantly Laurentian terranes beneath the Carolina superterrane. (2) This channel has an upper boundary consisting of the overlying Carolina superterrane, a lower boundary consisting of the mid-Palaeozoic Brevard fault zone, and less confinement towards the SW permitting along-strike, SW-directed extrusion of partially melted, mid-crustal material.

536

R.D. HATCHER & A. J. MERSCHAT

(3) Attributes that permit our primary conclusions are: contrasting structural style with adjacent terranes, prolonged thermal activity at middle- to upper-amphibolite-facies conditions, extensive migmatization, SW-directed hot Type F imbricate thrust sheets, curved flow, and gentle dip of meso- to macroscale planar structures. (4) Orogenic channels can form in transpressional regimes, but they produce more complex flow patterns than gravity-driven channels in head-on collision zones. (5) The IP provides unique insights into the complexities of an exhumed, tectonically driven orogenic channel. Because of the possible along-strike exposure of all levels of the channel, it is particularly important in understanding heterogeneous, non-plane channel flow at different structural levels.

Cabo Ortegal complex (NW Spain). Journal of Structural Geology, 24, 1971-1989. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & MEDVEDE, S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109. DOI: 10.1029/ 2003JB002809. BEAUMONT, C., NGUYEN, M. H., JAMIESON, R. A. & ELLIS, S. 2006. Crustal flow modes in large hot orogens. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications,

US National Science Foundation grants GA-1409, GA20321, EAR-810852, EAR-8206949, EAR-8417894, EAR-9004604 and EAR-9814800 supported a great deal of fieldwork in the Carolinas and NE Georgia Inner Piedmont by R.D.H. and graduate students during the 1970s, 1980s, 1990s and early 2000s. Additional support during the late 1960s and early 1970s was provided by the South Carolina Geological Survey, H. S. Johnson, Jr and N. K. Olson, state geologists. Fieldwork in the late 1990s and early 2000s has been supported by the EDMAP programme administered by the US Geological Survey. The University of Tennessee Science Alliance Center of Excellence has provided considerable support for R.D.H. and students since 1986. Chris Beaumont (Dalhousie University) has kindly shared the graphics of several of his finite-element models with us. He and his colleagues' models have provided additional impetus for attempting to apply the channel flow concept to the Appalachian Inner Piedmont. Sid Jones (Tennessee Department of Conservation, DOE Oversight Division) provided useful insight into the chemical engineering literature on transport processes, particularly plate flow. Reviews by J. Grocott (Kingston University), R. Law (Virginia Tech) and N. Culshaw (Dalhousie University) considerably improved the manuscript and are gratefully acknowledged. Additional comments by Djordje Grujic (Dalhousie University) were also useful and resulted in additional improvement of the paper. The authors, however, remain culpable for all errors of fact and interpretation. Finally, R.D.H. enjoyed numerous debates with Doug Nelson, to whom this volume is dedicated, about southern Appalachian tectonics prior to his 1NDEPTH success (e.g. Nelson 1988; Hatcher & Hooper 1988). These debates were at times heated, but we always remained on friendly terms - the way scientific debate should be conducted. Doug is missed.

BENTLEY, R. D. & NEATHERY, T. N. 1970. Geology of the Brevard zone and related rocks of the Inner Piedmont of Alabama. Alabama Geological Society, Eighth Annual Field Trip Guidebook. BIER, S. E. 2001. Geology of the southeastern South Mountains, North Carolina. MS thesis, University of Tennessee. BIER, S. E., BREAM, B. R. & GIORGIS, S. D. 2002. Inner Piedmont stratigraphy, metamorphism, and deformation in the Marion-South Mountains area, North Carolina. In: HATCHER, R. D. JR & BREAM, B. R. (eds) Inner Piedmont geology in the South Mountains-Blue Ridge Foothills and the southwestern Brushy Mountains, central-western North Carolina. Carolina Geological Society Guidebook, 65-100. BIRD, P. 1991. Lateral extrusion of lower crust from under high topography in the isostatic limit. Journal of Geophysical Research, 96, 1027510286. BREAM, B. R. 2002. The southern Appalachian Inner Piedmont: New perspectives based on recent detailed geologic mapping, Nd isotopic evidence, and zircon geochronology. In: HATCHER, R. D. JR & BREAM, B. R. (eds) Inner Piedmont geology in the South Mountains-Blue Ridge Foothills and the southwestern Brushy Mountains, central-western North Carolina. Carolina Geological Society Guidebook, 45-63. BREAM, B. R. 2003. Tectonic implications of geochronology and geochemistry of para- and orthogneisses from the southern Appalachian crystalline core. PhD dissertation, University of Tennessee. BREAM, B. R., HATCHER, R. D. JR., MILLER, C. F. & FULLAGAR, P. D. 2001. Geochemistry and provenance of Inner Piedmont paragneisses, NC and SC: Evidence for an internal terrane boundary? Geological Society of America Abstracts with Programs, 33, 65. BREAM, B. R., HATCHER, R. D. JR., MILLER, C. F. & FULLAGAR, P. D. 2004. Detrital zircon ages and Nd isotopic data from the southern Appalachian crystalline core, GA-SC-NC-TN: New provenance constraints for Laurentian margin paragneisses. In:

268, 91 - 145.

References

ACNESON, D. J. 1990. Elementary Fluid Dynamics. Oxford University Press, Oxford. AZCARRAGA, J., ABALOS, B. & CIL IBARGUCH1, J. I. 2002. On the relationship between kilometerscale sheath folds, ductile thrusts and minor structures in the basal high-pressure units of the

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT TOLLO, R. P., CORRIVEAU, L., MCLELLAND, J. & BARTHOLOMEW, M. J. (eds) Proterozoic Evolution of the Grenville orogen in North America. Geological Society of America, Memoir 197, 459-475. BRODKEY, R. S. & HERSHEY, H. C. 1988. Transport phenomena: A unified approach. McGraw-Hill, New York. BRUNEL, M. 1986. Ductile thrusting in the Himalayas: Shear sense criteria and stretching lineations. Tectonics, 5, 247-265. BURCHFmL, B. C. 2004. New technology; new geological challenges. GSA Today, 14(2), 4-10. BURCHFIEL, B. C. & ROYDEN, L. H. 1985. North-south extension within the convergent Himalayan region. Geology, 13, 679-682. BUTLER, J. R. 1991. Metamorphism. In: HORTON, J. W. JR & ZULLO, V. A. (eds) The Geology of the Carolinas. Carolina Geological Society 50th Anniversary Volume. University of Tennessee Press, Knoxville, 127-142. CARRIGAN, C. W., BREAM, B. R., MILLER, C. F. & HATCHER, R. D. JR. 2001. Ion Microprobe analyses of zircon rims from the eastern Blue Ridge and Inner Piedmont, NC-SC-GA. Geological Society of America Abstracts with Programs, 33(2), 7. CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. CONEY, P. J. 1980. Cordilleran metamorphic core complexes: An overview. In: CRITTENDEN, M. O. Jr, CONEY, P. J., & DAvIs, G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America, Memoir 153, 7-34. CONLEY, J. F. & DRUMMOND, K. M. 1981. Geologic map and mineral resources summary of the northeast 1/4 of the Marion quadrangle, North Carolina. North Carolina Geological Survey, GM 210-NE, scale 1/24,000. COSTAIN, C. K. & DECKER, E. R. 1987. Heat flow at the proposed Appalachian ultradeep core hole (ADCOH) site. In: HOPSON, J. L. 8r HATCHER, R. D. JR (eds) Appalachian ultradeep core hole project site investigation: Data set I. ADCOH Project Publication 5. University of Tennessee, Department of Geological Sciences, Studies in Geology, 17. COSTAIN, J. K., SPEER, J. A., GLOVER, L., PERRY, L., DASHEVSKY, S. & MCKINNEY, M. 1986. Heat flow in the Piedmont and Atlantic Coastal Plain of the southeastern United States. Journal of Geophysical Research, 91, 2123-2135. CULSHAW, N. G., DAVIDSON,A. & NADEAU, L. 1983. Structural subdivisions of the Grenville province in the Parry Sound-Algonquin region. Geological Survey of Canada Paper, 83-1B, 243-252. CULSHAW, N. G., JAMIESON,R. A., KETCHUM,J. W. F., WODICA, N., CORRIGAN, D. & REYNOLDS, P. H. 1997. Transect across the northwestern Grenville orogen, Georgian Bay, Ontario: Polystage convergence and extension in the lower orogenic crust. Tectonics, 16, 966-982. CURL, D. C. 1998. Stratigraphy and structure of Wellford and Reidville quadrangles in part of the

537

eastern Inner Piedmont, near Spartanburg, South Carolina. MS thesis, University of Tennessee. DALLMEYER, R. D. 1989. Late Palaeozoic thermal evolution of crystalline terranes within portions of the U.S., Appalachian orogen. In: HATCHER, R. D. JR, THOMAS, W. A. & VIELE, G. W. (eds) The Appalachian-Ouachita orogen in the United States. The Geology of North America F-2. Geological Society of America, Boulder Colorado, 417-444. DALLMEYER,R. D., WRIGHT, J. E., SECOR, D. T. JR. 8r SNOKE, A. W. 1986. Character of the Alleghanian orogeny in the southern Appalachians: Part II. Geochronological constraints on the tectonothermal evolution of the eastern Piedmont in South Carolina. Geological Society of America Bulletin, 97, 1382-1392. DAVIS, T. L. 1993a. Lithostratigraphy, structure, and metamorphism of a crystalline thrust terrane, western Inner Piedmont, North Carolina. PhD dissertation, University of Tennessee. DAVIS, T. L. 1993b. Geology of the Columbus Promontory, western Piedmont, North Carolina, southern Appalachians. In: HATCHER R. D. JR & DAVIS, T. L. (eds) Studies of Inner Piedmont Geology with a Focus on the Columbus Promontory. Carolina Geological Society Guidebook. North Carolina Geological Survey, 17-43. DAVIS, T. L., HATCHER, R. D. JR., LIU, A. & TABOR, J. R. 1991. Southern Appalachian western Inner Piedmont: A progressive crustal-scale shear zone. Geological Society of America Abstracts with Programs (Annual Meeting), 23, A138. DENNIS, A. J. 1989. Tectogenesis of an accreted terrane: the Carolina arc in the Palaeozoic. PhD dissertation, University of South Carolina. DENNIS, A. J. & WRIGHT, J. E. 1997a. The Carolina terrane in northwestern South Carolina, U.S.A.: Late Precambrian-Cambrian deformation and metamorphism in a peri-Gondwanan oceanic arc. Tectonics, 16, 460-473. DENNIS, A. J. & WRIGHT, J. E. 1997b. Middle and Late Palaeozoic monazite U - P b ages, Inner Piedmont, South Carolina. Geological Society of America Abstracts with Programs, 29(3), 12. ENGLAND, P. & HOUSEMAN, G. 1989. Extension during continental convergence, with application to the Tibetan Plateau. Journal of Geophysical Research, 94(B 12), 17,561-17,579. ESPENSHADE, G. H., RANKIN, D. W., SHAW, K. W. & NEUMAN, R. B. 1975. Geologic map of the east half of the Winston-Salem quadrangle, North CarolinaVirginia. US Geological Survey Map 1-709-B, scale 1:250,000. GIORGIS, S. D. 1999. Inner Piedmont geology of the northwestern South Mountains near Morganton, North Carolina. MS thesis, University of Tennessee. GIORGIS, S. D., MAPES, R. W. • BREAM, B. R. 2002. The Walker Top Granite: Acadian granitoid or eastern Inner Piedmont basement? In: HATCHER, R. D. JR & BREAM, B. R. (eds) Inner Piedmont geology in the South Mountains-Blue Ridge Foothills and the southwestern Brushy Mountains,

538

R.D. HATCHER & A. J. MERSCHAT

central-western North Carolina. Carolina Geological Society Guidebook, 33-43. GOLDSMITH, R. 1981. Structural patterns in the Inner Piedmont of the Charlotte and Winston-Salem 2-degree quadrangles, North Carolina and South Carolina. In: HORTON JR, J. W., BUTLER, J. R. & MILTON, D. M. (eds) Geological Investigations of the Kings Mountain Belt and Adjacent Areas in the Carolinas. Carolina Geological Society Field Trip Guidebook, 9-27. GOLDSMITH, R., MILTON, D. J. & HORTON, J. W. JR. 1988. Geologic map of the Charlotte 1-degree • 2-degree quadrangle, North Carolina and South Carolina. US Geological Survey Map 1-1251-E, scale 1:250,000. GRANT, W. n. 1958. The geology of Hart County, Georgia. Georgia State Division of Conservation, Geological Survey, Bulletin 67. GRIFFIN, V. S. JR. 1969. Migmatitic Inner Piedmont belt of northwestern South Carolina. South Carolina Division of Geology, Geologic Notes, 13, 87-104. GRIFFIN, V. S. JR. 1971a. Stockwork tectonics in the Appalachian Piedmont of South Carolina and Georgia. Geologischen Rundschau, 60, 868-886. GRI~IN, V. S. JR. 1971b. The Inner Piedmont belt of the southern crystalline Appalachians. Geological Society of America Bulletin, 82, 1885-1898. GRIFFIN, V.S. JR. 1974a. Geology of the Walhalla quadrangle, S.C. South Carolina Division of Geology, MS-19, scale 1:24,000. GRIFFIN, V. S. JR. 1974b. Analysis of the Piedmont in northwest South Carolina. Geological Society of America Bulletin, 85, 1123-1138. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KUNDIG, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan crystallines in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. GRUJIC, D., HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel: Insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. HADLEY, J. B. & NELSON, A. E. 1971. Geology of the Knoxville 1 • 2 degree quadrangle, North Carolina, Tennessee, and South Carolina. US Geological Survey Map 1-654, scale 1:250,000. HALLER, J. 1956. Probleme der Tiefentektonik: Bauformen im Migmatit-Stockwerk der ostgrtnl~indischen Kaledonien. Geologischen Rundschau, 45, 159-167. HANMER, S. 1988. Ductile thrusting at mid-crustal level, southwestern Grenville province. Canadian Journal of Earth Science, 25, 1049-1059. HATCHER, R. D. JR. 1972. Developmental model for the southern Appalachians. Geological Society of America Bulletin, 83, 2735-2760. HATCHER, R. D. JR. 1978. Tectonics of the western Piedmont and Blue Ridge: Review and speculation. American Journal of Science, 278, 276-304. HATCHER, R. D. JR. 1993. Perspective on the tectonics of the Inner Piedmont, southern Appalachians. In: HATCHER, R. D. JR & DAVIS, T. L. (eds) Studies of Inner Piedmont geology with a focus on the

Columbus Promontory. Carolina Geological Society Guidebook, North Carolina Geological Survey, 17-43. HATCHER, R. D. JR. 2001. Rheological partitioning during multiple reactivation of the Palaeozoic Brevard Fault Zone, Southern Appalachians, USA. In: HOLDSWORTH, R. E., STRACHAN, R. A., MACLOUGHLIN, J. F. & KNIPE, R. J. (eds) The Nature and Significance of Fault Zone Weakening. Geological Society, London, Special Publications, 186, 255-269. HATCHER, R. D. JR. 2002. An Inner Piedmont primer. In: HATCHER, R. D. JR & BREAM, B. R. (eds) Inner Piedmont geology in the South Mountains-Blue Ridge Foothills and the southwestern Brushy Mountains, central-western North Carolina. Carolina Geological Society Guidebook, 1-18. HATCHER, R. D. JR. 2004. Properties of thrusts and the upper bounds for the size of thrust sheets. In: MCCLAY, K. R. (ed.) Thrust Tectonics and Hydrocarbon Systems. American Association of Petroleum Geologists, Memoir 82, 18-29. HATCHER, R. D. JR & ACKER, L. L. 1984. Geology of the Salem quadrangle, South Carolina. South Carolina Geological Survey, MS-26, scale 1/24,000. HATCHER, R. D. Jr & GOLDBERG, S. A. 1991. The Blue Ridge Geologic Province. In: HORTON, J. W. Jr & ZULLO, V. A. (eds) The Geology of the Carolinas. Carolina Geological Society 50th Anniversary Volume, The University of Tennessee Press, Knoxville, 11-35. HATCHER, R. D. JR & HOOPER, R. J. 1992. Evolution of crystalline thrust sheets in the internal parts of mountain chains. In: MCCLAY, K. R. (ed.) Thrust Tectonics. Chapman and Hall, London, 217-234. HATCHER, R. D. JR, BREAM, B. R., MILLER, C. L., ECKERT, J. O. JR. FULLAGAR, P. D. & CARRIGAN, C. W. 2004. Palaeozoic Structure of Southern Appalachian Blue Ridge Grenvillian Internal Basement Massifs. In: TOLLO, R. P., CORRIVEAU, L., MCLELLAND, J. & BARTHOLOMEW, M. J. (eds) Proterozoic Evolution of the Grenville Orogen in North America. Geological Society of America, Memoir 197, 525-547. HAUCK, M. L., NELSON, K. D., BROWN, L. D., ZHAO, W. & Ross, A. R. 1998. Crustal structure of the Himalayan orogen at c. 90 ~ east longitude from Project INDEPTH deep reflection profiles. Tectonics, 17, 481-500. HEYN, T. 1984. Stratigraphic and structural relationships along the southwestern flank of the Sauratown Mountains anticlinorium. MS thesis, University of South Carolina. HIBBARD, J. P., TRACY, R. J. & HENIKA, W. S. 2003. Smith River allochthon: A southern Appalachian peri-Gondwanan terrane emplaced directly on Laurentia? Geology, 31, 215-218. HIGGINS, M. W., CRAWFORD, T. J., ATKINS, R. L. & CRAWFORD, R. F. 2003. Geologic map of the Atlanta 30' • 60' quadrangle, Georgia. US Geological Survey Geologic Investigations Series Map 1-2062, scale 1:100,000. HODGES, K.V. 2006. A synthesis of the channel flow extrusion hypothesis as developed for the

CHANNEL FLOW 1N THE APPALACHIAN PIEDMONT Himalayan-Tibetan orogenic system. In: LAW, R. D., SEARLE, M.P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 71-90. HOOPER, R. J. & HATCHER, R. D. JR. 1988. The Pine Mountain terrane, a complex window in the Georgia and Alabama Piedmont: Evidence from the eastern termination: Reply. Geology, 16, 1049-1050. HOOPER, R. J. & HATCHER, R. D. JR. 1992. The origin of the Pine Mountain window: An internal basement massif within the Piedmont of Alabama and Georgia. In: MASON, R. (ed.) International Basement Tectonics Association. Basement Tectonics 7. Kiuwer Academic, Dordrecht, 375-383. HOOPER, R. J., HATCHER, R. D., JR., TROYER, P. K., DAWSON, R. J. & REDMOND, C. G. 1997. The character of the Avalon terrane in central Georgia. In: CLOVER, L. III & GATES, A. E. (eds) Central and Southern Appalachian Sutures: Results of the EDGE Project and Related Studies. Geological Society of America, Special Paper 314, 1 - 14. HoPSON, J. L. 1984. Stratigraphy and structure of the Alto allochthon, Ayersville quadrangle, Georgia. MS thesis, University of South Carolina. HoPSON, J. L. 1994. Structure of part of the Blue Ridge, northeastern Georgia and North Carolina. PhD dissertation, University of Tennessee. HOPSON, J. L. & HATCHER, R. D. JR. 1988. Structural and stratigraphic setting of the Alto allochthon, northeast Georgia. Geological Society of America Bulletin, 100, 339-350. JAMIESON, R. A., CULSHAW, N., BEAUMONT, C., NGUYEN, M. H. & SLAGSTAD, T. 2004. Hot nappes and lumpy channels: Mid-crustal flow modes in the Western Grenville orogen. In: SEARLE, M. P., LAW, R. D. & GODIN, L. (convenors) Channel flow, ductile extrusion and exhumation of lower-mid crust in continental collision zones, Abstract Book. Geological Society, London. JAMIESON, R. A., BEAUMONT, C. & MEDVEDEV, S. & NGUYEN, M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109. DOI: 10.1029/2003JB00002811. KING, P. B. 1955. A geologic section across the southern Appalachians: An outline to the geology in the segment in Tennessee, North Carolina, and South Carolina. In: RUSSELL, R. J. (ed.) Guides to Southeastern Geology. Geological Society of America, New York, 332-373. KISH, S. 1997. The Cat Square charnockite--a Palaeozoic charnockite in the Inner Piedmont of North Carolina. Geological Society of America Abstracts with Programs, 29(3), 28. LAUBSCHER, H. P. 1988. Material balance in Alpine orogeny. Geological Society of America Bulletin, 100, 1313-1328. LEMMON, R. E. & DUNN, D. E. 1973a. Geologic map and mineral resources of the Bat Cave quadrangle, North Carolina. North Carolina Geological Survey, GM 202-NW, scale 1/24000.

539

LEMMON, R. E. & DUNN, D. E. 1973b. Geologic map and mineral resources of the Fruitland quadrangle, North Carolina. North Carolina Geological Survey, GM 202-NW, scale 1/24000. LIu, A. 1991. Structural geology and deformation history of the Brevard fault zone, Chauga belt, and Inner Piedmont, northwestern South Carolina and adjacent areas. PhD dissertation, University of Tennessee. LUTH, W. D., JAHNS, R. H. & TUTTLE, O. F. 1964. The granite system at pressures of 4 to 10 kilobars. Journal of Geophysical Research, 69, 759-773. MCCONNELL, K. I. 1990. Geology and geochronology of the Sauratown Mountains anticlinorium, northwestern North Carolina. PhD dissertation, University of South Carolina. MCSWEEN, H. Y. JR, SANDO, T. W., CLARK, S. R., HARDEN, J. T. & STRANGE, E. A. 1984. The gabbro-metagabbro association of the southern Appalachian Piedmont. American Journal of Science, 284, 437-461. MCSWEEN, H. Y. JR, SPEER, J. A. & FULLAGAR, P. D. 1991. 7. Plutonic rocks. In: HORTON, J. W. JR & ZULLO, V. A. (eds) The Geology of the Carolinas. Carolina Geological Society 50th Anniversary Volume. The University of Tennessee Press, Knoxville, 109-126. MANCKTELOW, N. S. 1995. Nonlithostatic pressure during sediment subduction and the development and exhumation of high pressure metamorphic rocks. Journal of Geophysical Research, 100, 571-583. MAPES, R. W. 2002. Geochemistry and geochronology of mid-Palaeozoic granitic plutonism in the southern Appalachian Piedmont terrane, North Carolina-South Carolina-Georgia. MS thesis, Vanderbilt University. MAPES, R. W., MILLER, C. F., FULLAGAR, P. D. & BREAM, B. R. 2001. Nature and origin of Acadian plutonism, Piedmont terane, NC-GA. Geological Society of America Abstracts with Programs, 33(6), 92. MART~NEZ-MART~NEZ,J. M., SOTO, J. I. & BALANYA 1997. Crustal decoupling and intracrustal flow beneath domal exhumed core complexes, Betics (SE Spain). Terra Nova, 9, 223-227. MAYBIN, A. H. III. 1995. Geologic map of the Simpsonville, Quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 92, scale 1:24,000. MAYBIN, A. H. III. 1997. Geologic map of the Pelham quadrangle, South Carolina. South Carolina Geological Survey, Open File Map, scale 1/24,000. MEHNERT, K. R. 1968. Migmatites and the Origin of Granitic Rocks. Elsevier, Amsterdam. MERSCHAT, A. J. 2003. Inner Piedmont tectonics in the southwestern Brushy Mountains, North Carolina: Field and laboratory data revealing 3-D crustal flow and sillimanite I and H metamorphism. MS thesis, University of Tennessee. MERSCHAT, A. J., HATCHER, R. D. JR. • DAVIS, T. L. 2005. The northern Inner Piedmont, southern Appalachians, USA: Kinematics of transpression

540

R.D. HATCHER & A. J. MERSCHAT

and SW-directed mid-crustal flow. Journal of Structural Geology, 27, 1252-1281. MILLER, C. F., HATCHER, R. D., JR., AYERS, J. C., COATI4, C. D. & HARRISON, T. M. 2000. Age and zircon inheritance of eastern Blue Ridge plutons, southwestern North Carolina and Northeastern Georgia, with implications for magma history and evolution of the southern Appalachian orogen. American Journal of Science, 300, 142-172. MIRANTE, D. C. & PATINO-DOUCE, A. E. 2000. Melting and migmatization in the southern Appalachian Inner Piedmont of northeast Georgia; the Athens gneiss. Geological Society of America Abstracts with Programs, 33, 297. MOSER, D. E., HEAMAN, L. M., KROGH, T E. & HANES, J. A. 1996. Intracrustal extension of an Archean orogen revealed using single-grain U - P b geochronology. Tectonics, 15, 1093-1109. NEILSON, M. J. 1988. The structure and stratigraphy of the Tallassee Synform, Dadeville Belt, Alabama. Southeastern Geology, 29, 41-50. NELSON, A. E., HORTON, J. W. & CLARKE, J. W. 1998. Geologic map of the Greenville 1~ x 2 ~quadrangle, Georgia, South Carolina, and North Carolina. US Geological Survey, Map 1-2175, scale 1:250,000. NELSON, K. D. 1988. The Pine Mountain terrane, a complex window in the Georgia and Alabama Piedmont: Evidence from the eastern termination: Comment. Geology, 16, 1049. NELSON, K. D., ZHAO, W., BROWN, L. D. ETAL. 1996. Partially molten middle crust beneath southern Tibet; synthesis of Project INDEPTH results. Science, 274, 1684-1688. NIEWENDORP, C. A. 1995a. Geology of the Bush River quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 84, scale 1/24,000. NIEWENDORP, C. A. 1995b. Geology of the Joanna quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 85, scale 1/24,000. NIEWENDORP, C. A. 1996. Geology of the Laurens North quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 89, scale 1/24,000. NIEWENDORP, C. A. 1997. Geology of the Paris Mountain Quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 99, scale 1:24,000. NIEWENDORP, C. A. & MAYBIN, A. H. III. 1994a. Geology of the Laurens South quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 77, scale 1/24,000. NIEWENDORP, C. A. & MAYBIN, A. H. III. 1994b. Geology of the Cross Hill quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 78, scale 1/24,000. OWNBY, S. E., MILLER, C. F., BERQUIST, P. J., CARRIGAN, C. W., WOODEN, J. L. & FULLAGAR, P. D. 2004. U - P b geochronology and geochemistry of a portion of the Mars Hill terrane, North Carolina-Tennessee: Constraints on origin, history, and tectonic assembly. In: TOLLO, R. P., CORRIVEAU, L., MCLELLAND, J. & BARTHOLOMEW, M. J. (eds) Proterozoic Tectonic Evolution of the Grenville

orogen in North America. Geological Society of America, Memoir 197, 609-632. PLATT, J. P. & BEHRMANN, J. H. 1986. Structures and fabrics in a crustal-scale shear zone, Betic Cordillera, SE Spain. Journal of Structural Geology, 8, 15-33. RANKIN, D. W., ESPENSHADE,G. J. & NEUMAN, R. B. 1972. Geologic map of the west half of the WinstonSalem quadrangle, North Carolina, Virginia, and Tennessee. US Geological Survey, Map 1-709-A, scale 1:250,000. ROSENBERC, C. L. & HANDY, M. R. 2005. Experimental deformation of partially melted granite revisited: implications for the continental crust. Journal of Metamorphic Geology, 23, 19-28. SAMSON, S. D. & SECOR, D. 2000. New U - P b geochronological evidence for a Silurian magmatic event in central South Carolina. Geological Society of America Abstracts with Programs, 32(2), A71. STELTENPOHL, M. G. (ed.) 2005. Southernmost Appalachian terranes, Allabama and Georgia. Southeastern Section Geological Society of America Field Trip guide. Alabama Geological Society. STELTENPOnL, M. G. & MOORE, W. B. 1988. Metamorphism in the Aabama Piedmont. Geological Survey of Alabama, Mineral Resources Circular 138. STELTENPOHL, M. G., KISH, S. A. & NEILSON, M. J. 1990. Geology of the southern Inner Piedmont, Alabama and southwest Georgia. Alabama Geological Survey & Alabama Geological Society, Guidebook for Field Trip VII. STIEVE, A. L. 1989. The structural evolution and metamorphism of the southern portion of the Tallulah Falls Dome, northeast Georgia. PhD dissertation, University of South Carolina. STUDENT, J. J. & SINnA, A. K. 1992. Carboniferous U - P b ages of zircons from the Box Ankle and Ocmulgee faults, central Georgia: Implications for accretionary models. Geological Society of America Abstracts with Programs, 24(2), 69. VAUCHEZ, A., BABAIE, H. A. & BABAE1, A. 1993. Orogen-parallel tangential motion in the Late Devonian-Early Carboniferous southern Appalachians internides. Canadian Journal of Earth Sciences, 30, 1297-1305. WEGMANN, C. E. 1935. Zur Deutung der Migmatite. Geologicschen Rundschau, 26, 20-350. WEST, T. E., JR. 1996. Geology of the Shoals Junction quadrangle, South Carolina. South Carolina Geological Survey, Open File Map 88, scale 1/24,000. WEST, T. E., JR. 1997. Structural studies along the Carolina-Inner Piedmont terrane boundary in South Carolina and Georgia: implications for the tectonics of the southern Appalachians. PhD dissertation, University of South Carolina. WEST, T. 1998. Structural analysis of the CarolinaInner Piedmont terrane boundary: Implications for the age and kinematics of the central Piedmont suture, a terrane boundary that records Palaeozoic Laurentia-Gondwana interactions. Tectonics, 17, 379-394.

CHANNEL FLOW IN THE APPALACHIAN PIEDMONT WHISONANT, J. S. 1979. Geologic map of the SE 1/4 of the Marion quadrangle, North Carolina. North Carolina Geological Survey, GM 210-SE, scale 1/24,000. WILLIAMS, S. T. 2000. Structure, stratigraphy, and migmatization in the southwestern South Mountains, North Carolina. MS thesis, University of Tennessee.

541

WILLIS, J. D. 1984. Geology of the Cross Anchor area-the boundary between the Carolina terrane and Inner Piedmont in northwestern South Carolina. MS thesis, University of South Carolina. YANAGIHARA, G. M. 1994. Structure, stratigraphy, and metamorphism of a part of the Columbus Promontory, western Inner Piedmont, North Carolina. MS thesis, University of Tennessee.

An argument for channel flow in the southern Canadian Cordillera and comparison with Himalayan tectonics RICHARD

L. B R O W N t

& H. D A N I E L G I B S O N 2

1Department of Earth Sciences, Carleton University, Ottawa, Ontario, Canada, K1S 5B6 (e-mail: richard.brown @ gmail, corn) 2Department of Earth Sciences, Simon Fraser University, Burnaby, British Columbia, Canada, V5A 1S6

Abstract: Crustal thickening in excess of 55 km, and high heat flow, suggest that a high-standing plateau region in the Cordilleran hinterland was present in the Late Cretaceous. A low strength middle crust developed beneath the plateau, and parts of this layer were exhumed to upper crustal levels in Late Cretaceous to Eocene time. During Late Cretaceous time, structures in the hinterland were reactivated. Strata, buried to mid-crustal depths since the Jurassic, began to flow upward to higher levels; earlier structures were refolded and tightened, and a new transposition fabric developed. Some 10-20 km of the middle crust was involved in high temperature ductile flow. The lower boundary of the ductile zone lies with thrust sense on top of Precambrian rocks of Canadian Shield affinity, and splays upwards to the NE where it closely coincides with highly strained rocks in the hanging wall of the Purcell Thrust Fault. The upper boundary is marked by a normal-sense high strain zone, above which only minor Cretaceous deformation occurred. The boundaries were reactivated at upper crustal levels after cessation of flow in the mid-crustal channel. This reactivation resulted in formation of ductile to brittle extension faults such as the Okanagan Fault System. During final stages of flow, the Precambrian basement gneisses at the base of the channel became domed and exhumed to upper crustal levels. Comparisons with Himalayan tectonics are clearly drawn, but there are significant contrasts such as the long residence time of the proposed Cordilleran channel, and the nature of the channel boundaries.

Deformation in the hinterland of the Rocky Mountain Thrust and Fold Belt of the southern Canadian Cordillera coincided with Early Jurassic terrane obduction onto the North American plate (Monger et al. 1982; Brown et al. 1986; Fig. la & b). F r o m Late Jurassic to Eocene, deformation migrated northeastward into the foreland giving rise to the Rocky Mountain Thrust and Fold Belt (Price & Mountjoy 1970). Parts of the metamorphic hinterland (southern Omineca Belt, Fig. la) were deeply buried and then e x h u m e d to high structural levels before the beginning of the Cretaceous (Brown & Tippet 1978; Parrish 1995; Colpron et al. 1998), but other areas within the Omineca Belt were not e x h u m e d until the Late Cretaceous and early Tertiary (e.g. Sevigny et al. 1990). Recent structural analysis and geochronology by Gibson et al. (2004, 2005; see also Crowley et al. 2000; Gibson, 2003) across the Big Bend area of the southern part of the hinterland, point to the existence of a middlecrustal zone some 10 to 20 k m thick, which was at upper amphibolite facies from as early as Middle Jurassic time until exhumation in the Late Cretaceous. The upper and lower boundaries of this ductile zone exhibit structural and thermal gradients that support a channel flow model of midcrustal deformation. In this paper we outline the

evidence for channel flow and discuss possible relationships between exhumation of the channel, formation of Tertiary gneiss domes, and developm e n t of low-angle brittle-ductile detachment faults. Finally, we discuss similarities and contrasts with Himalayan tectonics.

Geological setting The southern Omineca Belt of the Canadian Cordillera exposes a highly deformed region of metasedimentary, plutonic and metavolcanic rocks of mainly North American affinity, which form the southwestern hinterland of the Rocky Mountain Thrust and Fold Belt. A veneer of accreted terranes that was thrust northeastward onto the North American continental margin in the Jurassic is preserved in structural lows within this hinterland. T h e wide southern belt contrasts with the narrower belt north of 53~ where metamorphic grade is generally lower. This difference is a reflection of the considerable Eocene extension and tectonic denudation of the southern region and its diminished impact to the north. Obduction of the accreted terranes in the Jurassic was accompanied by crustal thickening and deep burial of North American rocks in a southwesterly-vergent

Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 543-559.

From: LAW,R. D., SEARLE,M. P. & GODIN,L. (eds)

0305-8719/06/$15.00

9 The Geological Society of London 2006.

544

R . L . BROWN & H. D. GIBSON

CHANNEL FLOW IN THE CANADIAN CORDILLERA pro-wedge setting. The pro-wedge (Willett et al. 1993) grew southwestward (deformation in North American upper plate, with vergence toward the outboard NE-dipping subducting plate) accompanied by retro-wedge growth (deformation of upper plate with vergence towards the cratonic margin of the North American plate) NE across a regional zone of structural divergence (Brown et al. 1993; Gibson 2003; Fig. 6a). By Late Jurassic time, some of these deeply buried rocks had been exhumed and carried northeastward as the retro-wedge expanded. The Jurassic, SW-vergent structures are preserved at high structural levels along the western margins of the Omineca Belt where their northeastern boundary delimits a local zone of structural divergence known as the Selkirk fan (Figs lb, 2 & 3; Wheeler 1963, 1965; Price & Mountjoy 1970; Brown & Tippett 1978; Brown et al. 1993; Colpron et al. 1998; Gibson et al. 2005). Structural culminations within the southern Omineca Belt expose regions of upper-amphibolitefacies metamorphic and plutonic rocks, bounded by generally outward-dipping low-angle normal faults that contain low grade rocks in their hanging walls (Figs lb & 2). These culminations have the characteristics of metamorphic core complexes and have been discussed in some detail in the literature (Coney 1980; Armstrong 1982; Parrish 1995; Vanderhaeghe & Teyssier 1997). Monashee and Valhalla are two complexes that expose important compressional shear zones at deep structural levels (Read & Brown 1981; Carr et al. 1987; Parrish 1995; Brown 2004; Carr & Simony 2006). These ductile zones carried hot and mobile middle-crustal rocks northeastward relative to the underlying basement rocks of the complexes, and are interpreted to extend northeastward into the discrete basal detachment of the Rocky Mountain Thrust and Fold Belt (Brown et al. 1986; Cook et aL 1992; Parrish 1995). Precambrian basement rocks of one of these domal complexes, the Monashee Complex (Fig. lb), were not deeply buried by the advancing orogen until the Latest Cretaceous (Parrish 1995; Crowley et al. 1999, 2001; Gibson et al. 1999), as evidenced by the lack of metamorphism and absence of Cordilleran plutonism until this time. In contrast, the structurally overlying rocks contain evidence of a protracted and diachronous involvement in the deeper levels of the orogen that extend from the Middle Jurassic to

545

Tertiary time. The allochthonous rocks comprise highly deformed equivalents of the late Precambrian sedimentary and volcanic pile that filled the rift basins as they developed along the western North American margin during continental break-up (Windermere Supergroup and overlying Cambrian to Ordovician strata) (Fig. lb). These strata were deposited outboard of the North American continental rocks of the Monashee Complex and were displaced northeastward over the complex during Cretaceous contraction and thickening of the orogen (Gabrielse & Yorath 1991a; Brown et al. 1993). The allochthonous rocks have been collectively referred to as the 'Selkirk allochthon' in the literature, and the boundary between these allochthonous rocks and the relatively autochthonous basement is known as the Monashee D~collement (see Parrish (1995) for review; see also Williams & Jiang (2005) for a contrary view). The metamorphosed and locally migmatized rocks that form the lower part of the Selkirk allochthon were highly ductile and penetratively deformed as they flowed northeastward. In the Late Cretaceous the Monashee Complex was deeply buried by the overriding rocks of the Selkirk allochthon. In the following sections we describe relationships within this mid-crustal ductile zone and propose that deformation within the zone can, in part, be interpreted in terms of channel flow.

The Cretaceous mid-crustal ductile zone Regional studies have established that the southern Omineca Belt exposes three distinct structural levels (Figs 1 & 2; Brown & Carr 1990; Cart 1991). The lowest level exposures are of middle to lower crustal rocks coring domal metamorphic complexes that were deeply buried in the Palaeocene and rapidly exhumed in the Eocene (Parrish 1995). Structurally above and flanking these outward-dipping domal exposures are the midcrustal rocks within the Selkirk allochthon that were mobile and at high metamorphic grade through most of Cretaceous time (Gibson 2003; Gibson et aL 2005). The highest levels exposed within the Selkirk allochthon contain strata that were variably deformed and metamorphosed in the Middle Jurassic and exhumed to upper crustal

Fig. 1. (a) Morphologic belts of the Canadian Cordillera. (b) Tectonic assemblage map of the southern Omineca Belt (modified after Wheeler & McFeely 1991) showing lithological map units of the autochthonous Monashee Complex (North American basement) and overlying Selkirk allochthon. 'Monashee Complex' exposes the deepest structural level. 'Selkirk allochthon' includes the metamorphic rocks of the middle-crustal layer in dark grey (interpreted as channel flow in the Selkirk and Monashee mountains), and the light grey panel is the upper crustal level overlying the channel. FCD, Frenchman Cap Dome; OFS, Okanagan Fault System; TOD, Thor Odin Dome.

546

R.L. BROWN & H. D. GIBSON

Fig. 2. Generalized geological map of the northern Selkirk Mountains illustrating simplified lithostratigraphy, regional metamorphic isograds, major structures, and structural levels discussed in the text. The middle crustal zone is bounded by the Selkirk Detachment Fault (SDF) at the top and the Purcell Thrust Fault (PTF) at the base. Box inset straddling SDF locates Figure 5. Section lines are for Figure 3 (e.g. A-A'). CRF, Columbia River Fault; MC, Monashee Complex. SRMT, Southern Rocky Mountain Trench. Mineral abbreviations for isograds are after Kretz (1983). See Figure 1 for location of Figure 2. levels by the Late Jurassic. The present stacking of the three structural levels is the combined result of Jurassic-Cretaceous NE-over-SW imbrication and ductile flow during crustal thickening (Brown et al. 1986), followed by Late Cretaceous midcrustal extrusion (Johnston et al. 2000) and Tertiary normal-sense low-angle extension faulting (Parrish et al. 1988). The Big Bend area

The exposed area of the Cretaceous mid-crustal ductile zone in the Big Bend area is shown in Figure 4. Upper-amphibolite-facies migmatitic schists and gneisses predominate, and are derived from transposed equivalents of the Proterozoic Windermere Supergroup together with local slices of infolded lower Palaeozoic Hamill Group (Wheeler 1965; Brown & Tippett 1978; Wheeler & McFeely 1991). Characteristically these highly

strained rocks exhibit boudinage of competent units, isolated fold hinges and complex fabrics indicative of at least two generations of transposition and later superimposed folding. These late folds (F3) are overturned to the NE on the northeastern side of the Selkirk Fan (Fig. 3; Brown & Tippett 1978; Brown et al. 1993). The folds were generated in the Late Cretaceous during the final stages of transposition while the rocks were still at upper amphibolite facies (Fig. 3; Gibson 2003; Gibson et al. 2005). Within the Big Bend area the exhumed midcrustal zone lies in the hanging wall of the Purcell Thrust Fault (PTF). These relationships are complicated by superimposed Tertiary normal faults that extend along the physiographic trend of the Southern Rocky Mountain Trench (SRMT; Figs 1 & 2). To the NE lie the Main Ranges of the Rocky Mountains with characteristic structures of the upper crustal Thrust and Fold Belt (Price & Mountjoy 1970;

CHANNEL FLOW IN THE CANADIAN CORDILLERA

547

Fig. 3. Composite structural cross-section that transects the studied area, illustrating the geometry of the Selkirk fan structure modified after Brown & Tippet (1978), Simony et al. (1980), Perkins (1983), Colpron et al. (1995) and Gibson et al. (2005). PTF, Purcell Thrust Fault; SDF, Selkirk Detachment Fault; CRF, Columbia River Fault. Rocks lying within the proposed channel are in the footwall of the SDF and the hanging wall of the PTF. Section lines are located on Figure 2. Geochronology results of Gibson et al. (2005) have been projected along-strike into the line of section. Note: subdivisions of Windermere Supergroup are included to highlight structural geometry.

Wheeler & McFeely 1991). At this latitude, northeasterly directed thrusting on the Purcell Thrust Fault and associated erosion has most likely been in part responsible for exhumation of the metamorphic rocks of the Omineca Belt that lie in its hanging wall (Simony e t al. 1980; Wheeler & McFeely 1991). The thrust is 'out of sequence' in that the orogenic front had migrated northeastward well into the Rocky Mountain foreland before Late Cretaceous time (Price & Mountjoy 1970); brittle faults associated with the Purcell Thrust Fault cut across these earlier thrusts. It is recognized that the trace of the Purcell Thrust Fault has a steeper dip than the ductile rocks in its hanging wall, and that locally exposed metamorphic rocks in the footwall probably represent a northeastward continuation of the ductile zone (see Fig. 6c). As such, the most recent motion on the Purcell Thrust Fault, together with erosion, were responsible for exhumation of the proposed channel. The lower boundary of the

channel, while it was active, is inferred to have been the northeastward continuation of the Monashee D6collement (or a related shear zone). The distinction between the current boundaries, that were presumably active during exhumation of the ductile zone, and the boundaries that prevailed during active flow is important in understanding the tectonic evolution of the channel. This point is discussed more completely in the final sections of the paper. The upper boundary of the ductile zone, herein named the Selkirk Detachment Fault (SDF; Figs 2, 5a & 6c), is well exposed within the Big Bend area where it is marked by ductile shearing, rotation of pre-existing structures, ubiquitous lenses of pegmatite and leucogranite, stratigraphic omissions and cut-offs, together with local brittle faulting. The area was first mapped by Wheeler (1965) and later by Brown and graduate students in the early 1970s (see Brown & Tippett 1978). A study of metamorphic assemblages by Leatherbarrow (1981)

548

R . L . BROWN & H. D. GIBSON

CHANNEL FLOW IN THE CANADIAN CORDILLERA across the zone suggested a pressure change from 5 kbar on the western side to 7 kbar on the eastern side, but the times at which the rocks attained these pressures were not known. It now appears that the 5 kbar assemblages most likely originated in Jurassic time, whereas the 7 kbar assemblages are Late Cretaceous. These diachronous relationships were not understood until geochronologic work by Gibson (2003), and Gibson et al. (2005). U - T h - P b monazite and zircon dating of metamorphic rocks and related intrusions across the structural thickness of the ductile zone has established that rocks to the east of the shear zone remained hot and deeply buried until the midCretaceous (c. 100 Ma), but strata to the west, in the hanging wall, were at upper crustal levels in the Middle to Late Jurassic (172-167 Ma; see also Brown et al. 1992b; Colpron et al. 1996). Within the Big Bend area Neoproterozoic rocks of the Windermere Supergroup generally lie to the NE of the shear zone. Lower Palaeozoic rocks of the Lardeau Group characterize the strata of the southwestern panel. Highly strained quartzite and carbonate lenses within the shear zone are most likely remnants of Eocambrian strata of the Hamel Group and Ordovician units of the Badshot Formation, respectively. Upper units of the Windermere Supergroup are cut out in the footwall of the shear zone. Based on these observations the shear zone is interpreted to be a major extensional ductile fault zone. The fault zone at the surface dips steeply to the SW (Fig. 3); it is inferred to flatten at depth and be truncated by the northeasterly dipping Columbia River normal fault (Figs 1 & 6c). As shown in the cross-section of Figure 3, the folded transposition fabric of the channel rocks intersects the trace of the fault at a high angle, but is parallel with the F3 axial planar fabric. We interpret this relationship as postchannel flow normal faulting along the SDF that may have been concurrent with generation of the post-transposition F3 structures. Future workers may discover that beyond our study area the tectonic boundary lies at higher or lower stratigraphic levels than presently defined. The hanging wall rocks, which were exhumed in Jurassic time, are locally at high metamorphic grade, and in the absence of detailed geochronologic constraints,

549

juxtaposition of highly strained amphibolite-facies rocks exhumed in the Jurassic with similar grade rocks exhumed in the Late Cretaceous may not be readily identified in the field. The relationship of the Adamant pluton (Figs 1 & 2) to the channel boundaries requires additional study. Metamorphic grade decreases south of the Adamant pluton (Fig. 2) and there is little evidence of high-temperature ductile flow or melt generation in these lower grade rocks. The Adamant pluton was emplaced in the Jurassic, and its metamorphosed boundary is highly deformed (Shaw 1980; Gibson 2003). At the southwestern margin there is evidence of mainly Jurassic deformation and metamorphism with a weak Cretaceous overprint, but data are not yet available for the timing of deformation and metamorphism at the pluton's eastern margin. The transposition fabric of the ductile zone (channel) is concordant with the northern boundary of the pluton and wraps around its eastern end. These data suggest that the pluton lies in the hanging wall of the ductile zone. Apparently the folded carapace of the ductile zone plunges southward beneath the Adamant pluton and presumably continues southward in the subsurface. A marble unit (middle marble of Fig. 2) lies within the ductile zone of the study area. At the eastern end of the Adamant pluton it is preserved as a highly strained unit in sheared contact with the pluton boundary. The marble apparently continues southeastward beyond the limits of the current study where it is less deformed and is associated with lower grade stratigraphy.

Timing o f deformation and exhumation within the Big B e n d area Exhumation of the mid-crustal zone to upper crustal levels by Late Cretaceous to early Tertiary time is well constrained by monazite and zircon geochronology (Gibson 2003; Gibson et al. 2004, 2005). Determination of when the rocks were flowing at middle-crustal levels is more difficult to establish. As previously mentioned, Gibson concluded that the NE-verging F3 folds formed in the Late Cretaceous and that metamorphic minerals orientated in the pre-F3 transposition fabric ($1-$2) is also Late

Fig. 4. Location and general summary of geochronological results from the proposed mid-crustal channel within the southern Omineca belt and from overlying cover rocks. The channel rocks are shown in dark grey, the cover rocks are in light grey, and the underlying Monashee Complex is white. Boxes that provide only a range of ages constrain the time of deformation (def.) and metamorphism (met.) for their respective locations. Additional text is included in the boxes for locations where the ages also constrain the cooling history or a metamorphic overprint, or only relate to the time of deformation (e.g. 'foliated pluton'). Details of these locations may be found by referring to the numbered reference. MD, Monashee D6collement; PTF, Purcell Thrust Fault; OFS, Okanagan Fault System.

550

R . L . B R O W N & H. D. G I B S O N

CHANNEL FLOW IN THE CANADIAN CORDILLERA

551

Fig. 6. Generalized cross-sections demonstrating (a) Late Jurassic configuration of the evolving orogen above an easterly dipping subduction zone. The pro-wedge illustrates obducted and accreted terrane at the highest structural level with underlying rocks of North American affinity. The dark grey layer schematically illustrates development of the mid-crustal channel above attenuated basement rocks of the North American plate. (b) Late Cretaceous-Palaeocene configuration of middle crustal channel prior to extensional faulting. (c) Present-day geometry across Frenchman Cap Dome; geometry of the Moho is constrained by Lithoprobe seismic reflection profile of Cook et al. (1992). The middle-crustal layer is interpreted to have been a zone of channel flow in the Late Cretaceous. Formation of Frenchman Cap Dome and displacement on Columbia River Fault (CRF) post-date the proposed channel flow. MC, Monashee Complex; MD, Monashee Drcollement; OFS, Okanagan Fault System; PT, Purcell Thrust Fault; SDF, Selkirk Detachment Fault; SFA, Selkirk Fan Axis. See text for additional explanation and references. These diagrams are modified from Gibson 2003.

Cretaceous in age (Figs 3 & 4). The F3 folds are tight to open flexural flow folds that have shortened the original width o f the ductile zone, and have locally folded metamorphic isograds (Simony e t al. 1980;

Leatherbarrow 1981). Monazite and zircon analyses indicate that at least some o f the rocks in the zone were hot and ductile as early as c. 140 M a (Gibson 2003; Gibson e t aL 2005) and there is local evidence

Fig. 5. (a) View of the upper boundary of the proposed mid-crustal channel as exposed within the northern Selkirk Mountains. See Figure 2 for location. The Selkirk Detachment Fault (SDF) is drawn below the cliffs that expose marbles of the Ordovian Badshot Formation. These marbles are truncated by the SDF. Neoproterozoic units of the Windermere Supergroup form the cliffs below the SDF. The Badshot Formation and overlying Lardeau Group seen above the SDF were exhumed to high structural levels in the Jurassic. Underlying Windermere Supergroup rocks were at mid-crustal levels in the Late Cretaceous. These footwall rocks are highly transposed, migmatized and intruded by leucogranite and pegmatite (outlined by black lines). See text for further explanation. (b) Outcrop view of typical highly deformed unit of the Windermere Supergroup in the proposed mid-crustal channel. First- and second-generation folds are rootless and lithologic units are transposed. The exposure is within the Windy Range approximately 10 km NE of the SDF shown above.

552

R.L. BROWN & H. D. GIBSON

of high temperature flow as early as c. 160 Ma (Crowley et al. 2000). Together these results indicate the presence of a mid-crustal ductile zone that was active, either continuously or discontinuously, from Middle Jurassic time until Late Cretaceous, a period of approximately 100 million years. Parrish (1995) demonstrated that the ages of peak metamorphism within the middle-crustal zone are younger in the structurally deeper levels compared to the higher levels. Brown (2004) pointed out that during the prolonged period of middle-crustal deformation, new strata would be incorporated into the deforming zone by progressive underplating as the orogenic front advanced. The development of transposition fabrics was apparently diachronous; the earlier formed tectonites must have been redeformed and transported northeastward by the younger deformation events within the evolving orogen. Generation of the F3 folds marks the final stages of ductile flow as the middle crust was exhumed.

The ductile z o n e w e s t o f the M o n a s h e e Complex

The distribution of rocks that were at mid-crustal depth in Late Cretaceous time is shown in Figure 4. Structures and metamorphic relationships within this region have been described in detail by several authors (Simony et al. 1980; Carr 1991; Johnson & Brown 1996; Johnston et al. 2000). Mineral assemblages characteristically indicate upper-amphibolite-facies conditions; the rocks are commonly migmatitic and contain plutonic sheets ranging in age from Palaeozoic through Mesozoic to as young as Eocene. The upper boundary of the zone is marked by the Okanagan Fault System (OFS, Fig. 1). The fault zone has been interpreted to be a low-angle extensional structure that was active in the Eocene during a period of crustal extension and core complex formation (Templeman-Kluit & Parkinson 1986; Parrish et al. 1988; Cook et al. 1992). There is evidence in the footwall rocks of rapid cooling and exhumation of high-temperature migmatitic rocks as early as the Late Cretaceous (Scammell 1993; Parrish 1995; Johnson & Brown 1996). Mylonitic fabrics in these footwall rocks exhibit a pronounced east-west trending mineral (commonly sillimanite) and stretching lineation; kinematic indicators are consistent with upper-platedown-to-the-west sense of motion. Adjacent to the hanging wall the rocks are commonly retrogressed to chlorite grade and overprinted by brittle fabrics. Rocks in the upper plate of the OFS are generally of low metamorphic grade, but also contain high-grade assemblages. These metamorphic rocks were exhumed by the end of the Jurassic and were at upper crustal levels by mid-Cretaceous time

(Schiarizza & Preto 1987; Johnson & Brown 1996). The sillimanite-K feldspar-bearing migmatitic rocks in the footwall of the OFS extend eastward and structurally down-section through a thickness of approximately 10 km. At the lowest structural level the migmatites are in sheared contact with the Precambrian basement rocks of the Monashee Complex. This lower boundary has previously been discussed extensively in the literature and is known as the Monashee Drcollement (Read & Brown 1981; Brown et al. 1986; Journeay 1987; Cook et al. 1992; Parrish 1995). For a distance of several kilometres from the boundary, kinematic indicators are well developed in the highly transposed migmatites, which clearly show top-to-the-NE sense of shear (Joumeay 1987; Brown et al. 1992a; Johnson 1994; Johnston et al. 2000). Locally superimposed on these fabrics are top-to-the-west extensional shear bands and associated normal faults (Johnston et al. 2000). Within the underlying rocks of the Monashee Complex metamorphic grade is generally lower but maintains amphibolite facies. In the northem part of the complex (Frenchman Cap Dome; Fig. lb) the effects of the middle-crustal deformation appear to decrease structurally downwards away from the d~collement (Gibson et al. 1999; Crowley et aL 2001). However, to the south in the Thor Odin Dome (Fig. lb) there is a very intense Eocene migmatitic and deformation event that appears to have masked these earlier relationships (Parrish et al. 1988; Carr 1991; Vanderhaeghe & Teyssier 1997; Johnston et al. 2000).

Continuity of the mid-crustal ductile zone The regional extent of the Cretaceous mid-crustal ductile zone in the southern Omineca Belt is illustrated in Figures lb, 2 and 4. Within the Big Bend area the zone is bounded above by the Eocene age Columbia River extension fault (Fig. 2). This fault leaves the river valley at its northern extremity and loses displacement as it curves into the Selkirk Mountains in the vicinity of Birch Creek. Footwall rocks of the mid-crustal zone extend northwestward beyond Birch Creek and cross the Columbia River valley into the ranges of the Monashee Mountains. These rocks of the Windermere Supergroup then swing westward around the northern end of the Monashee Complex and are continuous with the midcrustal zone that has been mapped on the western margin of the complex (Figs lb, 2 & 4). A detailed study of these strata at the northern end of the Monashee Complex was carried out by Scammell (1993), and at higher structural levels by Sevigny et al. (1989, 1990). We have also worked in the region and have found that the style of deformation

CHANNEL FLOW IN THE CANADIAN CORDILLERA and intensity of transposition is similar to the midcrustal zone within the Big Bend area. Results of field and geochronological studies in the region are reviewed in Parrish (1995). He concluded that the crustal zone has a diachronous thermal history with higher levels indicating older ages of peak metamorphism and generation of leucogranitic melt. The deepest structural level of the zone gives ages that range from Late Cretaceous to Palaeocene, while at higher levels thermal peak conditions were reached in Early to mid-Cretaceous time. This diachroneity may be explained in part by progressive underplating and exhumation, as suggested by Brown (2004). The middle-crustal zone appears to have been ductile and highly mobile throughout Cretaceous time, and the zone was a continuous crustal feature from as far west as the bounding Okanagan Fault System, to as far east as the Purcell Thrust Fault, which is interpreted to be a splay of the Monashee D6collement exposed in the southern Rocky Mountain Trench. Since the Big Bend panel is continuous with the exposed ductile zone to the west of the Monashee Complex, it is reasonable to assume that the boundaries of the zone may also be correlated. Palinspastic reconstruction by Johnson & Brown (1996) illustrates the continuity of the mid-crustal zone after restoration of the Columbia River Fault. It appears to be a requirement of this reconstruction that the Selkirk Detachment Fault correlates westward with the Okanagan Fault System.

Interpretation of the Cretaceous mid-crustal zone Johnston et al. (2000) concluded that the panel of midcrustal rocks exposed west of the Monashee Complex was extruded northeastward in Late Cretaceous time. The argument for extrusion is based largely on a vergence reversal of F3 folds from southwesterly in the upper part of the panel to northeasterly in the lower part. Before generation of the F3 folds, the panel developed F1/F2 penetrative structures that indicate northeasterly vergence across the width of the panel. The model presented in Johnston et al. (2000, Fig. 10) proposes that in F1/F2 times northeastward ductile thrusting carried hinterland rocks (Selkirk allochthon) towards the Rocky Mountain foreland, and ductile flow in the hinterland was balanced by contraction on thrust faults within the foreland. We interpret the upper boundary of ductile deformation (the upper crust-middle crust boundary) through Cretaceous time to have been located in the vicinity of the future Okanagan Fault System. The lower boundary lies within the upper part of the basement rocks of the Monashee Complex. During F1/F2 times, the upper levels of the orogen together with

553

the ductile rocks of the middle crust were carried northeastward across the more competent basement rocks of the Monashee Complex. Whether or not the upper and lower boundaries were defined in Cretaceous time by discrete shear zones or more diffuse crustal-scale gradients remains unclear. In this paper the term Monashee D6collement is retained (see Brown et al. 1992a, and references therein) to refer to this lower boundary of the mid-crustal ductile zone. Similarly the upper boundary is called the Okanagan Fault System, even though the boundary is not known to have been a discrete extensional fault zone until Tertiary time.

D o the data s u p p o r t a c h a n n e l flow model?

In the Late Cretaceous, when F3 folds were being generated within the ductile zone to the west of the Monashee complex, channel flow appears to be qualitatively supported by the data for the panel west of the Monashee Complex (Johnston et al. 2000). In the Big Bend area, late to post-transposition folds (F3) are overturned towards the NE across the width of the belt and vergence reversals of minor folds reflect their superposition on the limbs of the major folds. In this region it appears that most of the high-temperature ductile flow and associated transposition occurred during F1/F 2 deformation, and vergence of these tight to isoclinal folds is also predominantly northeasterly across the width of the belt. Associated discrete mylonitic shear zones also exhibit kinematic indicators that suggest a NE-over-SW sense of shear. The F3 folding resulted in an anticlinorium within the central part of the ductile zone (Brown & Tippett 1978). The structurally deepest rocks expose the highest metamorphic grade (Fig. 2). These structural observations, in the Big Bend area, are compatible with northeasterly motion and folding of a crustal-scale ductile thrust sheet and do not in themselves require channel flow. During exhumation of the mid-crustal ductile zone, normal-sense motion on SDF and thrust-sense motion on the Purcell Thrust Fault appear to have overlapped in time. To this extent the mid-crustal layer was being structurally unroofed while it was still moving northeastward as an active thrust sheet. The timing constraints presented in this paper and elsewhere indicate that concurrent motion on the SDF and Purcell Thrust Fault occurred in the Late Cretaceous when, at mid-crustal levels in the thrust sheet, migmatites, local melts and ductile transposition fabrics were being generated. These data lead the authors to a working hypothesis of Late Cretaceous mid-crustal channel flow.

554

R.L. BROWN & H. D. GIBSON

Tectonic model Figure 6c is a simplified cross-section illustrating the present geometry after doming and erosion. In this figure, the mid-crustal ductile zone is shown bounded by the Monashee D6collement and Purcell Thrust Fault at its base and by the Okanagan Fault System and Selkirk Detachment Fault at the top. Flow within the ductile zone is envisaged to have occurred in two distinct periods. The early period (F1/F2), which extended from Jurassic to Late Cretaceous, is the time of crustal thickening involving westward underplating of the orogen by the sediments of the foreland and underlying cratonic rocks of North America (Fig. 6a & b). The earliest deformed rocks of the hinterland were exhumed from mid-crustal depths as early as Middle Jurassic time and were carried northeastward above the underlying rocks of the mid-crustal ductile zone (Fig. 6a; Brown et al. 1986; Colpron et al. 1996). Crustal thickening on the order of 55 km was achieved by Late Cretaceous time (Johnson & Brown 1996; Foster et al. 2004). We assume that this thickness, together with high heat flow in the hinterland of the orogen, would have created an extensive plateau region (Fig. 6b). Extensional deformation in the hinterland occurred in Eocene time (Parrish et al. 1988), and the present crustal thickness had been reduced to less than 35 km (Cook et al. 1992). Whether or not there was significant crustal thinning before the Tertiary is not clearly resolved. However, within the ductile zone to the north of the Monashee Complex, pressuretemperature-time results of Scammell (1993) suggest that part of the zone was rapidly exhumed as early as 100 Ma. He envisaged a mechanism of ductile thrusting and thrust-parallel extension within the zone to account for decreases in pressure during progressive ductile flow. It is in this time period, continuing into the Tertiary, that channel flow is thought to have been ongoing in the mid-crustal zone. The possibility that erosion at the front of the thrust sheet may have facilitated channel flow at depth should be considered. This time was a period of rapid sediment accumulation in the foredeep and some of these deposits are presumed to have been derived from the Omineca Belt. However, testing of such a hypothesis is beyond the scope of this paper.

Gneiss dome formation The Monashee Complex exposes Precambrian basement in the cores of the gneiss domes of Frenchman Cap and Thor Odin (Fig. lb). Geochronology and plutonic history have demonstrated that these core

rocks are part of the North American Precambrian crust that underlies the Rocky Mountain Thrust and Fold Belt (Crowley 1999). Field structural studies over the past 30 years have elucidated the geometry of these basement rocks and an unconformably overlying, but highly transposed, Precambrian to Palaeozoic cover sequence (e.g. Journeay 1987; Scammell & Brown 1990; Parrish 1995; Gibson et al. 1999; Crowley et al. 2001). Above these crustal rocks is the Monashee D6collement, which carries allochthonous rocks of the Cordilleran orogen in its hanging wall. The main body of the proposed mid-crustal channel that is discussed in this paper lies above the basement rocks that core the domes (Fig. 6c). However, during emplacement of the channel rocks onto the basement rocks of the Monashee Complex in the latest Cretaceous to Eocene, these basement rocks were metamorphosed, locally migmatized and highly deformed. Restoration of sections to pre-doming configuration, together with timing constraints from geochronology, indicates that flow within the mid-crustal channel was completed before exhumation of the domes (Johnson & Brown 1996). The migmafites within the basement rocks of the domes are of two distinct origins. The first episode of migmatization occurred in the Palaeoproterozoic and is part of the pre-Cordillera basement history (Crowley 1999). The second migmatization occurred in Palaeocene to Eocene time; thermal and structural data from the Frenchman Cap Dome point to a decrease downwards in heat and deformation during this time interval (Parrish 1995; Gibson et al. 1999; Crowley 1999; Crowley et al. 2001). These observations indicate that the hottest part of the crust was above the basement rather than within it. Structuralmetamorphic data combined with detailed geochronology indicate that these hot migmatites were flowing northeastwards across the cooler and less mobile underlying basement rocks. In parts of the core of Thor Odin Dome there is evidence of partial melting in the Eocene, and it has been suggested that this melt weakening led to important ductile flow during exhumation associated with crustal thinning (Vanderhaeghe & Teyssier 1997). Although closely associated in time, it is important to distinguish between the Late Cretaceous to Palaeocene proposed channel flow, that was active mainly in the ductile zone above the basement gneisses, and heating of the basement gneisses in the Eocene after major flow within the channel was completed. Flow in the channel culminated in the earliest Palaeocene c. 65 Ma; Tertiary migmatites in Thor-Odin were primarily generated at c. 55 Ma. It is likely that by this time, heating of the basement rocks was related to deep crustal or subcrustal processes associated with the onset of crustal extension (Ranalli et al. 1989).

CHANNEL FLOW IN THE CANADIAN CORDILLERA A scenario that appears to best fit the data available in the literature, as well as our own observations, includes the following: Late Cretaceous to Palaeocene initiation of doming (Brown & Journeay 1987; Scammell 1993; Johnson & Brown 1996) possibly related to thinning of the overlying channel; Eocene boudinage of the middle crust, possibly including the lower crust, (Price et al. 1981) with associated partial melting in the Eocene (Vanderhaeghe & Teyssier 1997); Eocene formation of low-angle normal faults in the upper crust associated with rapid exhumation of the domes (Parrish et al. 1988; Parrish 1995).

Comparison with channel flow in the High Himalaya Similarities in the tectonic model proposed for the southern Canadian Cordillera and models for the High Himalaya are obvious, but important differences may be as significant as the similarities. At the latitude of the current study, the Purcell Thrust Fault marks the transition from exhumed mid-crustal rocks in its hanging wall to upper crustal rocks in its footwall. This boundary has similar characteristics to the Main Central Thrust zone at the base of the High Himalayan channel. The Okanagan Fault System marks the transition from exhumed middle crust in its footwall to upper crustal rocks in its hanging wall. This boundary has similar characteristics to the South Tibetan Detachment System at the top of the proposed Himalayan channel. In the Cordillera it is recognized that the proposed boundaries of the channel were modified during exhumation and are oblique to the transposition fabric in the channel rocks. This also appears to be the case in the Himalaya at least for the Main Central Thrust (Searle & Szulc 2005). Structures within both channels are characterized by transposition of original stratigraphic boundaries, rootless isoclinal folds, boudinage of competent units, and polyphase folding. Metamorphic assemblages, migmatites and leucogranitic sheets in both channels reflect temperature and pressure conditions compatible with residence at mid-crustal depths. Strata within the Cordilleran channel are primarily derived from sediments originally deposited on the western margin of the North American continent. Strata within the Himalayan channel are primarily derived from the northern margin of the Indian continent. Deformation of these sediments and crustal thickening in the hinterland of both orogens is a result of underthrusting of their cratonic margins. Gneiss domes are exposed in the hinterland of both orogens behind the erosion front of the extruded channel. The high elevation of the Tibetan Plateau, together with erosion along its southern flank are

555

modelled as being the driving forces of channel flow in the Himalayan orogen (Beaumont et al. 2001a). A similar gradient is proposed in the hinterland of the Cordillera for the time of flow within the Cordilleran channel. Erosion of the eastern flank of the hinterland during channel flow is suggested by synchronous accumulation of sediment in the foredeep (Price & Mountjoy 1970). Conversely, the mid-crustal zone in the southern Canadian Cordillera was, intermittently or possibly continuously, deforming at high temperature for approximately 100 million years. Channel flow in the Himalayas is thought to have been shortlived, occurring from as early as 30 Ma to about 17 Ma. Rapid erosion of the High Himalayan southern margin was facilitated by monsoon conditions. The eastern margin of the proposed Cordilleran plateau was an arid area in the lee of prevailing westerly Pacific winds. Channel flow occurred within the 'pro-wedge' of the Himalayan orogen, and in the 'retro-wedge' of the Cordillera. The Tibetan Plateau remains today as the largest high-standing area in the world. Modest elevations are currently present in the hinterland of the Cordillera, with areas in the Intermontane region almost at sea level. During proposed channel flow in the Himalaya, plate convergence was essentially orthogonal, but in the Cordillera Late Cretaceous and Tertiary convergence was oblique. The High Himalayan ductile zone is continuously exposed along the strike length of the orogen. In the Cordillera the complete along-strike exposure and structural characteristics of the mid-crustal zone remain to be established. It is clear from available data that the mid-crustal zone is discontinuously exposed in structural culminations. It may turn out that along parts of the strike length of the Cordillera, the mid-crustal layer behaved primarily as a ductile thrust sheet with limited or no evidence of channel flow (e.g. Carr & Simony 2006). Such variation along-strike may perhaps be attributed to oblique convergence, as recently proposed for the Kaoko Belt of Namibia (Goscombe et al. 2005).

Discussion The above comparison of the setting of channel flow in the southern Canadian Cordillera with the setting of channel flow in the High Himalaya raises several points for discussion. The apparent longevity of the mid-crustal ductile zone in the Cordillera, compared with that of the Himalaya, may well be explained by the different climatic setting of the two orogens. Beaumont et al. (2001b) demonstrated that in the absence of rapid erosion on the foreland side of an orogen, in this case the Canadian Cordillera, the ductile zone will

556

R.L. BROWN & H. D. GIBSON

remain deeply buried. The absence of a high plateau in the Intermontane region of the southern Canadian Cordillera is presumed to be a reflection of the intensity of crustal extension in the Eocene. Crust in the hinterland exceeded 55 kin in thickness in the Late Cretaceous and has since been reduced by c. 40% due to tectonic denudation and subsequent erosion. The Eocene is known to have been a time of relaxed orthogonal plate convergence and development of a dextral plate boundary setting (Gabrielse & Yorath 199 l b). In contrast, the Himalayan region is still experiencing orthogonal plate convergence. The High Himalaya lies in a pro-wedge setting (Willett et al. 1993) if it is accepted that the Indian plate is being subducted northward beneath the Tibetan Plateau (cf. Jamieson et al. 1996). In this interpretation the channel was extruded towards the subducting plate. Irrespective of how lithosphere is ultimately incorporated into the upper mantle, the observation that the Indian plate is underthrust northward relative to the overriding orogen is not in dispute. In the North American Cordillera the current polarity of subduction is clear in that oceanic rocks of the Pacific and Juan de Fuca plates are obliquely underthrusting the North American margin. This polarity is thought to have persisted from at least Jurassic time to the present (Oldow et al. 1990; Gabrielse & Yorath 1991b; Brown et al. 1993). In this respect, the Rocky Mountain Thrust and Fold Belt together with its hinterland evolved in a retrowedge setting; the front of the orogenic wedge migrated northeastward in the direction of the subducting plate (Willett et al. 1993). Despite these differences in the two orogens, both experienced underthrusting of continental crust beneath the orogenic wedge. A plateau region developed with growth of the wedge, and a mid-crustal ductile channel developed beneath the plateau regions. In both cases, flow within the channel was directed outwards from beneath the plateau region towards the evolving thrust belt that had developed above the underthrusting continental margin.

Conclusions A hot and low-strength, 10-20 km thick mid-crustal layer formed beneath a plateau region in the hinterland of the Rocky Mountain Thrust and Fold Belt. Ductile flow within this layer was active from initial crustal thickening in the Early Jurassic to its exhumation in Late Cretaceous to Palaeocene time. F~/F2 flow within the channel resulted in ductile folding and transposition of stratigraphy. Initially, the upper crust was carried passively above the ductile thrust sheet as it progressed northeastward towards the foreland. It is suggested that in the

later stages of F1/F2 deformation the passive roof to the ductile thrust sheet began to lag behind and the Okanagan Fault System became established as a normal-sense shear zone at the top of the midcrustal layer. F3 folding within the southwestem part of the ductile zone developed during channel flow. In the northeastern part of the ductile zone the F3 folds appear to post-date the proposed channel flow. The lower boundary of the channel was the Monashee Drcollement, and the upper boundary was the Okanagan Fault System that included the Selkirk Detachment Fault. The Purcell Thrust Fault is interpreted to be a splay of the Monashee Drcollement that was active during exhumation of the eastern part of the channel. Doming of the channel and underlying basement gneisses postdated the proposed channel flow, but the domes may have been localized by earlier necking of the channel. The authors gratefullyacknowledgea National Scienceand Engineering Research Council grant to R.L.B. The authors benefited from discussions with C. Beaumont, S. Carr, R. Jamieson, R. Price, P. Simony and P. Williams during the research for this publication. C. Teyssier and P. Williams are thanked for their constructivereviews.

References ARMSTRONG, R. L. 1982. Cordilleran metamorphic core complexes - from Arizona to southern Canada. Annual Review of Earth and Planetary Sciences, 10, 129-154. BEAUMONT, C., JAMIESON,R. A., NGUYEN,M. H. & LEE, B. 2001a. Himalayan tectonics explained by extrusion of a low viscosity crustal channel coupled to focussed surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON,R. A., NGUYEN,M. H. & LEE, B. 200lb. Mid-crustal channel flow in large hot orogens: results from coupled thermalmechanical models. In: COOK, F. & ERDMER, P. (eds) Slave-Northern Cordillera Lithospheric Evolution (SNORCLE) and Cordilleran Tectonics Workshop, Victoria, BC, 112-170. BROWN, R. L. 2004. Thrust belt accretion and hinterland underplating of orogenic wedges: an example from the Canadian Cordillera. In: MCCLAY,K. R. (ed.) Thrust Tectonics and Hydrocarbon Systems. American Association of Petroleum Geologists, Memoir 82, 51-64. BROWN,R. L. & CARR,S. C. 1990. Lithospheric thickening and orogenic collapse within the Canadian Cordillera. Pacific Rim '90 Congress. Australasian Institute of Mining and Metallurgy, Brisbane, Australia, 1- 10. BROWN, R. L. & JOURNEAY, J. M. 1987. Tectonic denudation of the Shuswap metamorphic terrane of southeastern British Columbia (Canada). Geology, 15, 142-146.

CHANNEL FLOW IN THE CANADIAN CORDILLERA BROWN, R. L. & TIPPETT, C. R. 1978. The Selkirk fan structure of the southeastern Canadian Cordillera. Geological Society of America Bulletin, 89, 548-558. BROWN, R. L., JOURNEAY, J. M., LANE, L. S., MURPHY, D. C. & REES, C. J. 1986. Obduction, backfolding and piggyback thrusting in the metamorphic hinterland of the southeastern Canadian Cordillera. Journal of Structural Geology, 8, 255 -268. BROWN, R. L., CARR, S. D., JOHNSON, B. J., COLEMAN, V. J., COOK, F. A. & VARSEK, J. L. 1992a. The Monashee d6collement of the southern Canadian Cordillera: a crustal-scale shear zone linking the Rocky Mountain Foreland to lower crust beneath accreted terranes. In: MCCLAY, K. R. (ed.) Thrust Tectonics. Chapman and Hall, London, 357-364. BROWN, R. L., McNICOLL, V. J., PARRISH, R. R. & SCAMMELL, R. J. 1992b. Middle Jurassic plutonism in the Kootenay Terrane, northern Selkirk Mountains, British Columbia. Geological Survey of Canada Paper, 91-2, 135-141. BROWN, R. L., BEAUMONT, C. & WILLETT, S. D. 1993. Comparison of the Selkirk fan structure with mechanical models: implications for interpretation of the southern Canadian Cordillera. Geology, 21, 1015-1018. CARR, S. D. 1991. Three crustal zones in the Thor-Odin-Pinnacles area, southern Omineca Belt, British Columbia. Canadian Journal of Earth Sciences, 28, 2003-2023. CARR, S. D. & SIMONY, P. S. 2006. Ductile thrusting versus channel flow in the southeastern Canadian Cordillera; evolution of a coherent crystalline thrust sheet. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 541-548. CARR, S. D., PARRISH, R. R. & BROWN, R. L. 1987. Eocene structural development of the Valhalla Complex, southeastern British Columbia. Tectonics, 6, 175-196. COLPRON, M., LOGAN, J. M., GIBSON, G. & WILD, C. J. 1995. Geology and Mineral Occurrences of the Goldstream River Area, Northern Selkirk Mountains (82M/9 and part of lO). British Columbia Ministry of Energy, Mines and Petroleum Resources Map 1995-2. COLPRON, M., PRICE, R. A., ARCHIBALD, D. A. & CARMICHAEL, D. M. 1996. Middle Jurassic exhumation along the western flank of the Selkirk fan structure: thermobarometric and thennochronometric constraints from the Illecillewaet synclinorium, southeastern British Columbia. Geological Society of America Bulletin, 108, 1372-1392. COLPRON, M., WARREN, M. J. & PRICE, R. A. 1998. Selldrk fan structure, southeastern Canadian Cordillera: tectonic wedging against an inherited basement ramp. Geological Society of America Bulletin, 110, 1060-1074. CONEY, P. J. 1980. Cordilleran metamorphic core complexes: An overview. In: CRITENDEN, M. L.,

557

CONEY, P. J. & DAVIS, G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America, Memoir, 153, 7-31. COOK, F. A., VARSEK, J. L., CLOWES, R. M. Er AL. 1992. Lithoprobe crustal reflection cross section of the southern Canadian Cordillera, 1. Foreland thrust and fold belt to Fraser River Fault. Tectonics, 11, 12-35. CROWLEY, J. L. 1999. U-Pb geochronologic constraints on Paleoproterozoic tectonism in the Monashee complex, Canadian Cordillera: Elucidating an overprinted geologic history. Geological Society of America Bulletin, 111,560-577. CROWLEY, J. L., GHENT, E. D., CARR, S. D., SIMONY, P. S. & HAMILTON, M. A. 2000. Multiple thermotectonic events in a continuous metamorphic sequence, Mica Creek area, southeastern Canadian Cordillera. Geological Materials Research, 2, 1-45. CROWLEY, J. L., BROWN, R. L. & PARRISH,R. R. 2001. Diachronous deformation and a strain gradient beneath the Selkirk allochthon, northern Monashee complex, southeastern Canadian Cordillera. Journal of Structural Geology, 23, 1103-1121. FOSTER, G., PARRISH, R. R., HORSTWOOD, M. S. A., CHENERY, S., PYLE, J. & GIBSON, H. D. 2004. The generation of prograde P-T-t points and paths; a textural, compositional, and chronological study of metamorphic monazite. Earth and Planetary Science Letters, 228, 125-142. GABRIELSE, H. & YORATH, C. J. 1991a. Introduction. In: GABRIELSE,H. & YORATH, C. J. (eds) Geology of the Cordilleran Orogen in Canada. Geology of Canada. Geological Survey of Canada, 3-11. GABRIELSE, H. & YORATH, C. J. 1991b. Tectonic synthesis. In: GABRIELSE, H. & YORATH, C. J. (eds) Geology of the Cordilleran Orogen in Canada. Geology of Canada. Geological Survey of Canada, 677-705. GIBSON, H. D. 2003. Structural and thermal evolution of the northern Selkirk Mountains, southeastern Canadian Cordillera: Tectonic development of a regional-scale composite structural fan. PhD thesis, Carleton University. GIBSON, H. D., BROWN, R. L. & PARRISH, R. R. 1999. Deformation-induced inverted metamorphic field gradients: An example from the southeastern Canadian Cordillera. Journal of Structural Geology, 21, 751-767. GIBSON, H. D., CARR, S. D., HAMILTON, M. A. 8z BROWN, R. L. 2004. Correlations between chemical and age domains in monazite, and metamorphic reactions involving major pelitic phases: an integration of ID-TIMS and SHRIMP geochronology with Y-Th-U X-ray mapping. Chemical Geology, 211, 237-260. GIBSON, H. D., BROWN, R. L. & CARR, S. D. 2005. U-Th-Pb geochronologic constraints on the structural evolution of the Selkirk fan, northern Selkirk Mountains, southeastern British Columbia. Journal of Structural Geology, 27, 1899-1924. GOSCOMBE, B., GRAY, D. & HAND, M. 2005. Extrusional tectonics in the core of a transpressional orogen; the Kaoko Belt, Na_mibia. Journal of Petrology, 46, 1203-1241.

558

R.L. BROWN & H. D. GIBSON

JAMIESON, R. A., BEAUMONT, C., HAMILTON, J. & FULLSACK, P. 1996. Tectonic assembly of inverted metamorphic sequences. Geology 24, 839-842. JOHNSON, B. J. 1994. Structure and tectonic setting of the Okanagan Valley fault system in the Shuswap Lake area, southern British Columbia. PhD thesis, Carleton University. JOHNSON, B. J. & BROWN, R. L. 1996. Crustal structure and early Tertiary extensional tectonics of the Omineca belt at 5 ldegrees N latitude, southern Canadian Cordillera. Canadian Journal of Earth Sciences, 33, 1596-1611. JOHNSTON, D. H., WILLIAMS, P. F., BROWN, R. L., CROWLEY, J. L. & CARR, S. D. 2000. Northeastward extrusion and extensional exhumation of crystalline rocks of the Monashee complex, southeastern Canadian Cordillera. Journal of Structural Geology, 22, 603-625. JOURNEAY, J. M. 1987. Stratigraphy, internal strain and thermo-tectonic evolution of northern Frenchman Cap dome: An exhumed duplex structure, Omineca Hinterland, southeastern Canadian Cordillera. PhD thesis, Queen's University. KRETZ, R. 1983. Symbols for rock-forming minerals. American Mineralogist, 68, 277-279. KUIPER, Y. 2003. Isotopic constraints on timing of deformation and metamorphism in the ThorOdin dome, Monashee Complex, southeastern British Columbia. PhD thesis, University of New Brunswick. Leatherbarrow, R. W. 1981. Metamorphism ofpelitic rocks from the northern Selkirk Mountains, southeastern British Columbia. PhD thesis, Carleton University. LORENCAK, M., BURG, J. P., SEWARD, D., VANDERHAEGHE,O. & TEYSSIER,C. 2001. Low-temperature cooling history of the Shuswap metamorphic core complex, British Columbia: Constraints from apatite and zircon fission-track ages. Canadian Journal of Earth Sciences, 38, 1615-1625. MONGER, J. W. H., PRICE, R. A. & TEMPELMANKLUIT, D. J. 1982. Tectonic accretion and the origin of the two major metamorphic and plutonic welts in the Canadian Cordillera. Geology, 10, 70-75. OKULITCH, A. V. 1979. Lithology, stratigraphy, structure and mineral occurrences of the ThompsonShuswap-Okanagan area, British Columbia. Geological Survey of Canada, Open File 637. OLDOW, J. S., BALLY, A. W. & AVE-LALLEMANT, H. G. 1990. Transpression, orogenic float, and lithospheric balance. Geology, 18, 991-994. PARRISH, R. R. 1995. Thermal evolution of the southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 32, 1618-1642. PARRISH, R. R., CARR, S. O. 8z PARKINSON,D. L. 1988. Eocene extensional tectonics and geochronology of the southern Omineca Belt, British Columbia and Washington. Tectonics, 7, 181-212. PERKINS, J. M. 1983. Structural geology and stratigraphy, Big Bend of the Columbia River, Selkirk Mountains, B.C. PhD thesis, Carleton University.

PRICE, R. A. & MOUNTJOY, E. W. 1970. Geological structure of the Canadian Rocky Mountains between Bow and Athabasca Rivers - A progress report. In: WHEELER, J. O. (ed.) Structure of the Southern Canadian Cordillera. Geological Association of Canada, Special Paper, 6, 7-25. PRICE, R. A., ARCHIBALD, D. • FARRAR, E. 1981. Eocene stretching and necking of the crust and tectonic unroofing of the Cordilleran metamorphic infrastructure, southeastern British Columbia and adjacent Washington and Idaho (abstract). Geological Association of Canada - Mineralogical Association of Canada Annual Meeting, A-47. RANALLI, G., BROWN, R. L. & BOSDACHIN,R. 1989. A geodynamic model for extension in the Shuswap core complex, southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 26, 1647-1653. READ, P. B. & BROWN, R. L. 1981. Columbia River fault zone: southeastern margin of the Shuswap and Monashee complexes, southern British Columbia. Canadian Journal of Earth Sciences, 18, 1127-1145. SCAMMELL, R. J. 1993. Mid-Cretaceous to Tertiary thermotectonic history of former mid-crustal rocks, southern Omineca belt, Canadian Cordillera. PhD thesis, Queen's University. SCAMMELL, R. J. & BROWN, R. L. 1990. Cover gneisses of the Monashee Terrane: a record of synsedimentary rifting in the North American Cordillera. Canadian Journal of Earth Sciences, 27, 712-726. SCHIARIZZA, P. & PRETO, V. A. 1987. Geology of the Adams plateau-Clearwater-Vavenby area. British Columbia Ministry of Energy, Mines, and Petroleum Resources, Paper 1987-2. SEARLE, M. P. & SZULC, A. G. 2005. Channel flow and ductile extrusion of the High Himalayan slab-the Kangchenjunga-Darjeeling profile, Sikkim Himalaya. Journal of Asian Earth Sciences, 25, 173-185. SEVIGNY, J. H., PARRISH, R. R. & GHENT, E. D. 1989. Petrogenesis of peraluminous granites, Monashee Mountains, southeastern Canadian Cordillera. Journal of Petrology, 30, 557-581. SEVIGNY, J. H., PARRISH, R. R., DONELICK, R. A. & GHENT, E. D. 1990. Northern Monashee Mountains, Omineca Crystalline Belt, British Columbia: Timing of metamorphism, anatexis, and tectonic denudation. Geology, 18, 103-106. SHAW, D. 1980. A concordant uranium-lead age for zircons in the Adamant Pluton, British Columbia. Geological Survey of Canada, Paper 80-1C, 243 -246. SIMONY, P. S., GHENT, E. D., CRAW, D., MITCHELL,W. & ROBBINS, D. B. 1980. Structural and metamorphic evolution of the northeast flank of Shuswap complex, southern Canoe River area, British Columbia. In: CRITENDEN, M. L., CONEY, P. J. & DAVIS, G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America, Memoir 153, 445-461. STEVENS, R. D., DELABIO, R. N. ~r LACHANCE, G. R. 1982. Age determinations and geological studies

CHANNEL FLOW IN THE CANADIAN CORDILLERA

K-Ar isotopic ages, Report 15. Geological Survey of Canada, Paper 81-02. TEMPLEMAN-KLUIT, D. J. & PARKINSON, D. 1986. Extension across the Eocene Okanagan crustal shear in southern British Columbia. Geology, 14, 318-321. VANDERHAEGHE,O. ~; TEYSSIER, C. 1997. Formation of the Shuswap metamorphic core complex during late-orogenic collapse of the Canadian Cordillera: Role of ductile thinning and partial melting of the mid- to lower crust. Geodinamica Acta, 10, 41-58. VANDERHAEGHE, O., TEYSSIER, C., McDOUGALL, I. & JAMES, W. 2003. Cooling and exhumation of the Shuswap metamorphic core complex constrained by 4~ thermochronology. Geological Society of America Bulletin, 115, 200-216.

559

WHEELER, J. O. 1963. Rogers Pass Map-area, British Columbia and Alberta (82N West Half). Geological Survey of Canada, Paper 62-32. WHEELER,J. O. 1965. Big Bend map-area, British Columbia. Geological Survey of Canada, Paper 64-32. WHEELER,J. O. & MCFEELY,P. 1991. Tectonic assemblage map of the Canadian Cordillera and adjacent parts of the United States of America. Geological Survey of Canada, Map 1712A. WILLIAMS,P. F. & JIANG, D. 2005. An investigation of lower crustal deformation: evidence for channel flow and its implications for tectonics and structural studies. Journal of Structural Geology, 27, 1486-1504. WILLETT, S., BEAUMONT, C. 8~ FULLSACK,P. 1993. Mechanical model for the tectonics of doubly vergent compressional orogens. Geology, 21, 371-374.

Ductile thrusting versus channel flow in the southeastern Canadian Cordillera: evolution of a coherent crystalline thrust sheet S H A R O N D. C A R R 1 & PHILIP S. S I M O N Y 2

l Ottawa Carleton Geoscience Centre, Department o f Earth Sciences, Carleton University, 1125 Colonel By Drive, Ottawa, ON, K1S 5B6 Canada (e-mail: scarr@ earthsci, carleton, ca) 2Department o f Geology and Geophysics, University o f Calgary, 2500 University Drive NW, Calgary, AB, T2N 1N4, Canada Abstract: The Late Cretaceous Gwillim Creek shear zone (GCSZ) exposed in the core of the Valhalla complex, and located in the hinterland of the southern Canadian Rocky Mountain thrust belt, is a 5 - 7 km thick, easterly verging, ductile thrust zone. It was active after c. 90 Ma and during anatexis (800~ and 800 MPa), rose eastward in the direction of transport, and its base was refrigerated from below at c. 60 Ma by thrust translation onto a cold footwall. Extensional shear zones are younger than the GCSZ, and there is no evidence of channel flow or ductile extrusion. Instead, a 30 km thick, coherent sheet was translated on the GCSZ, which at depth was linked to the Foreland thrust belt such as to form a composite crystalline thrust sheet. Doming of the Valhalla complex may be related to Eocene thrusting beneath the complex during the last stage of shortening. A channel flow, proposed by others for the region north of the Valhalla complex, could have evolved within the crystalline sheet by activation of lateral transition zones and an upper detachment, but is no wider than 250 km, does not represent the dominant orogenic process and may represent a nascent channel.

The origin of > 1 0 k m thick, gently dipping, gneissic, migmatitic and ductilely sheared midcrustal zones exposed in collisional orogenic belts is a topic of m u c h debate, as is the linkage between hinterland ductile structures and structures in the corresponding foreland thrust and fold belt. In the Canadian Cordillera, there are ductile thrust models where: (i) ductile shear zones in the hinterland, that formed during migrating transient shear, are indirectly linked to foreland thrusts (Brown 2004); and (ii) long-lived ductile shear zones, at the base of crystalline thrust sheets, are directly connected to specific major foreland thrusts (Bally et al. 1966; C o o k et al. 1988, 1992; this work). On the other hand, there are channel flow and extrusion models for the Canadian Cordillera, which focus on h o w flow zones a c c o m m o d a t e d deformation in the hinterland (Johnston et al. 2000; Teyssier et al. 2005; Williams & Jiang 2005; B e a u m o n t et al. 2006; Brown & Gibson 2006; Kuiper et al. 2006; Williams et al. 2006). This paper presents a case study from the Canadian Cordillera where the geological relationships fit a composite crystalline thrust sheet model (cf. Hatcher & Hooper 1992), rather than a channel flow model, and a discussion of h o w we reconcile our crystalline thrust sheet model, at the latitude of the Valhalla complex (latitude 4 9 - 5 0 . 5 ~ with channel flow models proposed for the rocks to the north,

at the latitude of the Monashee complex (latitude 50.5-52~ Figs 1 & 2). For the purposes of testing the applicability of channel flow models to geological data, we consider a channel to be a sheet or zone of ductile flow with severe transposition leading to incoherence o f stratigraphy. The zone of ductile flow has an upper ductile detachment zone, with sense of shear opposite to that of a lower ductile detachment, such that material in the channel flows ahead of underlying and overlying crust. The upper and lower shear zones are the same age or have significant age overlap. Similarly, in the model of Williams & Jiang (2005) and Williams et al. (2006), folds of opposite vergence, taken to represent particular levels in the channel or detachment flow zone, must overlap in age. A zone of ductile deformation, with an upper detachment zone that has far less accumulated displacement than the lower detachment, m a y represent incipient channel flow quenched at an early stage, within a crystalline thrust sheet. A zone of deformation that has internal stratigraphy, with significant coherence, and a thin upper ductile detachment, is the ductile equivalent of tectonic delamination rather than an example of channel flow. Models of channel flow (Teyssier et al. 2005; B e a u m o n t et al. 2006; Williams et al. 2006) discuss channel evolution with time. They are

From: LAW, R. D., SEARLE,M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 561-587. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

562

S.D. CARR & P. S. SIMONY

Fig. 1. Tectonic assemblage map of the southern Omineca Belt (simplified from Wheeler & McFeely 1991; modified after Carr 1991; Johnson 1994). The shaded area of the inset locates the map within the morphogeological belts of the Canadian Cordillera: the Insular and Coast belts constitute the western orogen, whereas the Intermontane, Omineca and Foreland belts constitute the eastern orogen. Note the generally north-south striking extensional faults superimposed (e.g. OER ----Okanagan Eagle River extensional shear zone system) on the arcuate trace of compressional structures. Important high-grade culminations include: Frenchman Cap (FC) and Thor Odin (TO) domes in the Monashee complex (MC), bounded on the east by the Columbia River fault (CRF); the Malton gneiss (MG); and the Valhalla complex. The Early Eocene Valkyr (VSZ)-Slocan Lake (SLF) extensional fault system bounds the Valhalla metamorphic core complex. The Late Cretaceous Gwillim Creek shear zone (GCSZ), interpreted as a ductile thrust, is exposed at the deepest exposed structural levels within Valhalla complex. Late Cretaceous-Eocene thrust faults in the Rocky Mountains and those that are related to the Lewis Thrust are shown; other faults and folds in the Rocky Mountain thrust belt are not shown. The dashed 'tip line' represents the map projection of the subsurface location of the transition from ductile detachment or channel flow to ductile thrusting.

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA

563

Fig. 2. West-east cross-section of the SE Canadian Cordillera through southern Valhalla complex (Fig. 1) showing the linkage between the Gwillim Creek shear zone (GCSZ) in the hinterland with the Lewis thrust and related structures in the foreland. Subsurface geometry from the Purcell Anticlinorium eastward is from Cook & van der Velden (1995). otherwise two-dimensional, utilizing the structural thickness and length of the channel in the flow direction. The third dimension, or width, is important for determining the relative significance of the channel as a mechanism in orogenesis. Width is the dimension normal to the dominant flow direction and thickness, and may be approximated by the dimension along-strike. Figure 3a illustrates the terminology used here for a three-dimensional channel with thickness (T), width (W) and length (L). Additional parameters are the duration of flow and flow distance of rock particles. Channels may have basal detachments, such as the Main Central Thrust in the Himalaya, that carry the flowing channel up over the evolving frontal thrust belt, or the channel may be linked at depth to the frontal thrust belt as shown in Figures 2 and 3a, which illustrate an incipient stage in the development of the channel. As flow evolves, material in the channel moves forward, under the suprastructure, towards the tip line where a space problem arises from the fact that the suprastructure above the evolving channel is continuous with the strata of the thrust belt, as well as being continuous with the suprastructure and infrastructure of the adjacent detachment flow zone. Some combination of

structures has to develop in the tip line region to accommodate the excess volume, and such structures are schematically illustrated in Figure 3b. At a nascent stage, these structures would be small, easily overlooked, or open to interpretations that do not recognize channel flow. However, as the volume of the channel flow becomes tectonically significant these structures would become obvious. In Figure 3b, channel extrusion to the surface is shown as one possible end result. The structures shown in Figure 3b were selected merely because they can balance the volume flowing to the tip line. It is of interest that they bear some resemblance to structures in the models of Beaumont et al. (2006), with 'tunnelling' channels. This volume-balancing problem arises only in the case of channel flow, not detachment flow. It is possible that channel flow may be quenched at any stage. In the case where the channel was narrow and channel flow was quenched at an early stage, a nascent channel would be preserved; however, it does not represent the dominant mechanism for the orogen. In homogenous channel flow models for the Himalayan orogen (Grujic et al. 2002; Searle et al. 2003, and references therein), the thickness of the

564

S.D. CARR & P. S. SIMONY

Fig. 3. (a) Block diagram, with central portion cut out, illustrates a channel, of width (W), thickness (T) and length (L), nested inside a flow zone with detachment flow (Williams et al. 2006) outside the channel. The 'tip line' schematically represents the zone of transition from flow in the hinterland to ductile thrusting in the internal portion of a thrust belt. ISC, infrastructure with channel flow; ISD, infrastructure with detachment flow; SS, suprastructure. The zigzag line, on the frontal face of the block, schematically represents the lateral transition from channel flow (ISC) to detachment flow (ISD). Two failed tectonic rock bolts illustrate the difference in flow inside and outside the channel. Channel flow is at an incipient stage. (b) Views of the left wall of central cut-out of the block diagram illustrating structures that could form to balance excess volume brought to the tip line region as channel flow evolved. (i) The tip line region was inserted as a wedge under the rear of the thrust belt with the activation of an 'upper detachment' retro-thrust. (ii) Development of an asymmetric bulge. Layering in the ductile upper detachment is extended into boudins and then warped; a marker horizon is shown folded in front of the bulge. (iii) Evolution of such a bulge leading to extrusion of the channel toward the surface between a normal-sense upper detachment and thrust-sense lower detachment.

High H i m a l a y a n slab (T) is 10 to 20 km; channel width (W) is 1 0 0 0 - 2 0 0 0 km; displacement on the upper South Tibetan d e t a c h m e n t is large and is contemporaneous with large thrust displacement on the basal detachment, the Main Central Thrust; and duration o f flow is 10 to 15 million years (Searle et al. 2003). The H i m a l a y a n orogen is classified as a large hot orogen ( B e a u m o n t et al. 2006). The Canadian Cordillera has large belts o f batholiths and intensely deformed, high-grade m e t a m o r p h i c and migmatitic

rocks, and was therefore classed by Beaumont et al. (2006) as having characteristics that approach those of large hot orogens. Thus, it is o f particular interest to evaluate evidence for, and against, models o f channel flow in the Cordillera. The North A m e r i c a n Cordillera did not involve c o n t i n e n t - c o n t i n e n t collision as in the Himalayan orogen, but was f o r m e d by convergence of ancient North A m e r i c a with an oceanic plate, and accretion o f arc and oceanic terranes. In Canada,

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA

565

Fig. 3b. (Con~nued)

the Cordillera consists of two orogens. The western orogen comprises arc and oceanic terranes that became welded together, and in which the Coast plutonic complex developed (Insular and Coast belts; inset, Fig. 1). The eastern orogen consists of arc, oceanic and pericratonic terranes that were accreted to cratonic North America (Intermontane, Omineca and Foreland belts; inset, Fig. 1). We consider a segment of the eastern orogen, between latitudes 49~ the C a n a d a - U S A border, and 53~

Our focus is on a c. 10 km thick sheet of gently dipping, gneissic, migmatitic and ductilely deformed mid-crustal rocks in the Valhalla complex (Figs 1 & 4), the shear zone that bounds the base of the sheet, and the position of the sheet within an essentially continuous, > 2 5 km thick section. The rocks within the gneissic sheet are folded, but contain a coherent stratigraphic succession in which a right-way-up plutonic edifice was emplaced. This is one of the few places where

566

S.D. CARR & P. S. SIMONY

Fig. 4. Geological map of Valhalla complex and surrounding area modified after Simony & Cart (1997). Normal faults with top-down-to-the-east displacement include the brittle Champion Lakes fault (CLF), the ductile-brittle Columbia River (CRF) and Slocan Lake (SLF) fault zones, and the ductile Valkyr shear zone (VSZ). Note the three culminations of the Valhalla, Passmore and China Creek domes within the Valhalla complex, with the Gwillim Creek shear zone (GCSZ) exposed in their cores. The box in the southern area delineates the location of Figure 6, and lines A - N to C - C I indicate the locations of the cross-sections shown in Figure 5.

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA

567

Fig. 5. West-east geological cross-sections across Valhalla complex in the: (a) Valhalla (A-A'), (b) Passmore (B-B I) and (c) China Creek (C-C') culminations, showing the relationship between the Gwillim Creek shear zones, sheet-like rocks units and extensional structures. See Figures 4 and 6 for location of section lines and legends. Note the difference in scale of section C-C' relative to the other sections. CLF, Champion Lakes fault zone; SLF, Stocan Lake fault zone; VSZ, Valkyr shear zone.

contact relationships can be established at both the top and the base of such a mid-crustal gneissic sheet (Figs 4 & 5). The base is the ductile Gwillim Creek shear zone, with > 40 km of top-to-the-east, Cretace o u s - P a l a e o c e n e age thrust displacement. Near the top of the sheet is a ductile, top-to-the-east, Palaeoc e n e - E o c e n e age extensional shear zone system, for which magnitude of displacement decreases southward from c. 10 kin, in the north, to c. 1 km in the south, where it has not significantly disrupted stratigraphic continuity. These relationships are inconsistent with those predicted by channel

flow models. The shear zones bounding the top and bottom of the deformed zone have the same sense of shear and are of different ages; therefore, we reject channel flow and interpret the shear zone at the base as a major ductile thrust that carried a c. 30 km thick coherent crystalline thrust sheet (Fig. 2). We suggest that doming of the Valhalla complex, which post-dated flow in the gneissic sheet, may be related to Eocene thrusting beneath the complex that accommodated the last stage of shortening in the Rocky Mountain thrust belt.

568

S.D. CARR & P. S. SIMONY

Tectonic evolution of the southeastern Canadian Cordillera The southeastern Canadian Cordillera consists of three major tectonic elements (Fig. 1). (i) The extern a l z o n e of the orogen, the eastern part of the Rocky Mountain Foreland fold and thrust belt, comprises the eastern Rocky Mountains and Foothills. It consists of Palaeozoic and Mesozoic strata of the North American margin and Mesozoic to Palaeocene strata of the Alberta Foreland Basin that were deformed during Cretaceous to Eocene thinskinned shortening above a basal d~collement. (ii) The internal z o n e of the fold and thrust belt occurs in the western Rocky Mountains and ranges to the west. It includes Neoproterozoic and Palaeozoic strata of the North American succession, and their equivalents in the pericratonic Kootenay terrane, that were multiply deformed in the Mesozoic-Palaeogene. (iii) The h i n t e r l a n d z o n e consists of the eastern margin of the inner accreted terranes, namely the late Palaeozoic Slide Mountain marginal basin terrane and mid-Triassic-Lower Jurassic Quesnel volcanic arc terrane, which were accreted to the internal zone in the Middle Jurassic, and multiply deformed in the Mesozoic. Metamorphism and plutonism were superimposed on a belt of rocks that straddles the internal and hinterland zones and is termed the Omineca belt (inset, Fig. 1). It contains low- and medium-grade rocks overprinted by belts of high-grade metamorphism, Cordilleran metamorphic core complexes exhumed during Eocene extension, and clusters of Middle Jurassic, mid-Cretaceous and Palaeocene-Eocene granitoid plutons. The orogen formed during Mesozoic to Eocene transpressional convergence between the North American plate, moving westward as the Atlantic Ocean opened, and an east-dipping subduction zone. Early to Middle Jurassic accretion of terranes via low-angle, low-grade faults, with large displacement, was followed by crustal shortening and thickening in the hinterland and internal zones. In the internal zone, west-verging recumbent nappes, with limbs as large as 3 0 - 4 0 kin, formed in association with low-angle thrusts and regional low-grade metamorphism. During continued progressive deformation, kilometre-scale, Middle Jurassic age, 'F2' west-verging and east-verging folds and fold fans, and widespread greenschist-lower amphibolitefacies metamorphism, were superimposed on existing structures. The folds and faults of this more ductile regime are locally associated with structures such as parallel buckle folds and kink folds, and thrust duplexes typical of the Foreland belt. Folds, with steep axial planes at high structural level curving to gentler dips at depth, are pinned by Middle Jurassic plutons that were quenched and, in

part, exhumed in the Middle Jurassic (Colpron et al. 1996). Between latitudes 48 ~ and 50~ the western margin of the fold and thrust belt is a belt of complex, folded thrusts and fight folds that has a curved or arc-shaped map pattern, termed the Kootenay Arc (Fig. 1). During the Cretaceous-Eocene, deformation propagated eastward and the external Rocky Mountain thrust belt was formed. In the internal zone, Cretaceous-Eocene plutons, produced via crustal melting, are undeformed at high structural levels and sheared at depth, indicating that concomitant deformation in the internal zone was generally restricted to high-grade ductile zones at midcrustal or deeper levels. At the scale of the orogen, the cessation of compression and onset of a regional transtensional regime occurred in the Eocene (Price & Carmichael 1986). In the Omineca belt, this was marked by regional extension and exhumation of Cordilleran metamorphic core complexes, intrusion of syenite plutons and mafic dykes, and normal displacement on steep brittle faults (Parrish et al. 1988, and references therein).

Three crustal zones and the suprastructure-infrastructure association Three crustal zones (Carr 1991) are exposed in the Omineca belt (inset, Fig. 1). The deepest zone is weakly deformed, autochthonous or parautochthonous North American basement, the subsurface continuation of the Canadian Shield (e.g. parts of Frenchman Cap dome). The middle zone consists of rocks that were mobile and at high metamorphic grade mainly in the Cretaceous and/or Palaeocene (e.g. Thor Odin dome, Valhalla complex). The rocks of the upper zone were variably deformed at low metamorphic grade mainly in the Middle Jurassic and Early Cretaceous (e.g. Purcell Anticlinorium). These three zones can be compared to the three levels in the suprastructure-infrastructure association (SIA) of Williams et al. (2006), and it is convenient to refer to the upper and middle levels as the suprastructure and infrastructure (Wegmann 1935; Campbell 1973; Murphy 1987), respectively. However, in order to apply this classification to the southern Canadian Cordillera, it is important to note that the three zones may have undergone penetrative deformation and metamorphism at different times. The infrastructure may be younger, with its metamorphism and deformation superimposed on a pre-existing suprastructure-infrastructure architecture. The suprastructure is dominated by moderately inclined to upright folds, with axial-planar

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA cleavage, and by sharply defined ductile thrusts, in rocks of greenschist and subgreenschist facies. Cross-cutting Middle Jurassic plutons indicate the deformation and metamorphism to have been Middle Jurassic in age. The infrastructure is characterized by tight, recumbent folds, rootless folds, boudinage and strongly foliated and lineated migmatite, paragneiss and orthogneiss. Timing constraints on structures and metamorphism of the infrastructure demonstrate that, in general, the infrastructure formed and/or was reactivated after suprastructure formation. In some regions, such as the NW corner of Figure 1, infrastructure was reactivated and was hot and mobile in the Early Cretaceous only (Scammell 1993; Parrish 1995, and references therein; Simony & Carr 1997; Digel et al. 1998). In some regions, such as the area east, west and NW of Frenchman Cap dome (Scammell 1993; Brown & Gibson 2006, and references therein), the infrastructure was hot and mobile in the Early and Late Cretaceous, while elsewhere, such as in the Thor Odin and Valhalla domes, only Late Cretaceous (Spear & Parrish 1996; Johnston et al. 2000) or Palaeogene (Hinchey 2005) ages have been obtained. Locally, however, in the NW corner of Figure 1, and perhaps in the southern part of the Valhalla complex (Figs 1 & 4), a transition zone is preserved from the Jurassic suprastructure to a Jurassic infrastructure with recumbent folds and gently dipping foliation (Pell & Simony 1982; Murphy 1987; Digel et al. 1998). The infrastructure is exposed in large metamorphic-structural culminations (Crowley et al. 2000; Gibson et al. 2005; Ghent & Simony 2005), and in fault-bounded extensional core complexes in the Omineca belt (Parrish et al. 1988). There are important variations from complex to complex, and no single mechanism explains all the infrastructure exposures. The deepest crustal zone, the deepest level of the suprastructure-infrastructure association, is only exposed in Frenchman Cap dome, the northern part of the Monashee complex (Fig. 1). Basement exposed elsewhere in the Omineca belt (e.g. Malton gneiss, Thor Odin dome) has been remobilized in the middle crustal zone or infrastructure.

Architecture of a coherent thrust sheet with the Valhalla complex near its base A c. 10 km thick section of complexly folded, amphibolite-facies paragneiss, and variably deformed intrusive rocks is exposed in the Valhalla complex (see Reesor 1965; Cart et al. 1987; Simony & Cart 1997; Schaubs & Cart 1998). The complex is a fault-bounded, doubly plunging, broad antiform

569

with three distinct culminations: the Valhalla, Passmore and China Creek domes (Fig. 4). A downward-increasing, Late Cretaceous strain gradient culminates in the Gwillim Creek shear zone at the base of the section, exposed in the cores of all three domes (described in a later section). The Valhalla complex is a Cordilleran metamorphic core complex, and represents infrastructure that was exhumed in the Eocene on an extensional fault system (Parrish et al. 1988). The fault system has two components: the ductile Valkyr shear zone, which is arched over the Valhalla complex and forms its upper boundary; and the ductilebrittle, east-dipping, Slocan Lake-Champion Lakes normal fault system, bounding the eastern margin of the complex (Figs 4 & 5; Cart et al. 1987; Simony & Carr 1997, and references therein). This top-to-the-east composite fault system is hinged at the southern end. Magnitude of displacement decreases from > 1 0 km in the north, at the latitude of Valhalla dome, to 1.5 to 0.5 km in the southern China Creek dome (Simony & Carr 1997). In the south, where extensional displacement is small, the pre-extensional relationships between rocks of the complex and those in the hanging wall of the extensional fault system are preserved, showing that there is no major tectonic boundary between the infrastructure and overlying suprastructure. Therefore, the Valhalla complex, and structurally overlying rocks, together constitute a coherent thrust sheet, on the order of 30 km thick, with the Late Cretaceous Gwillim Creek shear zone at its base. This coherent crystalline thrust sheet is mainly composed of rocks of the Quesnel arc terrane, such as the Triassic Slocan Group and volcanic rocks of the Early Jurassic Rossland Group, and underlying Permo-Carboniferous strata. During the early Mesozoic accretion and crustal shortening episode, they were repeated by thrust faults; folded into upright to steeply inclined, gently plunging folds; metamorphosed at chlorite, biotite and garnet grade; and then intruded by Middle Jurassic granodiorite-tonalite plutons. In the lower part of the thrust sheet, these rocks and structures underwent episodes of intrusion and penetrative deformation during Cretaceous-Palaeogene times. In China Creek dome, the basement to the Quesnel arc terrane is present, consisting of the Devonian Trail gneiss and Carboniferous to Triassic Mount Roberts Formation (Figs 6 & 7). Rocks of the Quesnel terrane overlie them unconformably, and all are stitched together by the Middle Jurassic plutonic suite. A more inboard section of the composite accreted terrane occurs in the Valhalla and Passmore domes, within the Valhalla assemblage of Schaubs & Carr (1998) (Fig. 4). The assemblage includes rocks correlated with Palaeozoic North

570

S.D. CARR & P. S. SIMONY

Fig. 6. Detailed geological map of the southern Passmore dome and the China Creek dome, showing relationships between map units and structures such as the Eocene Valkyr (VSZ)-Slocan Lake (SLF)-Champion Lakes (CLF) extensional fault system that bounds the complex, the upper limit of Late Cretaceous strain and the Late Cretaceous Gwillim Creek shear zone (GCSZ), a zone of high strain exposed in the deepest structural level in the core of the China Creek culmination. All rock units structurally below the Valkyr shear zone, that in map view lie between the trace of the Valkyr shear zone and the Slocan Lake-Champion Lakes fault zone, comprise significant proportions of Early Cretaceous grey gneiss at depth and Palaeocene-Eocene leucogranite, pegmatite and aplite sheets at higher levels. Line C-C' shows the location of cross-section in Figure 5c.

American margin and pericratonic stratigraphy, as well as a slice of allochthonous Slide Mountain terrane and part of the Quesnel arc. The trace of Early Jurassic faults, that would have accommodated the collapse of the marginal basin and accretion of Slide Mountain and Quesnel terranes, likely occur within the Valhalla assemblage, but have been transposed by younger deformation a n d / o r obscured by intrusions. Plutonic rocks are important within the crystalline thrust sheet as they demonstrate stitching relationships and act as strain markers. Their spatial disposition is illustrated in Figures 4 - 7 . In the upper part of the sheet, plutons of Middle Jurassic, mid-Cretaceous and Eocene ages are c o m m o n and are generally not penetratively deformed. In contrast, in the lower part of the edifice, in the Valhalla complex, the igneous rocks were penetratively deformed. Most important for this study are Middle

Jurassic biotite hornblende tonalite-granodiorite plutons that generally structurally overlie the Valhalla complex, but also occur at all depths in the China Creek dome. In map view, they ring the complex, cover an area of some 4000 to 5000 k m 2 (Fig. 4), and may have extended across the complex prior to erosion. These plutons are composite laccoliths and sheets with flat bases, and are linked by feeder sills, all at about the same structural level. In Figure 7, for the sake of clarity, Middle Jurassic plutons are shown well above the Gwillim Creek shear zone, but, in fact, the lower parts of the Mackie and Bonnington plutons are involved in the strain related to the Gwillim Creek shear zone. In the China Creek dome, a Late Jurassic or Early Cretaceous foliated, biotite granodiorite orthogneiss, termed the 'grey gneiss', is ubiquitous and occurs as 30 to 100 m thick sheets (Fig. 5c). Grey

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA

571

Fig. 7. Schematic sketch of a generalized cross-section, looking west, through the c. 30 km thick thrust sheet transported by the Gwillim Creek shear zone (GCSZ) at its base. The sketch illustrates the relationships between rock units, particularly sheets of igneous rock and plutons, and shows the division of the thrust sheet into its suprastructure and infrastructural Valhalla complex. Dashed wavy bold lines represent unconformities.

gneiss sheets were emplaced within the previously deformed Trail Gneiss complex and infolded Mount Roberts Formation. Sheets of this suite also intruded Middle Jurassic tonalite, and, at the deepest levels exposed in the China Creek dome, grey gneiss included slabs from feeder sills of the Middle Jurassic Mackie pluton. The grey gneiss was itself intruded by the mid-Cretaceous, c. 110 Ma (Parrish 1995) Kinnaird phiton. The Kinnaird pluton, south and west of Castlegar, consists of multiple, gently dipping porphyritic granite and granodiorite sheets, each 50 to 200 m thick (Halwas & Simony 1986). It includes screens and enclaves of the grey gneiss and lies mainly at and below the level of the Middle Jurassic plutons that it locally cuts. The c. 63 Ma porphyritic Airy quartz monzonite occurs as a sheet that occupies much of the Passmore dome (Can" et al. 1987; figs 5b & 6). It occurs just north of the Kinnaird body, at and below its level. Downward-increasing strain gradients occur in the Kinnaird and Airy quartz monzonite sheets and are attributed to motion on the Gwillim Creek shear zone (Simony & Carr 1997, and references therein; this work). While these sheets of igneous rocks were injected and deformed in lower levels of the crystalline thrust sheet, the

Late Cretaceous to Eocene plant-bearing clastic rocks were being unconformably deposited on the top (Little 1982). The Middle Jurassic, Late Jurassic to Early Cretaceous, mid-Cretaceous and Palaeocene plutonic suites, linked by intrusive relationships, form an interconnected, right-way-up plutonic edifice that extends from the Gwillim Creek shear zone, at the base, up to low-grade Jurassic strata. They bolt the coherent but deformed tectonostratigraphic succession into an upright, coherent sheet with a Cretaceous top-to-the-east shear sense and associated penetrative strain that increases downward to the Gwillim Creek shear zone at its base (Fig. 8). Individual plutons, such as the Trail and Kinnaird plutons, extend from the level affected by shear strain up into the suprastructure where they are not strained, and where originally gently dipping roof structures and gently dipping intrusive sheets are preserved. The strained basal contacts of these plutons are oriented at low angles to their unstrained upper contacts. A sill complex interconnects the Middle Jurassic plutons at their bases in both east-west and north-south directions (Fig. 4.) It is unlikely that high strain could mimic such three-dimensional interconnections, and therefore

572

S.D. CARR & P. S. SIMONY

Fig. 8. Schematic sketch illustrating the gradient of downward-increasing strain within a representative Middle Jurassic pluton from above the upper limit of Cretaceous deformation to the Gwillim Creek shear zone. Fold style and metamorphic index minerals in metasedimentary rocks adjacent to the pluton are also indicated. (a) In this zone the pluton contains igneous textures (e.g. randomly oriented crystals) and is undeformed; foliation occurs only locally in the pluton. (b) Plutonic rocks are pervasively foliated but the foliation is weak and the igneous origin of the rock is obvious. (c) The pluton is strongly foliated and lineated but contains lenses and domains where the plutonic protolith is still obvious; these more weakly strained domains become smaller in size downward. (d) Strongly lineated and foliated orthogneiss. (e) Mylonite layers are present in the orthogneiss and the protolith is difficult to recognize. Leucosome occurs as foliated streaks, and lenses of foliated pegmatite are present. BT, biotite; CHL, chlorite; FSP, feldspar; GRT, garnet; MS, muscovite; QZ, quartz; SIL, sillimanite.

the feeder sill system as well as the flat floors of the plutons must be of intrusive origin, albeit strained. There is a complete transition from a suprastructure to an infrastructure, with penetrative flow, that carried the suprastructure eastward (Fig. 5c). Above the level of high strain associated with the Gwillim Creek shear zone, the individual plutons show no offset contacts or significant distortions. This demonstrates that no shear zone, either mapped, like the Valkyr shear zone, or as yet undetected, could have displacements of more than 5 0 0 - 1 5 0 0 m (Simony & Cart 1997). Such a 'bolted' suprastructure-infrastructure transition precludes any

channel flow model where the channel material flowed ahead with respect to its suprastructure.

Upper margin of the Valhalla complex (infrastructure), Ladybird granite and Valkyr shear zone Throughout the Valhalla complex, the upper limit of intense C r e t a c e o u s - P a l a e o c e n e strain and metam o r p h i s m (Figs 4 - 6 ) occurs near the base of the Middle Jurassic sills and laccoliths, and that level is occupied by 5 9 - 5 5 M a sheets of biotite

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA leucogranite or pegmatite with minor muscovite, garnet and tourmaline (Figs 4 & 5a,b). This is the Ladybird granite and it resembles OligoceneMiocene leucogranites exposed in the Himalaya (Searle 1999). In the China Creek dome, leucogranite occurs as individual sheets 5 to 30 m thick, associated with bodies of aplite and pegmatite injected just below and at the base of the Jurassic laccoliths. In the Valhalla and Passmore domes, the granite forms one continuous layer, 1 to 3 km thick. The sheet is at least 80 km long, 30-40 km wide and represents a volume of >4500 km 3. The injection of leucogranite appears to have been controlled everywhere by the bases of Middle Jurassic plutons, even though they are at a level 2.5 km deeper in the China Creek dome than in the Passmore and Valhalla domes (Simony & Carr 1997). The deepest leucogranite sheets locally cross-cut fabrics and layering associated with the Gwillim Creek shear zone. In Passmore dome, the Ladybird granite cuts the top of the 63 Ma Airy quartz monzonite sheet. These field relationships are consistent with U - P b age dates indicating that the Ladybird granite post-dated the Gwillim Creek shear zone and is not directly linked to its evolution and motion. The Valkyr shear zone lies along the contact between the top of the leucogranite and the base of the Middle Jurassic plutons. Field mapping and structural analysis (Cart et al. 1987; Simony & Carr 1997) showed that the Valkyr shear zone is a ductile, top-to-the-east extensional shear zone. Ductile strain in the Ladybird granite increases upward to the Valkyr shear zone and is related to it. Locally, late-phase granites intruded across the Valkyr shear zone and have little or no penetrative strain. Elsewhere, sheets of leucogranite emplaced into the upper part of the Valkyr shear zone have a strong lineation and foliation but are surrounded by Middle Jurassic tonalite with virtually no structural fabric. These relationships, taken together, are interpreted to show that the Valkyr shear zone propagated near the contact between the suprastructure, stiffened by old cold plutons, and the warmer infrastructure, further heated by the injection over 4 million years of a >4500 km 3 volume of leucogranite. Motion on the Valkyr shear zone and leucogranite injection overlapped in time. It is more likely that the top of the leucogranite localized the shear zone for the thermomechanical reasons suggested above, than the converse where the shear zone localized the leucogranite. The base of the thick Ladybird granite sheet in the Passmore dome is up to 3 km deeper than the level of Valkyr shear zone strain, and no evidence has been seen in the Valhalla complex for localization of leucogranite in dilatent zones related to the Valkyr shear zone. With the exception of the northwestern part of the Valhalla dome, the Ladybird leucogranite was emplaced

573

where neither the country rock composition nor its metamorphic state favoured in situ melting. The Ladybird granite may resemble OligoceneMiocene leucogranites in the Himalaya, but its tectonic relations differ. It is largely younger than peak metamorphism, shear and melting in the Gwillim Creek shear zone, which is the main crustal (detachment) flow zone. As in the Himalaya (Searle 1999), the Ladybird granite sheets lie close to an upper detachment zone; however, the Valkyr shear zone is not analogous to the South Tibetan detachment. Its shear sense is top-to-the-east, the same as the structurally deeper Gwillim Creek shear zone. Its displacement is much less than that of the Gwillim Creek shear zone, and it is entirely younger. In the China Creek dome, displacement on the Valkyr shear zone is probably less than 1 km (Simony & Carr 1997) and there is complete continuity from the suprastructure to the infrastructure. The Ladybird granite is not linked to channel flow, as has been proposed for the Himalayan leucogranites.

The Gwillim Creek shear zone The Late Cretaceous top-to-the-east Gwillim Creek shear zone marks the base of the Valhalla complex. Two strands of the shear zone are exposed in the Valhalla dome (Parrish et al. 1987a), where sheets of foliated and veined, c. 110 Ma Mulvey granodiorite gneiss (dated by Parrish et al. 1985, 1988; Heaman & Parrish 1991) are tectonically interleaved with paragneiss sheets (Figs 4 & 5a). All the rocks are penetratively deformed; however, high strain is localized at the lower boundaries of the two Mulvey gneiss sheets, and they are interpreted as having been repeated in an antiformal stack duplex (Fig. 5a). In Passmore dome, the shear zone is marked by a gradient of downwardincreasing strain within migmatitic paragneiss, in the core of the culmination (Schaubs & Carr 1998). It projects directly into a gently arched package of strong seismic reflectors, termed the Valhalla reflector, imaged on a Lithoprobe seismic profile through the Passmore dome (Cook et al. 1988; Eaton & Cook 1990). In the deepest levels of the China Creek culmination, a gradient of generally Late Cretaceous downward-increasing strain, interpreted as the Gwillim Creek shear zone (Simony & Carr 1997; this work), is superimposed on the lower part of the stratigraphic and plutonic edifice described above. In China Creek dome, the approximate location of a Cretaceous metamorphic and deformation 'front' is delineated in the map and cross-section on Figures 5c & 6. Where the deeper parts of the Middle Jurassic plutons are exposed on the sides of valleys, a strong ductile deformation fabric is

574

S.D. CARR & P. S. SIMONY

superimposed on the igneous textures and fabrics (Fig. 8). This characteristic of the Middle Jurassic Trail, Mackie and Bonnington plutons (Figs 4-6), as well as deformation in the mid-Cretaceous Kinnaird pluton, indicates that intense post-midCretaceous ductile deformation took place at depth. Although the weaker foliation at high levels in the Jurassic plutons could be Early Cretaceous and pre-date the Kinnaird pluton, we show that this deformation at depth was mainly Late Cretaceous in age, accompanied by high-grade Barrovian metamorphism and associated with shear on the Gwillim Creek shear zone. The progressive change in fabric and the increasing amount of strain with depth are similar in all of the plutons and broadly correlate with fold style and metamorphism in adjacent metasedimentary units, as summarized in Figure 8. The Mackie and Kinnaird plutons best illustrate the complete transition with depth from undeformed plutonic rocks with igneous fabrics to intense ductile deformation fabrics. At the deepest level an intense foliation and lineation, an L-S tectonite fabric, is pervasive; mylonite layers occur between gneissic layers and the plutonic origin of the rocks can only be surmised. In paragneiss and orthogneiss, like the Trail gneiss and grey gneiss, the rocks become streaky and banded with an intense transposition foliation, local mylonite layers and small rootless folds. This gradient of downward-increasing strain is also present in the c. 110 Ma Kinnaird body and the grey gneiss below it, where sheared asymmetric boudins, sheared and dismembered veins and granitic sheets, sheared-out fold limbs and shear strain fabrics in feldspars were interpreted to indicate zones of top-to-the-east shear (Halwas & Simony 1986). Such highly deformed rocks resemble those of the Late Cretaceous Gwillim Creek shear zone where it is exposed in Passmore dome (Schaubs et al. 2002) and the Gwillim Creek shear zone was mapped on that basis on the valley bottom in China Creek dome (Fig. 6). The most southerly control point on the Gwillim Creek shear zone is provided by a strain gradient in the basal 500 m of the SE part of the Trail pluton, exposed in the Columbia Valley in the footwall of the Champion Lakes fault (Fig. 4). Comparison with our 'standard gradient' (Fig. 8) suggests the Gwillim Creek shear zone may be 500 to 1500 m below the valley floor.

Timing constraints on the Gwillim Creek shear zone An internally consistent data set, based on field and microstructure relationships and geochronology studies, constrains movement of the Gwillim

Creek shear zone to be Late Cretaceous to Palaeocene in age (c. 90 to 60 Ma). Deformation was younger than the c. 110 Ma Kinnaird and Mulvey granitoids, synchronous with or younger than the c. 90 Ma China Creek pegmatite (see below), and synchronous with high-grade metamorphism and anatexis (Schaubs et al. 2002). A younger age limit for motion on the shear zone is constrained by a downward-increasing strain gradient within the c. 63 Ma Airy quartz monzonite. Its base, located about 1 km above the Gwillim Creek shear zone, is foliated with flattened and lineated quartz and aligned feldspar (Figs 4, 5b & 6). At c. 63 Ma, the Airy quartz monzonite had just been intruded and was hot, ductile and weaker than the surrounding rocks. Its age therefore provides an upper limit for the age of ductile shearing on the Gwillim Creek shear zone below. Late Palaeocene and Early Eocene granitoids clearly cross-cut the Gwillim Creek shear zone. These include sheets of the c. 59-55 Ma Ladybird granite (Carr et al. 1987) and a c. 52 Ma pegmatite dyke in the Valhalla dome (Parrish 1992), as well as undeformed and unfoliated apophyses of the 53 Ma College Creek pluton in foliated grey gneiss of the China Creek dome. By c. 59 Ma motion on the Gwillim Creek shear zone had ceased, between c. 59 and 55 Ma the extensional Valkyr-Slocan Lake shear zone was active, and by 52 Ma the entire Valhalla complex had cooled below the brittle-ductile transition and was involved in Eocene crustal extension concomitant with emplacement of mildly alkaline plutons and dykes with chilled margins and brittle fracture-controlled intrusive contacts. Formation of the Valhalla antiform had probably ended because the dykes do not show obvious fanning or warping across it. The sequence of changes in intrusive style through time is consistent with the temperaturetime cooling curve produced by Parrish (Heaman & Parrish 1991) from the closure temperatures of several chronometers. Timing of motion on the Gwillim Creek shear zone was further refined by detailed petrography and extensive geochronology (Parrish et al. 1988; Heaman & Parrish 1991; Parrish 1995; Spear & Parrish 1996). The thermal peak of metamorphism occurred at c. 7 2 - 6 7 Ma. Schaubs et al. (2002) determined P and T conditions from microdomains where reactions were synchronous with deformation. Correlation of micro- and mesoscopic deformation fabrics provides P-T constraints on the Gwillim Creek shear zone at different structural levels in Passmore and Valhalla domes. Peak conditions of 825~ and 730 MPa in a zone above the Gwillim Creek shear zone, and 850~ and 840 MPa within the shear zone, were synkinematic with respect to the transposition foliation within the

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA zone. Schaubs et al. (2002) also demonstrated that net transfer and exchange reactions were quenched shortly after the metamorphic peak in the deepest samples, an observation consistent with cooling from below by translation onto a cold footwall (Spear & Parrish 1996). On the basis of diffusion modelling of retrograde net transfer reactions on garnet rims in pelitic migmatites, in conjunction with in situ Pb-Th ion microprobe geochronology on monazites from two samples near the Gwillim Creek shear zone in Passmore dome, Spear (2004) interpreted ages he obtained in the 85-75 million year age range to represent prograde metamorphism. He also proposed that near-peak conditions, with melt present, at 60 + 2 Ma were followed by a short period of rapid (200~ Ma-1) almost isobaric cooling. He concluded that this required refrigeration from below and, like Spear & Parrish (1996) and Schaubs et al. (2002), suggested that this cooling was accommodated by transport of the hot Valhalla complex up a thrust ramp and onto a cool thrust fiat at a rate of several centimetres per year. His two-dimensional thermal modelling suggested a 10-20 ~ ramp, some 15 km high. This interpretation is consistent with the hanging wall ramp identified on the basis of mapped geometry described in the following section. The high cooling rate of 200~ Ma -1 at the deepest exposed level (Spear 2004), and the substantial decrease in cooling rate with distance from the shear zone (Schaubs et al. 2002), suggest that the footwall lay not far (c. 1-2 km?) below the deepest currently exposed parts of the Passmore dome. This also suggests that by c. 60 Ma the Gwillim Creek shear zone had a relatively abrupt, highly strained and layered base with aphanitic mylonitic interlayers in micaceous gneiss that had an intense S foliation, giving a strong seismic anisotrophy, and perhaps explaining its pronounced seismic reflectivity (Jones & Nur 1982; Eaton & Cook 1990). In the core of the China Creek culmination, biotite granodiorite sheets of the grey gneiss are streaked with layers and lenses of strongly foliated leucosome, and lenses and sheets of pegmatite termed the China Creek pegmatite. The China Creek pegmatite cuts the Early Cretaceous grey gneiss and the base of the mid-Cretaceous Kinnaird pluton, and is in turn cut by the Eocene Ladybird pegmatite. The China Creek pegmatites are generally concordant with the foliation, particularly within the Gwillim Creek shear zone where they are strongly deformed, foliated, lineated and boudinaged, indicating that the pegmatite was pre- or syntectonic with respect to Gwillim Creek shear zone strain. A deformed China Creek pegmatite from the deepest well-exposed level in the core of the

575

China Creek culmination was mapped and sampled for U - P b geochronology to better constrain the age of the pegmatite, and to test the interpretation that shearing in the China Creek dome is indeed related to the Gwillim Creek shear zone. Details of the analytical methods, results and age interpretation are given in the Appendix. The best estimate of crystallization age for the pegmatite is 90 __+ 0.3 Ma; monazites in the pegmatite underwent metamorphic crystallization and/or recrystallization in the Late Cretaceous, likely between c. 77 and 88 Ma, and this dates at least part of the ductile shearing. These ages fall into the age range obtained for prograde metamorphism and associated deformation in the Passmore dome, and this is consistent with the structural interpretation that the Gwillim Creek shear zone is exposed in the China Creek dome. The apparent spatial association of leucosome and pegmatite with the Gwillim Creek shear zone suggests that the c. 90 Ma age of the China Creek pegmatite dates an early episode of motion on the Gwillim Creek shear zone.

Geometry and depth of the Gwillim Creek shear zone The identification of the Gwillim Creek shear zone at deep structural levels in all three culminations (Fig. 4) is consistent with the geometry of the Valhalla, Passmore and China Creek culminations and intervening structural depressions. The trajectory of Gwillim Creek shear zone strain in the three east-west cross-sections of Figure 5 suggests that the shear zone progressively ramped up through the plutonic-metamorphic edifice in the direction of eastward transport, In all three sections, the eastward rise of the Gwillim Creek shear zones suggests preservation of a hanging wall ramp with a 'ramp angle' of about 5 ~ to 10~ In the China Creek culmination, the depth of the Gwillim Creek shear zone during flow is determined from the thickness of section between the Late Cretaceous shear zone and a synchronous unconformity surface. Late Cretaceous plant-bearing sandstone and conglomerate lie unconformably on the Lower Jurassic Rossland Group in the SW corner of Figure 4 (Little 1982). Taking into account the typical structural thickness of the Rossland Group, Mount Roberts Formation, Trail gneiss complex and Middle Jurassic plutonic sheets, we can make a rough estimate of not more than 15 km for the Late Cretaceous depth of the Gwillim Creek shear zone, leading to a pressure of some 450 MPa. This is consistent with the metamorphic mineral assemblage of quartz + biotite + muscovite + s'dlimanite + garnet in local semipelitic layers, and the

576

S.D. CARR & P. S. SIMONY

SOUTH

NORTH future China Creek culmination

future Passmore culmination

Late Cretaceous ground surface

_

_

Approximateroof/eve/ofMiddleJurassicp / u t o ~ ~ - = ~ , ~

Nelson Batholith

BonningtonPluton

Tra#

~.~-..P3~n

10_ 9die jura 15. wel o

~

r

~

~

-

. / ~ " ~ ' " ~ MPa

20_

25800 MPa

30km

Late Cretaceous geometry of the Gwillim Creek shear zone (GCSZ) - looking east in the direction of transport

Fig. 9. Schematic north-south sketch showing the southward rise of the Gwillim Creek shear zone with depth relative to a Late Cretaceous land surface as represented by the unconformity at the base of Late Cretaceous conglomerate. The geometry is illustrated as it existed while the Gwillim Creek shear zone was active in the Late Cretaceous, just prior to the sheet being carried up the frontal ramp and onto the upper flat. persistence of muscovite to the deepest levels exposed where flattened and elongate aggregates of quartz + plagioclase + minor K-feldspar have formed within grey gneiss. In the Valhalla and Passmore culminations, pelitic paragneiss contains the syntectonic mineral assemblage quartz + biotite § garnet § sillimanite § K-feldspar § granitic melt, indicating peak metamorphic conditions of approximately 800~ and 750 to 800 MPa in the Late Cretaceous (Spear & Parrish 1996; Schaubs et al. 2002; Spear 2004). Compared with the China Creek culmination, this suggests a southward rise in the Late Cretaceous Gwillim Creek shear zone of some 10 to 12 km from a depth of 26 to 28 km in the Passmore culmination, to c. 15 km in the China Creek culmination 9 Figure 10 depicts the relative positions of geological elements in the Late Cretaceous before development of the Valhalla antiform and the broad regional southward plunge of much of the complex. This southward rise of the Gwillim Creek shear zone through the plutonic edifice in its hanging wall can be deduced from the maps, and by comparing section B - B ' through the Passmore culmination with C - C ' through the China Creek culmination (Fig. 5). This southward rise

represents a lateral ramp. Its height, estimated at 10 to 13 km, also approximates the height of the frontal ramp because, even in inffastructural zones, frontal ramps at early stages of their evolution have heights comparable to that of their related lateral ramps. The ramp probably formed at the site of a lateral rheological contrast, induced not by the introduction of different crustal material into the flow zone, but by a horizontal thermal gradient where east-dipping isothermal surfaces separated hotter, weaker rocks in the west from cooler, stiffer rocks in the east and south, all within the same gneiss sheet. Ramps and flats are structures typical of the external thrust belt of the foreland, but ramps like the one documented here are, in fact, predicted for the infrastructure of the internal zone by the models of Beaumont et al. (2006). The evidence for both the frontal and lateral ramps is preserved in the hanging wall sheet because the well-defined plutonic edifice permits tracking the rising shear zone, and the north to south decrease in metamorphic pressures is recorded in distinct mineral assemblages. The lateral ramp is illustrated in Figure 9 as it would have appeared before the hanging wall thrust

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA sheet travelled up and over the ramp. The step in the shear zone is not preserved in the Valhalla complex because the hanging wall sheet flowed up the ramp and onto the upper flat where the ramp became inverted. The 800 and 450 MPa Late Cretaceous mineral assemblages, from Passmore and China Creek domes, respectively, are now at the same elevation, and the base of the Middle Jurassic plutons drops southward toward the Gwillim Creek shear zone. The complementary footwall ramps were left behind in the subsurface tens of kilometres to the west, where they were obliterated by younger ductile shear. Therefore this ramp geometry is not seen in the Lithoprobe seismic profile across Passmore dome (Eaton & Cook 1990).

Large displacement on the Gwillim Creek shear zone and its link to thrust fault systems in the Rocky Mountains Given that the frontal and lateral ramps of the Gwillim Creek shear zone were 10-12 km high, and had an estimated dip of 5 - 1 0 ~ then a displacement of at least 40 km was required to move the Valhalla complex rocks up the ramp and onto the 'fiat' to the east. This part of the motion occurred between c. 65 and 60 Ma; however, since the shear zone was active between c. 90 and 60 Ma then the total displacement must have been much greater, and an estimate of 80 to 100 km is not unreasonable. Clearly the Gwillim Creek shear zone was a major structure, accommodating significant relative eastward transport of the hinterland toward the foreland. Therefore, it is reasonable to expect that it has lateral continuity, or width, well beyond the 90 km strike length established by mapping of the Valhalla complex, and that large displacement must be accommodated by linkage with large structures to the east. Late Cretaceous to Palaeocene motion on the Gwillim Creek shear zone corresponds to a period of major thrusting in the external part of the Foreland fold and thrust belt, and on this basis, linkage between the Gwillim Creek shear zone and Rocky Mountain thrusts has been proposed (Parrish et al. 1987a; Spear & Parrish 1996; Schaubs et al. 2002; Spear 2004). Varsek & Cook (1994) and Cook & van der Velden (1995) proposed direct linkage with transfer of displacement beneath the Purcell Anticlinorium to the Rocky Mountain basal drcollement. The crustal-scale geologicaland geophysical-based cross-section of Figure 2 is modified from that of Cook & van der Velden (1995), and extended westward across the Kootenay Arc, into the Quesnel terrane, and across the China Creek culmination of the Valhalla complex. It illustrates a possible direct link of the Gwillim

577

Creek shear zone to the Lewis thrust of the Rocky Mountains, one of the foreland thrusts with large displacement and strike length (Price 1981; Fermor 1999). The Lewis thrust has a displacement of 60 to 100 km near latitude 49~ and a strike length of c. 200 kan on either side of latitude 49~ Apatite fission track dating (Ozadetz et al. 2004) indicates that at c. 75 + 5 Ma a major ramp was activated, and Mesoproterozoic and Palaeozoic strata of the Lewis thrust sheet were transported up the ramp and onto a broad flat on CampanianMaastrichtian foreland clastic sediments, and then cooled. 4~176 dating of clay gouge on the Rundle thrust, a fault linked to the Lewis thrust near its northern termination, suggests motion at c. 73 Ma (van der Pluijm et al. 2001). Given the large size of the Lewis thrust sheet, and the large displacement on the Lewis thrust, motion may well have begun before c. 80 Ma and continued after c. 70 Ma. Motion that took place on the Gwillim Creek shear zone before c. 80 Ma could correspond to motion on older thrusts immediately west of the Lewis thrust, that are linked to the same d~collement. Shear between c. 65 and 60 Ma could correspond to motion on thrusts east of the Lewis thrust that also share the same drcollement (Fermor & Moffat 1992). Given that the Gwillim Creek shear zone in Valhalla complex is a major ductile thrust zone of Late Cretaceous-Palaeocene age, it has to have a subsurface trajectory somewhat like that shown in Figure 2. On the east side of the Valhalla complex, it dips to the east and has to pass under the eastern margin of Quesnel Terrane and under the Kootenay Arc (Fig. 2). The main deformation in the Kootenay Arc, the Purcell Anticlinorium and the western ranges of the Rocky Mountains was Jurassic and Early Cretaceous, and pre-dated intrusion of largely undeformed mid-Cretaceous plutons (Archibald et al. 1983; Colpron et al. 1996; Larson et al. 2004). The Gwillim Creek shear zone cannot rise into the Kootenay Arc, Purcell Anticlinorium or Western Rockies because no major shear zones or thrusts of Late Cretaceous age have been documented, although there are secondary structures that could be Late Cretaceous in age. The shear zone is drawn such as to suggest that it reactivated the Early to Middle Jurassic terrane-accretion thrust, truncated and sheared the steep Kootenay Arc structures at depth and decapitated previously emplaced basement slices under the Purcell Anticlinorium. Although the exact trajectory of the Gwillim Creek shear zone is speculative, there is little choice but to link it with the Lewis thrust and related faults in the Rocky Mountains. The thrust faults of the western Rocky Mountains and the basal drcollement of the Lewis thrust pass

578

S.D. CARR & P. S. SIMONY

westward under the Purcell Anticlinorium, perhaps as a distinct, mylonitic, ductile thrust that is represented by a single line at the scale of Figure 2. Somewhere under the western Purcell Anticlinorium or under the Kootenay Arc, it must pass into the westward-thickening and ductile Gwillim Creek shear zone, a zone many kilometres in thickness, where eastward-directed 'detachment flow' (Williams et al. 2006) carried a thick sheet and 'balanced' the thrusting further east. This conclusion is not altered by the very real problem, discussed by Williams et al. (2006) of calculating shortening (or stretching) in a flow zone. The problem does not arise as long as the suprastructure is merely carried by the flow zone. Location of the transition from ductile detachment flow in a sheet to thrusting of stiff sheets on ductile thrust zones may be approximated by the 'tip line' of the thick Gwillim Creek shear zone shown in Figure 2. Models that may be mechanically more appealing (e.g. Brown 2004) invoke transient high-strain zones, in deep-seated rocks, that migrate through the thickening wedge, indirectly driving thrusting in the foreland. In the Valhalla dome, several kilometres of complexly and ductilely deformed metasedimentary rocks overlying the Gwillim Creek shear zone may be consistent with this model. However, the Gwillim Creek shear zone itself has a history of high strain lasting 20-30 million years. During that time it carried a coherent, internally deforming thrust sheet. The interpreted crosssection (Fig. 2) implies that the Lewis and related thrusts formed at the leading edge of a major composite crystalline thrust sheet of the kind proposed by Hatcher & Hooper (1992), with a width, measured along-strike, of some 600 km and a length, measured across-strike, of at least 400 km. The Gwillim Creek shear zone, a ductile detachment flow zone that formed well below the brittle-ductile transition, gradually passed eastward into the basal d~collement of the Rocky Mountain thrust belt. Spear (2004), quoting Hollister & Crawford (1986), suggested that melting and associated reactions would facilitate major deformation in the shear zone. The Gwillim Creek shear zone is marked by abundant strained pegmatite; abundant pegmatite fluid would also have contributed to weakening of the zone. The weakening bears some analogy to the melt weakening of a middle-crustal layer that leads to channel flow in models of Beaumont et al. (2006). However, the weakened zone in the SE Cordillera carried the suprastructure along with it. Within the sheet, from west to east, there was every transition from weakening by melt reactions and the presence of melt in the flow zone, to weakening of ductile zones by metamorphic reactions with aqueous fluids, and ultimately to weakening of brittle

thrust surfaces in the external thrust belt by pore water acting as a chemically inert fluid exerting pressure that reduced effective normal stress across thrust planes. The different mechanisms overlapped in space and time, and all contributed to facilitate the eastward motion of the great sheet that had the Lewis thrust at its leading edge.

Origin of the Valhalla structural culmination Arching in the Valhalla complex is a young feature that post-dated deactivation of the Gwillim Creek shear zone, as well as intrusion by the Palaeogene Ladybird granite suite, and formation of lineations in the Eocene Valkyr-Slocan Lake extensional shear zone system (Carr et al. 1987), but predated cooling of the complex at c. 52Ma (Heaman & Parrish 1991). Figure 2 depicts a basement slice with basal detachment beneath the Valhalla complex. Ramping and eastward advance of the basement slice at c. 55-52 Ma would account for arching of the Valkyr shear zone, the Valhalla complex and the Gwillim Creek shear zone. If the arching of the complex is related to duplexing of basement slices beneath it, then it is likely that variations in the basement slice and underlying ramp may explain the geometry of subculminations within the complex. This model is consistent with previous interpretation of Lithoprobe geophysical data (Eaton & Cook 1990), and timing of late thrusting in the external zone of the Foreland belt (e.g. Rocky Mountains and Foothills). Doming mechanisms related to channel and/or detachment flow as suggested by models of Beaumont et al. (2006) are ruled out since the doming is entirely younger than flow of the infrastructure within the complex.

Northward continuation of the Gwillim Creek shear zone Given the high strain and possible large displacement of the Gwillim Creek shear zone in the Valhalla complex, it is to be expected that the shear zone and the thrust sheet it carried extend north of the Valhalla complex. The Monashee basement complex, to the NW (Fig. 1), projects beneath the Valhalla complex, on the basis of both structural mapping between the two areas (Cart 1991, 1995) and subsurface geometry revealed by Lithoprobe seismic reflection profiles (Cook et al. 1992). These same geological and geophysical data also indicate a pattern of structural depressions and culminations that bring the Gwillim Creek shear zone to the present erosional surface on the south and SW margins of the Monashee complex (Fig. 1).

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA Northward from there, it becomes involved in younger folding and flow, and loses its definition. The highly strained and metamorphosed rocks lying on the west flank of Thor Odin dome (TO, Fig. 1), in the southern Monashee complex, must include the Gwillim Creek shear zone and the thrust sheet that it carried, albeit transposed and overprinted by younger deformation. The upper, low-grade, supracrustal part of the thrust sheet is preserved on the west, down-thrown side of the Eocene extensional Okanagan-Eagle River normal fault system (OER, Fig. 1), that is underlain by a zone of ductile top-down-to-the-west shear (Johnson 2006). The gneissic sheet carried by the Gwillim Creek shear zone and by deeper zones of flow and shear in the antiformal Monashee complex (Johnson & Brown 1996; Williams & Jiang 2005; Brown & Gibson 2006) dips eastward on the east side of the complex. It must continue eastward and lie buried under low-grade rocks in the hanging wall of the top-to-the-east Columbia River normal fault (Fig. 1), in a manner that is similar to the way the Gwillim Creek shear zone passes under the Kootenay Arc on the east side of the Valhalla complex (Fig. 2). The low-grade suprastructure that overlies the gneiss sheet has an older deformation history (Crowley & Brown 1994) and the gneiss in question must continue eastward in the subsurface. In a west to east transect near latitude 51~ the gneiss sheet must thin eastward. Sixty kilometres east of the Monashee complex there is only some 5 km of room for it between autochthonous basement below, and Neoproterozoic rocks at low metamorphic grade above (Colpron et al. 1996). At the west margin of the Rocky Mountains, conservative down-plunge projection and fault restorations (Lickorish 1993; Kubli & Simony 1994) leave no room for the zone of gneisses between basement and overlying low-grade rocks. The 'tip line' of the gneissic infrastructural flow zone must, therefore, lie at a depth of 12 to 15 km, about 70 km east of the Monashee complex and is most probably linked to the 'tip line' of the Gwillim Creek shear zone in Figure 2. Such a 'tip line' is illustrated in Figure 3a and the map projection of that subsurface 'tip line' is illustrated in Figure 1. It represents schematically the zone where the infrastructure thins and its ductile eastward flow is transferred to eastward motion on the thrusts in the external zone, that define the front of the great composite crystalline sheet. The infrastructural gneissic sheet is at its thickest (15-20km; Johnston et al. 2000) and perhaps hottest (Norlander et al. 2002) in the Thor Odin dome (Hinchey 2005; Hinchey et al. 2006). It is of interest that it is NE of this thick, hot, central zone that the Late Cretaceous-Palaeocene thrusts

579

of the Rocky Mountain Front Ranges have their greatest displacement. That displacement decreases northwestward, along-strike, and is significantly reduced by latitude 53~ (Price & Mountjoy 1970; Mountjoy 1980; Fermor 1999). In the corresponding hinterland to the west, Late CretaceousPalaeocene pegmatite and granitic leucosome become less abundant northwestward, such that in the NW corner of Figure 1, north of latitude 52.5 ~ N, only Middle Jurassic and Early Cretaceous ages have been reported to date (Parrish 1995; Digel et al. 1998). Thus, large Late CretaceousPalaeocene thrust displacement in the foreland is linked to Late Cretaceous-Palaeocene metamorphism, pegmatite emplacement and melting in a ductile zone in the corresponding hinterland. Weakening of the base of the composite crystalline thrust sheet by high P-T conditions, presence of pegmatitic fluids and reactions associated with melting (Hollister & Crawford 1986) led to the development of a detachment flow zone. Transport of the rocks above the flow zone, towards the foreland, approximately balanced shortening of the same age on the foreland thrust system. Such a temporal linkage of metamorphism and melting in the hinterland and internal zones with thrust displacement in the external zone, supports the contention that different mechanisms responsible for weakening the base of the great crystalline thrust sheet operated in overlapping space and time. It also requires the existence of the kind of subsurface 'tip line' for the infrastructural flow zone illustrated in Figures 1-3, with transfer of motion as proposed above. In Figure 3a, the right face of the block diagram, illustrating the gneissic infrastructure, geometrically resembles models where the channel is shown to 'tunnel' at its leading edge (Beaumont et al. 2006). The resemblance is only geometric, not kinematic, as the detachment flow zone was bolted to its suprastructure and carried it forward.

Reconciliation of crystalline thrust sheet model for the Valhalla complex (latitude 49-50.5~ with proposed channel flow models to the north (latitude 50.5-52~ Channel flow models have been proposed for some of the rocks of the infrastructural flow zone that lie to the north of the Valhalla complex. In a transect through the Frenchman Cap dome, the northern part the Monashee complex, Brown & Gibson (2006) interpreted the top-down-to-the west ductile shear zone, below the gently west-dipping Okanagan-Eagle River fault zone, as the upper part of a mid-Cretaceous-Palaeocene channel. They interpreted the base of the channel to be

580

S.D. CARR & P. S. SIMONY

arched over the dome, and to lie structurally above North American basement in the core of the dome. East of the Columbia River fault, in the hanging wall of a top-down-to-the-east normal fault that bounds the dome, the channel is buried under low-grade rocks (Colpron et al. 1996; Johnson & Brown 1996). In their interpretation, it re-emerges further east in the Big Bend region (BB, Fig. 1), west of the Rocky Mountains, in the footwall of a top-down-to-the-west ductile shear zone (Brown & Gibson 2006). Highly strained and metamorphosed rocks west of the Thor Odin dome and highly transposed rocks within the dome, including North American basement, are interpreted in terms of channel flow by Williams & Jiang (2005), Kuiper et al. (2006) and Williams et al. (2006). Within the proposed channel, a c. 10 km thick lower zone with top-to-the-east transposition foliation and eastvergent drag folds was overprinted at high levels by west-verging folds. The west-verging structures are interpreted to mark the upper part of the channel, and indicate the development of channel flow, following a period of east-directed detachment flow (Williams & Jiang 2005; Kuiper et al. 2006). The infrastructural flow zone does not extrude; rather, eastward of the Monashee complex it has to continue beneath low-grade suprastructure rocks that have an older deformation history (Crowley & Brown 1994). For purposes of discussion, the name Monashee channel is used to refer to proposed channel flow involving the Monashee complex and surrounding rocks (Williams & Jiang 2005; Brown & Gibson 2006; Kuiper et aL 2006), and terminology and concepts of Williams et al. (2006) are used. The width of the proposed Monashee channel, along-strike, is constrained to be about 200-300 km by boundaries on both its northern and southern sides. These may have been lateral transition zones across which the effectiveness of the upper detachment decreased and channel flow passed laterally into detachment flow, in the manner schematically illustrated in Figure 3a. In the case of the northern boundary, detachment flow, as well as channel flow, decreased northward between latitudes 52~ and 52.3~ in the NW corner of Figure 1. North of that boundary zone, a structural and metamorphic edifice of Middle Jurassic and Early Cretaceous age (Parrish 1995, and references therein; Digel et al. 1998; Reid 2003) preserves rocks formed at more than 25 km below the surface (Digel et aL 1998; Ghent & Simony 2005). The southern margin of the channel must lie north of the Valhalla complex and SW of the Thor Odin dome. High-grade metamorphic rocks exposed between the Valhalla complex and the Intermontane belt to the west, all share the same suprastructure with

the Valhalla complex (e.g. Upper Palaeozoic, Triassic and Jurassic rocks of the Quesnel terrane, Fig. 1). Therefore, models of east-directed channel flow for that region are severely limited by the suprastructure-infrastructure coherence documented for the Valhalla complex. The proposed Monashee channel can be integrated into the model that we propose here, of a very large composite crystalline thrust sheet that is c. 600 km wide along-strike. The central 200 km wide segment within the detachment flow zone became thick and hot enough for the initiation of channel flow, such that detachment flow passed laterally into channel flow as illustrated in Figure 3a. There are relationships in the Big Bend region and in the vicinity of the Monashee complex that suggest that if channel flow took place in the Monashee channel, the nascent stage of that flow may be preserved. In the northern portions of the Big Bend region (Fig. 1), Neoproterozoic formations outline a pre-Middle Jurassic, westverging nappe structure. It was refolded and overprinted by younger metamorphism (Raeside & Simony 1983) but the preservation of its structure and coherence of stratigraphy inside the channel can only be understood if flow in the channel had not evolved too far. Devono-Mississippian orthogneiss sheets occupy similar stratigraphic-tectonic positions at the base of the suprastructure to the east and west of the Monashee complex (Johnson & Brown 1996). A similar orthogneiss sheet, in a comparable metasedimentary succession, lies in the upper part of the Monashee channel, just west of the Monashee complex (Johnston et aL 2000; Kuiper et al. 2006). This suggests that the transition from the lower suprastructure into the channel is at least partly preserved, and that the evolution of channel flow out of detachment flow (Williams & Jiang 2005) had not progressed too far. The proposed Monashee channel would have evolved within the infrastructural detachment flow zone and it would have shared the subsurface tip line with the detachment flow zone. The evolving channel would therefore have been 'tunnelling' (Beaumont et al. 2006), as illustrated in the left part of Figure 3a, in contrast to the detachment flow seen on the right face of the block. The suprastructure of the Monashee channel was continuous with the Neoproterozoic and Palaeozoic strata of the Rocky Mountain thrust belt and substantial flow in the channel would have required the formation of structures to balance the excess volume brought to the tip line (Fig. 3b; Beaumont et al. 2006). The fact that such structures are not obvious is consistent with the idea, proposed earlier, that channel flow was arrested at the nascent stage. The local rise of the Monashee channel to the present

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA surface, NE of the Frenchman Cap dome in the Big Bend area (BB, Fig. 1) proposed by Brown & Gibson (2006), does not necessarily involve exhumation of the 'tip line' and does not change this general conclusion. Even if channel flow did take place in the southeastern Canadian Cordillera, a fundamental difference remains between the Cordillera and the Himalaya, aside from the scale of the channel flow; no good analogue to the Himalayan Main Central thrust evolved in the Cordillera to direct the channel structurally above the Foreland thrust belt during flow. If erosion exerted a control on Cordilleran detachment or channel flow, it did so indirectly through erosion at the CretaceousPalaeocene Rocky Mountain front, not through erosion of extruding channel material.

Conclusions In the hinterland of the southeastern Canadian Cordillera, thick zones of gently dipping, gneissic, migmatitic and ductilely sheared rocks were hot and ductile in Cretaceous and Palaeocene times. They formed a widespread infrastructural sheet that carried a suprastructure already deformed, metamorphosed and intruded in the Middle Jurassic. The suprastructure was carried eastward behind the evolving external part of the thrust belt, on thick ductile shear zones or detachment flow zones. The question is whether all or part of the infrastructure became hot and weak enough to flow ahead of its suprastructure, and whether or not channel flow is a major mechanism in Cordilleran tectonism. In the Valhalla complex, a coherent stratigraphic succession and an edifice of interlinked plutons bolt the suprastructure to the infrastructure to constitute a c. 30 km thick thrust sheet. It contains a gradient of downward-increasing ductile strain and metamorphic grade to the base of the c. 7 km thick Gwillim Creek shear zone. Channel flow and ductile extrusion are models that are inconsistent with the coherence of this thrust sheet and its strain gradient. Within the thrust sheet, the Eocene Valkyr-Slocan Lake extensional shear zone bounds the upper margin of the Valhalla complex (infrastructure); however, its magnitude of displacement is not large enough to disrupt the coherence of the sheet, and motion is entirely younger than, and has the same shear sense as the motion on, the underlying Gwillim Creek shear zone. Therefore, the two shear zones cannot constitute the upper and lower bounding shear zones of a channel, and we conclude that the Gwillim Creek shear zone forms the base of a c. 30 km thick thrust sheet that translated both infrastructure and suprastructure towards the foreland.

581

Geochronology linked to petrological and microstructural studies demonstrate that the Gwillim Creek shear zone was a zone of ductile shear, pegmatite emplacement and melting from c. 90 to 60 Ma. A frontal ramp, and a related north-dipping lateral ramp, existed beneath the Gwillim Creek shear zone and the Valhalla complex in the Late Cretaceous, as documented by geological relationships in the hanging wall sheet. Rapid quenching at the base of the thrust sheet carried by the Gwillim Creek shear zone suggests refrigeration by its emplacement up a 1 0 - 1 2 k i n high westdipping ramp onto a cold footwall. Flow of hot ductile rocks over the ramp and their refrigeration on the upper flat require that the Gwillim Creek shear zone was a major ductile shear zone, or detachment flow zone, with an estimated thrustsense displacement of 80-100 kin. It can be directly linked to a system of thrusts in the foreland related to the Lewis thrust. Such a linkage implies the existence of a composite crystalline thrust sheet some 600 km wide, parallel to strike, more than 400 km long, in the transport direction, and some 30 km thick in its western part. The sheet was moving eastward from c. 90 to 60 Ma. Doming of the Valhalla complex is entirely younger than deformation and metamorphism in the infrastructure, as well as c. 59-55 Ma motion on the Valkyr extensional shear zone, and may be related to Eocene thrusting beneath the complex during the last stage of shortening in the Foothills of the Rocky Mountain thrust belt. Channel flow has been proposed for the infrastructural zone in the vicinity of the Monashee complex, to the north of the Valhalla complex. This channel is constrained by regional geological relationships to have a width of only 200-250 km parallel to strike. We suggest a three-dimensional model that could reconcile the proposed channel with the composite crystalline thrust sheet model. The central 200 km wide segment in the infrastructure of the crystalline sheet could have become thick, hot and weak enough to begin flowing ahead of its suprastructure. This was facilitated by activation of an upper detachment zone and of lateral transition zones. The channel did not extrude but 'tunnelled' to a tip line where its eastward flow was transferred to eastward ductile thrusting. The proposed channel is of limited width, may not have flowed very far, and would have evolved within a restricted part of the infrastructure of a great crystalline thrust sheet where detachment flow and ductile thrusting, linked to thrusting in the external zone, were the dominant mechanisms. We therefore conclude that channel flow was not the dominant orogenic mechanism in the SE Canadian Cordillera.

582

S.D. CARR & P. S. SIMONY

We gratefully acknowledge research funding from NSERC and Lithoprobe, fruitful discussions with our Cordilleran colleagues, in particular R. Brown, F. Cook, J. Crowley, A. Hinchey, R. Price, F. Spear and P. Williams, and constructive reviews of the manuscript from D. Gibson, R. Jamieson and R. Law. Publication number 42 of the Ottawa-Carleton Geoscience Centre, Isotope Geochemistry and Geochronology Research Facility.

Appendix U - P b geochronology o f the China Creek Pegmatite (ST-92-52) The pegmatite geochronology sample locality is west of Highway 22 between Trail and Castlegar (Fig. 6), 1 km north of the village of China Creek on the south side of the creek. The China Creek Pegmatite was deformed in the Gwillim Creek shear zone in the core of the China Creek culmination. The leucocratic muscovite biotite pegmatite is strongly foliated with an east-west lineation. Although deformed, and generally concordant with highly deformed banded gneisses and enclaves of Middle Jurassic granodiorite-tonalite, the granitic pegmatite cuts a package of layered biotitic gneiss, and is considered to be pre or syntectonic with respect to the Gwillim Creek shear zone. Although the timing of shear zone initiation is uncertain, significant strain post-dated emplacement of the pegmatite.

0.030

U - P b geochronology was carried out at Carleton University following isotope dilution procedures outiined by Parrish et al. (1987b) using microcapsules for mineral dissolution (Panfish 1987) and a mixed 233U-235U-2~ tracer (Parrish & Krogh 1987). Multicollector mass spectrometry was carried out on a Finnigan MAT 261 mass spectrometer (Roddick et al. 1987); error estimation used numerical error propagation (Roddick 1987); decay constants used are those recommended by Steiger & J~iger (1977); and discordia lines through analyses were calculated with the use of a modified York (1969) regression (Panfish et aL 1987b). Data are presented in Table 1 and the concordia diagram in Figure 9. Zircons are colourless, euhedral and sharply faceted, and many have large xenocrystic cores. Of the zircons without visible cores, the clearest, least magnetic zircons with the fewest inclusions and cracks were abraded prior to analysis. Of the four fractions analysed, fractions D and E have inherited Pb as indicated by large discordance of 83% and 91%, respectively, and their Proterozoic 2~176 ages (Table Al, Fig. AI). Although fractions B and C plot near the concordia curve around 79 Ma, they are 19% and 13% discordant, respectively, with 2~176 ages of c. 90 Ma. Fraction B is a multigrain fraction and fraction C is a single-grain fraction, and the crystals are all of similar size. Because the 2~176 as well as the 2~ and 2~ ages of both fractions are the same within error, the discordance was likely caused by a recent Pb-loss event, and the 90.3 _+ 0.6 Ma 2~176 date of fraction C is taken as the crystallization age of the pegmatite.

China Creek pegmatite $7"-92-52 1U

j

160

0.024

|

/ 140

o.o18

O

oo Y ' /

-~9

//loo

/ /,C J

y

80

.o o.o .

j

o.ooo /

o.o,.

, oo0.0

o.o15

//oo '

006

'

0 12"Wt 207

,/ "" 0.06

'

' 0.09

235 Pb/

U

Fig. A1. Concordia diagram of U - P b geochronology data for the China Creek pegmatite; inset (bottom right) shows enlargement of data near concordia curve. Zircons are labelled with letters A - E and monazites with numerals 1, 2 and 4 - 6 .

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA

d o

, ,+l

+1 +1 TI ~

~ ~ Z Z Z Z Z

+1+1+1+1+1+1+1+1+1

.

.

.

.

.

o

o

o

o

r oo

Z

cSe!.o~F~cSc5 +I +I +l +l +l +l +l +l +l

~ ~

~ ~

+, +, +, +, +, +, +, +, +,

~

,

~

~

o

N z ~ ~c~ (-q

oo p~ oo

oo ~

o~

oo oo

oo

oo

p--

oddddddd

("4

+I +I +i +i +i +I +I +I +I

"~

I

(D

C5 C5 c5 C5 c5 C5 C5 C5 C5

~

~

~

~

.

~

.,..~

,.-.,

.,.~

.-~ v~

oo

0

-~- c q

cq

cq

~

~. ~

~

~

~

9

oo

cq ,.-.,

o

,.,_v

I

m

< ~-~ cq ~

.~ .~

~

.~

o

583

584

S.D. CARR & P. S. SIMONY

Two multigrain and three single-grain fractions of pale yellow, irregular-shaped anhedral monazite crystals were analysed (Table A1, Fig. 1). The analyses are reversely discordant, plotting above the concordia curve (i.e. 2~ dates > 2~ dates). This is typical of 'young' monazites from the metamorphic terranes in British Columbia (Parrish 1990, 1995, and references therein) and is interpreted as being due to unsupported 2~ resulting from incorporation of excess 23~ upon crystallization (Scharer 1984; Parrish 1990); therefore, 2~ dates are reported here. Corrections of the 2~ dates for excess 2~ were not made using the method outlined by Sch~er (1984) because, as pointed out by Parrish (1990), there is little justification for the use of this correction for monazite from metamorphic terranes since the T h - U ratios of the fluids in which the mineral grew are unknown. The 2~ monazite dates range from 76.9 ___0.2 Ma to 86.5 ___ 1.2 Ma (Table A1, Fig. 1). The ages of three of the five fractions overlap within error; however, the spread in data is considered to be real. Diffusive Pb loss is ruled out because many studies have shown that monazite is highly resistant to this thermally induced process (Smith & Gilletti 1997; Foster & Parrish 2003, and references therein). The spread in dates may reflect either: (i) progressive growth of monazite crystals in the Late Cretaceous, in which case the 77-87 Ma 2~ dates reflect the timing of metamorphism; or (ii) two or more episodes of Late Cretaceous or younger monazite growth or recrystallization resulting in more than one intracrystalline age domain, as commonly occurs in monazite (Foster & Parrish 2003, and references therein; Gibson et al. 2004). These models could possibly be distinguished by dissolving parts of crystals for dissolution and anaylses or by using in situ analytical techniques in conjuction with grain imaging and chemical mapping (Gibson et al. 2004); however, both scenarios point to monazite growth during a Late Cretaceous metamorphic event following c. 90 Ma intrusion of the pegmatite. References

ARCHIBALD, D. A., GLOVER, J. K., PRICE, R. A., FARRAR, E. & CARMICHAEL,D. M. 1983. Geochronology and tectonic implications of magmatism and metamorphism, southern Kootenay Arc and neighboring regions, southeastern British Columbia. Part I: Jurassic to Mid-Cretaceous. Canadian Journal of Earth Sciences, 20, 18911913. BALLY, A. W., GORDY, P. L. & STEWART, G. A. 1966. Structure, seismic data and orogenic evolution of the southern Canadian Rockies. Canadian Society of Petroleum Geologists, Bulletin, 14, 3 3 7 - 381. BEAUMONT, C., NGUYEN, M. H., JAMIESON, R. J. & ELLIS, S. 2006. Crustal flow modes in large hot orogens. In: LAW, R. D., SEARLE, M. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 91 - 145. BROWN, R. L. 2004. Thrust belt accretion and hinterland underplating of orogenic w e d g e s - A n example from the Canadian Cordillera. In:

MCCLAY, K. R. (ed.) Thrust Tectonics and Hydrocarbon Systems. Association of American Petroleum Geologists, Memoir, 82, 51-64. BROWN, R. L. & GIBSON, D. H. 2006. An argument for channel flow in the southern Canadian Cordillera and comparison with Himalayan tectonics. In: LAW, R. D., SEARLE, M. GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 543-559. CAMPBELL, R. B. 1973. Structural cross-section and tectonic model of the southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 10, 1607-1620. CARR, S. D. 1991. Three crustal zones in the ThorOdin-Pinnacles area, southern Omineca Belt, British Columbia. Canadian Journal of Earth Sciences, 28, 2003-2023. CARR, S. D. 1995. The southern Omineca Belt, British Columbia: new perspectives from the LITHOPROBE Geoscience program. Canadian Journal of Earth Sciences, 32, 1720-1739. CARR, S. D., PARRISH, R. & BROWN, R. L. 1987. Eocene structural development of the Valhalla complex, southeastern British Columbia. Tectonics, 6, 175-196. COLPRON, M., PRICE, R. A., ARCHIBALD, D. A. & CARMICHAEL, D. M. 1996. Middle Jurassic exhumation along the western flank of the Selkirk fan structure: thermobarometric and thermochronometric constraints from the Illecillewaet synclinorium, southeastern British Columbia. Geological Society of America Bulletin, 108, 1372-1392. COOK, F. & VAN DER VELDEN, A. 1995. Three-dimensional crustal structure of the Purcell Anticlinorium in the Cordillera of southwestern Canada. Geological Society of America Bulletin, 107, 642-664. COOK, F., GREEN, A., SIMONY, P. ETAL. 1988. Lithoprobe seismic reflection structure of the southeastern Canadian Cordillera: Initial results. Tectonics, 7, 157-180. COOK, F., VARSEK, J., CLOWES, R. ET AL. 1992. Lithoprobe crustal reflection cross section of the southern Canadian Cordillera, 1, Foreland thrust and fold belt to Fraser River fault. Tectonics, 11, 12-35. CROWLEY, J. L. & BROWN, R. L. 1994. Tectonic links between the Clachnacudainn Terrane and Selkirk Allochthon, southern Omineca Belt, Canadian Cordillera. Tectonics, 13, 1035-1051. CROWLEY, J. L., GHENT, E. D., CARR, S. D., SIMONY, P. S. & HAMILTON, M. A. 2000. Multiple thermotectonic events in a continuous metamorphic sequence, Mica Creek area, southeastern Canadian Cordillera. Geological Materials Research, 2, 1-45.

DIGEL, S. G., GHENT, E. D., CARR, S. D. & SIMONY, P. S. 1998. Early Cretaceous kyanite-sillimanite metamorphism and Paleocene sillimanite overprint near Mount Cheadle, southeastern British Columbia: geometry, geochronology, and metamorphic implications. Canadian Journal of Earth Sciences, 35, 1070-1078.

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA EATON, D. W. S. & COOK, F. A. 1990. Crustal structure of the Valhalla complex, British Columbia from Lithoprobe seismic reflection and potential field data. Canadian Journal of Earth Sciences, 27, 1048-1060. FERMOR, P. 1999. Aspects of the three-dimensional structure of the Alberta Foothills and Front Ranges. Geological Society of America Bulletin, 111, 317-346. FERMOR, P. B. & MOFFAT, I. W. 1992. Tectonics and structure of the western Canada foreland basin. In: MACQUEEN, A. W. & LECKIE, D. A. (eds) Foreland Basins and Fold Belts. American Association of Petroleum Geologists, Memoir, 55, 81 - 105. FOSTER, G. & PARRISH, R. R. 2003. Metamorphic monazite and the generation of P-T-T paths. In: VANCE, D., Mt3LLER, W. & VILLA, I. M. (eds) Geochronology: Linking the Isotopic Record with Petrology and Textures. Geological Society, London, Special Publications, 220, 25-47. GHENT E. D. & SIMONY, P. S. 2005. Geometry of isogradic, isothermal, and isobaric surfaces: Interpretation and application. The Canadian Mineralogist, 42, 295-310. GIBSON, H. D., CARR, S. D., HAMILTON, M. A. & BROWN, R. L. 2004. Correlating Yttrium zones and age domains in monazite with metamorphic reactions involving major pelitic phases: an integration of IDTIMS and SHRIMP geochronology with Y-Th-U X-ray mapping. Chemical Geology, 211, 237-260. GIBSON, H. D., BROWN, R. L. & CARR, S. D. 2005. UTh-Pb geochronologic constraints on the structural evolution of the SeLkirk fan, northern Selkirk Mountains, southeastem British Columbia. Journal of Structural Geology, 27, 1899-1924. GRUJIC, D, HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel: Insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. HALWAS, D. & SIMONY, P. S. 1986. The Castlegar Gneiss Complex, Southern British Columbia. Geological Survey of Canada, Paper 86-1A, 583-587. HATCHER, R. D. & HOOPER, R. J. 1992. Evolution of crystalline thrust sheets in the internal parts of mountain chains. In: MCCLAY, K. R. (ed.) Thrust Tectonics. Chapman and Hall, London, 217-233. HEAMAN, L. & PARRISH, R. R. 1991. U.-Pb geochronology of accessory minerals. Short course on Applications of Radiogenic Isotopic Systems to Problems in Geology. Mineral Association of Canada, 59-102. HINCHEY, A. M. 2005. Thor-Odin dome: constraints on Paleocene-Eocene anatexis and deformation, leucogranite generation and the tectonic evolution of the southern Omineca belt, Canadian Cordillera. PhD thesis, Carleton University, Ottawa, Ontario. HINCHEY, A. M., CARR, S. D., McNEILL, P. D. & RAYNER, N. 2006. Paleocene-Eocene high-grade metamorphism, anatexis and deformation in Thor-Odin dome, Monashee complex, southeastern British Columbia. Canadian Journal of Earth Sciences, in press.

585

HOLLISTER, L. S. & CRAWFORD, M. L. 1986. Meltenhanced deformation-a major tectonic process. Geology, 14, 558-561. JOHNSON, B. J. 1994. Structure and early Tertiary extensional tectonics of the Okanagan Valley fault system, southern Omineca Belt, British Columbia. PhD thesis, Carleton University, Ottawa, Ontario. JOHNSON, B. J. 2006. Extensional shear zones, granitic melts and linkage of overstepping normal faults bounding the Shuswap metamorphic core complex, British Columbia. Geological Society of America Bulletin, 118, 366-382. JOHNSON, B. J. & BROWN, R. L. 1996. Crustal structure and early Tertiary extensional tectonics of the Omineca Belt at 51~ latitude, southern Canadian Cordillera. Canadian Journal of Earth Sciences, 33, 1596-1611. JOHNSTON, D. H., WILLIAMS, P. F., BROWN, R. L., CROWLEY, J. L. & CARR, S. D. 2000. Northeastward extrusion and extensional exhumation of crystalline rocks of the Monashee complex, southeastern Canadian Cordillera. Journal of Structural Geology, 22, 603-625. JONES, T. P. & NtJR, A. 1982. Seismic velocity and anisotropy in mylonites and the reflectivity of deep crustal fault zones. Geology, 10, 260-263. KUBLI, T. E. & SIMONY, P. S. 1994. The Dogtooth Duplex, a model for the structural development of the northern Purcell Mountains. Canadian Journal of Earth Sciences, 31, 1672-1686. KUIPER, Y. D., WILLIAMS, P. F. & KRUSE, S. 2006. Possibility of channel flow in the southern Canadian Cordillera; a new approach to explain existing data. In: LAW, R. D., SEARLE, M., & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 587-609. LARSON, K. P., PRICE, R. A. & ARCHIBALD, D. A. 2004. The southern termination of the Western Ranges and Main Ranges of the southern Canadian Rocky Mountains: tectonic and geochronologic implications of new stratigraphic and structural data and detailed mapping. Geological Society of America, Abstracts with Programs, 36, 208. LICKORISH, W. n. 1993. Structural evolution of the Porcupine Creek Anticlinorium western Main Ranges, Rocky Mountains, British Columbia. Journal of Structural Geology, 15, 477-489. LITTLE, H. W. 1982. Geology of the Rossland-Trail Map Area, British Columbia. Geological Survey of Canada, Paper 79-26. MOUNTJOY, E. W. 1980. Mount Robson, Alberta British Columbia (83 E). Geological Survey of Canada Map 14998A, 1:250 000 scale. MURPHY, D. C. 1987. Suprastructure/infrastructure transition, east-centrai Cariboo Mountains, British Columbia: geometry, kinematics, and tectonic implications. Journal of Structural Geology, 9, 13 -29. NORLANDER, B. H., WHITNEY, D. L., TEYSSIER, C. & VANDERHAEGHE, C. 2002. Partial melting and decompression of the Thor-Odin Dome, Shuswap

586

S.D. CARR & P. S. SIMONY

metamorphic core complex, Canadian Cordillera. Lithos, 61, 163-125. OSADETZ, K. G., KOHN, B. P., FEINSTEIN, S. & PRICE, R. A. 2004. Foreland belt thermal history using apatite fission-track thermochronology: Implications for Lewis thrust and flathead fault in the southern Canadian Cordilleran petroleum province. In: SWENNEN, R., ROURE, F. & GRANATH, J. W. (eds). Deformation, Fluid Flow and Reservoir Appraisal in Foreland Fold and Thrust Belts. The American Association of Petroleum Geologists, Hedberg Series, 1, 21-48. PARRISH, R. R. 1987. An improved micro-capsule for zircon dissolution in U.-Pb geochronology. Chemical Geology (Isotope Geoscience Section), 66, 99 - 102. PARRISH, R. R. 1990. U-Pb dating of monazite and its applications to geological problems. Canadian Journal of Earth Sciences, 27, 1431 - 1450. PARRIS~, R. R. 1992. U-Pb ages of Jurassic-Eocene plutonic rocks in the vicinity of Valhalla complex, southeastern British Columbia. In: Radiogenic Age and Isotope Studies, Report 5. Geological Survey of Canada, Paper 91-2, 115-134. PARRISH, R. R. 1995. Thermal evolution of the southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 32, 1618-1642. PARRISH, R. R. & KROGH, T. E. 1987. Synthesis and purification of 2~ for U - P b geochronology. Chemical Geology (Isotope Geoscience Section), 66, 103-110. PARRISH, R. R., CARR, S. D. & BROWN, R. L. 1985. Valhalla gneiss complex, southeast British Columbia: 1984 fieldwork. Geological Survey of Canada, Paper 85-1A, 81-87. PARRISH, R. R., PRICE, R. A. & CARR, S. D. 1987a. Late Cretaceous-Paleogene metamorphism in the core zone of the Cordilleran orogen: Valhalla complex of southeastern British Columbia. Geological Society of America, Abstracts with Programs, 19, 800. PARRISH, R. R., RODD1CK,J. C., LOVERIDGE,W. D. & SULLrVAN, R. W. 1987b. Uranium-lead analytical techniques at the Geochronology Laboratory. In: Radiogenic Age and Isotopic Studies, Report 1. Geological Survey of Canada, Paper 87-2, 3-7. PARRISH, R. R., CARR, S. D. & PARKINSON, D. 1988. Eocene extensional tectonics and the geochronology of the southern Omineca belt, British Columbia. Tectonics, 2, 181-212. PEEL, J. • SIMONY, P. S. 1982. Hadrynian Horsethief Creek Group/Kaza Group correlations in the southern Cariboo Mountains, British Columbia. Geological Survey of Canada, Paper 82-1A, 305-308. PRICE, R. A. 1981. The Cordilleran foreland thrust and fold belt in the southern Canadian Rocky Mountains. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Geological Society, London, Special Publications, 9, 427-448. PRICE, R. A. & CARMICHAEL,D. M. 1986. Geometric test for Late Cretaceous-Paleogene intracontinental transform faulting in the Canadian Cordillera. Geology, 14, 468-471.

PRICE, R. A. & MOUNTJOY, E. W. 1970. Geologic structure of the Canadian Rocky Mountains between Bow and Athabasca rivers' - A Progress Report. Geological Survey of Canada, Special Paper, 6, 7-25. RAESIOE, R. R. & SIMONY, P. S. 1983. Stratigraphy and deformational history of the Scrip Nappe, Monashee Mountains, British Columbia. Canadian Journal of Earth Sciences, 20, 639-650. REESOR, J. E. 1965. Structural evolution and plutonism in Valhalla gneiss complex, British Columbia. Geological Survey of Canada Bulletin, 129. REID, L. F. 2003. Stratigraphy, structure, geochronology and geochemistry of the Hobson Lake Area, (Cariboo Mountains, British Columbia) in relation to the tectonic evolution of the southern Canadian Cordillera. PhD thesis, University of Calgary, Calgary, Alberta. RODDICK, J. C. 1987. Generalized numerical error analysis with applications to geochronology and thermodynamics. Geochimica et Cosmochimica Acta, 51, 2129-2135. RODDICK, J. C., LOVERIDGE, W. D. & PARRISH, R. R. 1987. Precise U/Pb dating of zircon at the subnanogram Pb level. Chemical Geology (Isotope Geoscience Section), 66, 111-121. SCAMMELL, R. J. 1993. Mid-Cretaceous to Tertiary thermotectonic history of former rocks, southern Omineca Belt, Canadian Cordillera. PhD. Thesis, Queen's University, Kingston, Ontario. SCHARER, U. 1984. The effect of initial 23~ disequilibrium on young U.-Pb ages: the Makalu case, Himalaya. Earth and Planetary Science Letters, 67, 191-204. SCHAUBS, P. & CARR, S. 1998. Geology of metasedimentary rocks and late Cretaceous deformation history, in the Northern Valhalla complex, British Columbia. Canadian Journal of Earth Sciences, 38, 1018-1036. ScrtAtmS, P. S., CARR, S. D. & BERMAN, R. B. 2002. Structural and metamorphic constraints on ca. 70 Ma deformation of the northern Valhalla complex, British Columbia: Implications for the tectonic evolution of the southem Omineca Belt. Journal of Structural Geology, 24, 1195-1214. SEARLE, M. P. 1999. Emplacement of Himalayan leucogranites by magma injection along giant sill complexes; examples from Cho Oyu, Gyanchung Kang and Everest leucogranites (Nepal Himalaya). Journal of Asian Earth Sciences, 17, 773-783. SEARLE, M. P., SIMPSON, R. R., LAW, R. D., PARRISH, R. R. & WATERS, D. J. 2003. The structural geometry, metamorphic and magmatic evolution of the Everest Massif, High Himalayas of Nepal - south Tibet. Journal of the Geological Society, London, 160, 344-366. SIMONY, P. S. & CARR, S. D. 1997. Large lateral ramps in the Eocene Valkyr shear zone: extensional ductile faulting controlled by plutonism in southern British Columbia. Journal of Structural Geology, 19, 769-784. SMITH, H. A. & GILETTI, B. J. 1997. Pb diffusion in monazite. Geochimica et Cosmochimica Acta, 51, 1047-1055.

EVIDENCE AGAINST CHANNEL FLOW, CANADIAN CORDILLERA SPEAR, F. S. 2004. Fast cooling and exhumation of the Valhalla Metamorphic Core Complex, southeastern British Columbia. International Geology Review, 46, 193-209. SPEAR, F. S. & PARRISH, R. R. 1996. Petrology and cooling rates of the Valhalla complex, British Columbia, Canada. Journal of Petrology, 37, 733 -765. STEIGER, R. H, & J~,GER, E. 1977. Subcommision on geochronology: Conventions on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. TEYSSIER, C., FERRIC, E. C., WHITNEY, D. L., NORLANDER, B., VANDERHAEGHE, O. & PARKINSON, D. 2005. Flow of partially molten crust and origin of detachments during collapse of the Cordilleran orogen: In: BRUHN, D. & BURLINI, L. (eds) High-Strain Zones; Structure and Physical Properties. Geolo~cal Society, London, Special Publications, 245, 39-64. VAN DER PLUIJM, B., HALL, C. M., VROLIJK, P. & PEVEAR, D. 2001. Ar dating of faults in the southern Canadian Rockies; the Rundle and McConnel thrust systems. Geological Society of America, Abstracts with Programs, 33, 25. VARSEK, J. R. & COOK, F. 1994. Three-dimensional crustal structure of the eastern Cordillera,

587

Southwestern Canada and Northwestern United States. Geological Society of America Bulletin, 106, 803-823. WECMANN, C. E. 1935. Zur Deutung der Migmatite. Geologische Rundschau, 26, 305-350. WHEELER, J. O. & MCFEELY, P. 1991. Tectonic

assemblage map of the Canadian Cordillera and adjacent parts of the United States of America. Geological Survey of Canada, Map 1712A, scale 1:2,000,000. WILLIAMS, P. F. & J1ANG, D. 2005. An investigation of lower crustal deformation: Evidence for channel flow and its implications for tectonics and structural studies. Journal of Structural Geology, 27, 1486-1504. WILLIAMS, P. F. JIAN~, D. & LIN, S. 2006. Interpretation of deformation fabrics of infrastructure zone rocks in the context of channel flow and other models. In: LAW, R. D., SEARLE, M., & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 221-235. YORK, D. 1969. Least squares fitting of a straight line with correlated errors. Earth and Planetary Science Letters, 5, 320-324.

Possibility of channel flow in the southern Canadian Cordillera: a new approach to explain existing data Y V E T T E D. K U I P E R 1, P A U L F. W I L L I A M S 2 & S T E F A N K R U S E 2

1Department o f Geology and Geophysics, Boston College, Chestnut Hill, M A 02467, USA (e-mail: [email protected]) 2Department o f Geology, University o f N e w Brunswick, Fredericton, NB, E3B 5A3, Canada Abstract: Existing structural, metamorphic and geochronological data in and close to the

Shuswap Metamorphic Complex in the southern Canadian Cordillera are shown to be consistent with a channel flow model. Four general structural levels (domains) can be distinguished in the region, based on the orientation and vergence of folds. In the lowest three levels folds are mostly recumbent, whereas in the uppermost level they are upright. The lowest three levels are interpreted as a channel flow zone. NE-verging folds of the lowest level (Domain 1, e.g. the Monashee Complex) formed during top-to-the-NE detachment flow and/or in the lower part of a channel flow zone. When detachment flow changed to channel flow, the sense of shear changed in the upper part of the channel flow zone, resulting in overprinting of NE-verging folds by SW-verging folds (Domain 2, e.g. most parts of the Shuswap Metamorphic Complex to the west of the Monashee Complex). Temperature was probably increasing, weakening a progressively larger portion of the crust, and the crustal shear zone therefore widened. Thus, in the highest structural levels within the channel flow zone, SW-verging folds developed in areas where no NE-verging folds originally formed (Domain 3, e.g. the Cariboo Mountains). The channel flow model as presented here is compatible with many of the ductile structures and accommodates existing metamorphic and geochronological data in the part of the southern Canadian Cordillera described.

Rocks of the Monashee Complex (MC; an exposure of ancestral North American rocks within the Omineca Belt; Armstrong et al. 1991), and of the part of the Selkirk Allochthon (SA) that is adjacent to the MC to the west (Fig. 1), have been shown to be part of a high-grade nappe association (Williams & Jiang 2005). This association is characterized by 'high-grade metamorphism, a horizontal to shallow dipping transposition foliation (or its enveloping surface), a non-coaxial deformation history, and where divisible into sub units, sheet-like bodies (individual nappes) with boundaries that are parallel to the transposition foliation' (Williams & Jiang 2005). The structural association is interpreted as a product of detachment flow (a zone of approximately horizontal flow that transports overlying crust relative to underlying crust and/or mantle) and/or channel flow (where this zone tunnels through the crust in the direction of flow; see below). In the case of the MC and SA early detachment flow was followed by channel flow (Williams & Jiang 2005). The model for the MC and SA involves crustal thickening, by upright folding, and the development of a thick detachment zone of non-coaxial flow in the weakest rocks, between a strong upper crust and a strong lower crust and/ or mantle (cf. Fig. 2a, Domain 1). Fold vergence

is consistent with top-to-the-NE sense of shear. With progressive deformation, the zone weakened and the crust thickened, and detachment flow developed into channel flow. This resulted in a reversal of the sense of shear and fold vergence in the upper part of the zone (SA in Williams & Jiang 2005; Domain 3 in Fig. 2b, and Domain 2 in Fig. 2c). Channel flow has, for example, been modelled by Bird (1991), Mancktelow (1995), Davidson et al. (1997), Royden et al. (1997), Clark & Royden (2000), Beaumont et al. (2001, 2004), Grujic et al. (2002) and Jamieson et al. (2004). Channel flow as a result of crustal shortening has been proposed for the Himalayan orogen (Grujic et al. 1996, 2002; Beaumont et al. 2001, 2004; Jamieson et al. 2002, 2004) and for the Tibetan Plateau (Bird 1991; Westaway 1995; Royden 1996; Royden et al. 1997; Clark & Royden 2000; Shen et al. 2001) and has also been interpreted as occurring during crustal extension in the Basin and Range Province of the western USA (Bird 1991; Kruse et al. 1991; McKenzie et al. 2000). In a thickened or thickening crust, radioactive self-heating reduces the viscosity of the crust, creating a weakened zone in the middle to lower crust (Beaumont et al. 2004). Channel flow occurs in this weakened zone as a result of a horizontal gradient in lithostatic pressure due to

From: LAW, R. D., SEARLE,M. P. & GODIN,L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 589-611. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

590

Y.D. KUIPER E T A L .

Fig. 1. Tectonic assemblage map (after Journeay & Williams 1995). Area names referred to in the text are indicated in ellipses: CC, Clachnacudainn Complex; CCr, Cusson Creek; CM, Cariboo Mountains; FC, Frenchman Cap Dome; IS, Illecillewaet Synclinorium; JM, Joss Mountain; KA, Kootenay Arc; MC, Monashee Complex; MCr, Mica Creek (Antiform); MG, Malton Gneiss; OL, Okanagan Lake; P, Pinnacles Area; PA, Purcell Anticlinorium; PCA, Porcupine Creek Anticlinorium; Q, Quesnellia; SA, Selkirk Allochthon; SF, Selkirk Fan; SFA, Selkirk Fan Axis (dotted line), Scrip Nappe is in Mica Creek area (see MCr); SMC, Shuswap Metamorphic Complex; TO, Thor-Odin Dome; VC, Valhalla Complex; VR, Vidler Ridge. Fault names are indicated in rectangles: ANTF, Adams-North Thompson Fault; CRF, Columbia River Fault; GC, Gwillim Creek shear zone; OERF, Okanagan-Eagle River Fault; PT, Purcell Thrust; RMT, Rocky Mountain Trench; SLF, Slocan Lake Fault; VSZ, Valkyr Shear Zone. Location of cross-sections (Figs 4 & 5) indicated with capital letters. Inset: location of morphogeological belts of the Canadian Cordillera: 1, Insular Belt; 2, Coast Belt; 3, Intermontane Belt; 4, Omineca Belt; 5, Foreland Belt; MC, location of the Monashee Complex.

CHANNEL FLOW IN THE CANADIAN CORDILLERA (erosion)

4 4

Domain

i

...... ".. 9

2?

3

9

"

.'"'"

3

2

2

1

2?

i..'. (a) Pre-Middle Jurassic

(b)Middle Jurassic

(e) mid- to Late Cretaceous

Fig. 2. Illustration of detachment flow and channel flow and progressive fold overprint relationships in the four structural domains. The model explains fold overprinting relationships as rocks move upwards with respect to the channel flow zone, in time (see text for discussion). (a) Before the Middle Jurassic, detachment flow causes NE-verging folds in Domain 1 (and possibly Domain 2). (b) In the Middle Jurassic, channel flow is ongoing in Domains 1-3. The channel flow zone is wider than the pre-Middle Jurassic detachment zone, in order to explain SW-verging folds in Domain 3 that are not overprinting earlier NE-verging folds. (c) In the midto Late Cretaceous, rocks moved upwards with respect to the channel flow zone, due to further thickening and erosion. The same effect can be reached by the channel flow zone narrowing or moving downwards, or a combination of the two. Relative thicknesses of the domains are not to scale. Dotted, dashed and solid lines represent new, inherited and overprinted folds respectively. Folds may or may not be overprinted. Fold a in Domain 1 is overprinted twice, whereas fold b escapes overprinting, but becomes tighter. Folds of the same generation can form at different times (e.g. F1 folds al, cl and dl form at three different times) and folds of different generations can form at the same time (e.g. F1, F2 and F3 folds db c2 and a3 form in the mid- to Late Cretaceous). Folds that formed during detachment flow (Domain 1 in (a)) are referred to as detachment flow folds (DFFs) and folds that formed during channel flow (Domains 1-3 in (b) and Domains 1 and 2 in (c)) are referred to as channel flow folds (CFFs).

variations in crustal thickness (Bird 1991; Kruse e t al. 1991; McKenzie e t al. 2000; Beaumont e t aL 2004). Channel flow in the southern Omineca Belt is interpreted as having occurred in response to crustal thickening during the Cordilleran orogeny. The purpose of the study reported here is to determine whether the channel flow interpretation of ductile structures and metamorphism in the MC and SA is compatible with the geology of the

591

surrounding sections of the southern Canadian Cordillera (i.e. a much larger area than the area studied by Williams & Jiang 2005). The study is based mainly on data reported in the literature, but also on detailed mapping and analysis by the authors and co-workers of the margins of the T h o r - O d i n dome (the MC comprises two domes, the southern T h o r - O d i n dome, and the northern Frenchman Cap dome; Fig. 1), and on reconnaissance visits to other relevant areas. The area considered is indicated in Figures 1 and 3. Existing structural, metamorphic and geochronological data are shown to be compatible with a channel flow model rather than with previously proposed hypotheses, which were based on a Rocky Mountain fold and thrust belt type model. For example, previous models involved localized deformation in the Monashee Drcollement (see below) and related underlying shear zones to accommodate Cordilleran deformation in the MC (e.g. Read & Brown 1981; Brown & Journeay 1987; Brown e t al. 1992; Gibson e t al. 1999; Crowley e t al. 2001). These models are consistent with available age data, but not with all structural and general geological observations. We recognize several geometrical problems associated with application of the channel flow model, but the most important problem with the application of Rocky Mountain fold and thrust belt type models to high-grade rocks of the southern Omineca Belt, is that they do not adequately explain the observed penetrative transposition fabric within the nappes (cf. Williams & Jiang 2005). General application of Rocky Mountain fold and thrust belt type models to areas of shallowly dipping transposition has previously been criticized by Casey & Dietrich (1997) and Ramsay (1997) on the grounds that the observed geometry does not always fit the model and the shear zones commonly overprint the folds. The channel flow model presented below offers a better explanation of the penetrative ductile structures in the high-grade rocks of the area discussed (Fig. 1). Development of this model involved some reinterpretation of (especially geochronological) data and clearly further detailed work is needed to test the relative merits of the different hypotheses.

Geological background Current tectonic models for the southeastem Canadian Cordillera involve large-scale detachments near the base of the crust, beneath the Intermontane and Omineca belts (Fig. 1), that accommodated crustal shortening during Cordilleran deformation. In the southern Omineca Belt, a NE-directed detachment is drawn between the MC and the overlying rocks of the SA (e.g. Brown 1980; Read & Brown 1981; Journeay 1986; Brown e t al. 1992, and

592

Y.D. KUIPER ET AL.

Fig. 3. Domain division: Domain 1 (dark grey) is the deepest structural level currently exposed, where NE-verging channel flow folds (CFFs) and/or detachment flow folds (DFFs) are recorded; Domain 2 (intermediate grey) is an intermediate structural level, where NE-verging DFFs or CFFs are overprinted by SW-verging CFFs; Domain 3 (light grey) is a higher intermediate level, where SW-verging CFFs are present and NE-verging folds absent; Domain 4 (lightest grey) is the highest structural level where only upright folds are present. Overprinting post-metamorphic upright or steeply dipping folds are present in all domains. The boundaries between domains are mostly faults. The F3 vergence-reversal zone (dashed line), which forms the centre of the channel flow zone, is exposed SW of Thor-Odin. The boundary between Domains 1 and 2 SW of Thor-Odin is gradational. Areas not discussed are in white. Abbreviations as in Figure 1.

CHANNEL FLOW IN THE CANADIAN CORDILLERA references therein). This detachment, the Monashee D6collement, is interpreted as bounding all but the east side of the MC, where the boundary is marked by the late normal Columbia River Fault. The Monashee D6collement has been interpreted from seismic profiles as being continuous with the basal d6collement beneath the Rocky Mountain foreland fold and thrust belt (e.g. Brown et al. 1992; Cook 1995) although there is no clear, continuous trace matching the proposed discontinuity. Displacement on these large-scale detachments is interpreted as being large. For example, Brown & Cart (1990, and references therein) inferred 100 km displacement on the Monashee D6collement and 100 km on the interpreted unexposed underlying basement thrusts, to accommodate the 200 krn shortening measured in the Rocky Mountain foreland belt (Price & Mountjoy 1970; Price 1981). A problem with this interpretation is that it is made on the assumption that the thrust sheets are only locally strongly deformed and most of the deformation is concentrated along thrust sheet margins. In this interpretation, the thrust sheets start as sheets and preserve their sheet-like appearance throughout the deformation. However, in the MC and SA deformation is intense, and isoclinal folding is penetrative at all scales from the regional recumbent folds to microscopic folds. The sheetlike shapes of structural units are, at least in part, determined by deformation, not only by an original sheet morphology. Any tectonic model for this region must be able to account for this intense and complex deformation. A further problem with the fold and thrust belt model is that detailed field mapping by Reesor & Moore (1971), Johnston et al. (2000), Spark (2001) and Kruse et al. (2003, 2004) has failed to reveal any significant discontinuity that could be interpreted as a thrust separating the SA from the MC in the Thor-Odin area. Many discontinuities are seen at all scales within the MC and SA, but they do not join up and form the patterns characteristic of duplexes and other thrust structures. Many are in fact shear bands (Johnston et al. 2000). Locally these discontinuities resemble ramps, but if interpreted as such they indicate the opposite sense of shear (top-to-the-SW) to that required by the thrust model. As an alternative, an extrusion model was presented by Johnston et al. (2000) for the NW ThorOdin area. In this model, a wedge of rocks was extruded towards the NE, between two strong bounding layers, interpreted as moving towards one another during vice-like wedge extrusion. The evidence presented for this 'extrusional' flow is equally valid as supporting evidence for channel flow, because the kinematics of the two models are similar. However, the channel flow model offers a

593

better explanation for the penetrative ductile deformation in the MC and SA. The detachment/ channel flow model presented below explains Cordilleran ductile structures in the southern Omineca Belt and may require a smaller amount of crustal shortening there (cf. Williams & Jiang 2005) than the previous models based on discrete thrusting, because it does not involve long-distance transport of thrust sheets along thrust faults. Instead, thrust movement in the Rocky Mountain fold and thrust belt is accommodated by a ductile penetrative mechanism of thickening in the Omineca Belt, which does not require transport of material over long distances (see section on 'Shortening in the Omineca belt and the Rocky Mountain fold and thrust belt' below). Two other papers are presented in this volume that provide arguments for (Brown & Gibson 2006) and against (Carr & Simony 2006) channel flow in parts of the southeastern Canadian Cordillera.

The channel flow model applied to tectonic elements of the southern Omineca Belt The areas indicated in Figure 1 are discussed below in order to test the channel flow model, at a regional scale. The area west of the Okanagan-Eagle River Fault is beyond the scope of this contribution; however, available information on that area is consistent with our interpretation. The area surrounding the Valhalla Complex is also not discussed in detail, because these rocks are TriassicJurassic low-grade sedimentary and volcanic rocks, and intrusives that were probably not involved in channel flow (perhaps comparable to Domain 4; see below).

General statements about the channel f l o w model in the southern Canadian Cordillera

In this paper, NE-verging folds that develop during detachment flow (Fig. 2a, Domain 1) from preexisting folds or syndetachment flow drag folds are named detachment flow folds (DFFs), and NEor SW-verging folds that develop in the same way during channel flow (Fig. 2b & c) are named channel flow folds (CFFs), regardless of their local generation or 'phase' designations. NEverging DFFs may be overprinted by NE- or SWverging CFFs (Fig. 2), depending on their position in the channel. Structural domains where NEverging folds are overprinted by SW-verging CFFs, or where only SW-verging CFFs are present, are interpreted as having been originally at shallower structural levels (in the upper part of the channel flow zone), than domains in which

594

Y.D. KUIPER E T A L .

==

%

"6

e~

d::

~3 o

o

e~

~h t~

I

o

.=. 9

,.a O

only NE-verging folds (DFFs and CFFs) are present (Fig. 2). It is likely that the deformation zone changes width by both widening and narrowing (widening and narrowing used in the sense of Means (1995) to mean that a shear zone changes width by its boundaries migrating through the rock). The presence of SW-verging folds can be explained only by the channel being wider than the original (NE-verging) detachment flow zone. Further, the locus of weakest rocks, and therefore the position of the deformation zone, is likely to change with time as the crust thickens, melts or is exhumed. As a result, complicated overprinting and timing relationships may develop. We use the designations F1, F2 etc. for each area as used by the original authors, but note that: (1) folds with the same designation do not necessarily belong to the same stage in the deformation history, when comparing different areas (e.g. folds al and dl in Fig. 2); (2) in the infrastructure, even in one small area, because of the progressive nature of the strain, fold generations may be highly diachronous, and an F1 fold in one outcrop may be younger than an F3 fold in another outcrop (e.g. folds dl and a3 in Fig. 2). Despite the inadequacy of the generation approach in such rocks, the generation designations give some idea of relative age of the various stages of deformation. Upright or steeply inclined folds may: (1) have formed during early crustal thickening, and been preserved in the low-grade superstructure overlying the infrastructure shear zone, where they have been reorientated by the horizontal flow in the highergrade rocks; (2) represent the early stages of drag folds developing within the channel or detachment zones; or (3) be later folds overprinting all rocks during a separate deformation event. Thus, where metamorphic grade is high, isoclinal recumbent folds can be expected, plus less abundant upright and inclined folds, many of which are in various stages of being converted to isoclinal recumbent folds. Where grade is low, upright or steeply inclined folds are expected. Within the area discussed in this paper, four structural levels (which we refer to as domains) are recognized (Figs 3-6, Table 1) that are believed to have existed prior to crustal extension. The general characteristics of the domains, from lowest to highest structural level, are as follows. In Domain 1, DFFs and CFFs verge to the NE and formed during upper-amphibolite-facies metamorphism. In Domain 2, DFFs or early CFFs are NE-verging and later CFFs are SW-verging. Metamorphism was upper-amphibolite facies during both folding events. The possibility that deformation in both domains started at lower grades is not precluded, and in view of the lower-grade rocks above (Domain 3) it seems likely. This

CHANNEL FLOW IN THE CANADIAN CORDILLERA

595

Fig. 5. Cross-section of northern Domain 2 and adjacent Rocky mountains (E-F in Fig. 1). CFF is channel flow fold and DFF is detachment flow fold. Modified after Simony et al. (1980). suggestion, however, is not supported by direct evidence in the infrastructure. Domain 3 contains only SW-verging channel flow structures formed under (lower) amphibolite-facies conditions. The SWverging structures in Domain 3 formed in Middle Jurassic times, whereas the SW-verging structures in Domain 2 fonrted in mid-Cretaceous times. Domain 4 does not contain DFFs or CFFs but is characterized by upright folds that developed above the infrastructure (Figs 2 & 3). Postmetamorphic upright or steeply inclined folds occur at all levels. In general, transposition foliations are shallowly dipping. Thus, a constant vergence of structures across the shear zone is characteristic of detachment flow, while a vergence reversal, and possibly inverted metamorphic gradients, are characteristic of channel flow. High metamorphic grade and transposition are typical for both types of flow. Lithological layers are commonly, but not necessarily, dismembered (Fig. 7). Structural, metamorphic and geochronological data and interpretations for the southern Omineca Belt (Fig. 1) are described and discussed by domain below. More detailed descriptions can be found in Kuiper (2003). A summary of descriptions and interpretations is given in Table 1. We use the term crustal shear zone where we do not wish to designate the flow as detachment or channel flow. D o m a i n 1: s t r u c t u r e s a n d m e t a m o r p h i s m

Domain 1 includes the MC and the area directly south of it, the Malton Gneiss and the Valhalla Complex

(excluding the Palaeocene-Eocene granitoids). Tight to isoclinal F1 folds including regional-scale recumbent folds are common (Reesor & Moore 1971; Morrison 1982; Carr 1991; Schaubs & Carr 1998, Williams & Jiang 2005) and are interpreted as folds pre-dating the crustal shear zone and as drag folds formed within the shear zone. The inherited folds could have had any orientation initially, but were transposed during NE-directed flow. NE-verging tight to isoclinal F2 folds in Domain 1 (e.g. Morrison 1982; Carr 1991; Schaubs & Carr 1998; Johnston et al. 2000; Crowley et al. 2001; Schaubs et al. 2002; Fig. 4) and in Domain 2 (see below) may be examples of DFFs. Alternatively, they could be CFFs in the lower part of the channel flow zone. The F2 folds are interpreted as crustal-shear-zone-related drag folds, that have been transposed into the horizontal foliation with their axes rotating towards parallelism with the shear direction (cf. Williams & Jiang 2005). Open to isoclinal F3 folds, generally with SWdipping axial planes, verge to the NE in Domain 1 (e.g. Morrison 1982; Carr 1991; Schaubs & Can" 1998; Johnston et al. 2000; Crowley et al. 2001; Schaubs et al. 2002; Fig. 4). SW of the MC, F 3 folds are north-vergent to the north of Cusson Creek/Vidler Ridge (Figs 1 & 3; see Carr 1991) and south-vergent to the south of the ridge (i.e. higher up in the structural sequence, in Domain 2). The vergence reversal zone (dashed line in Fig. 3) dips to the south or SW and curves to the east around the MC, as does the transposition foliation associated with F2 and F3 folds. The location

596

Y.D. KUIPER E T A L .

Fig. 6. Summary of tectonic interpretation. (a) Location of areas discussed in the text, with respect to the channel flow zone, changing with time. Thickening of the crust is not shown. CM, Cariboo Mountains; FC, Frenchman Cap Dome; MG, Malton Gneiss; ND2, Northern Domain 2; SD2, Southern Domain 2; SF, Selkirk Fan; TO, Thor-Odin Dome; VC, Valhalla Complex. Areas indicated in small letters are in the back of each diagram, and areas indicated in large letters are in the front. (b) Horizontal projections to the back of the diagrams in (a). Geographical orientation, and orientation with respect to the channel flow direction, can be inferred from positions of areas relative to other areas in the same structural domain, and relative to the north arrow in (a). The relative geographical position of the various domains is uncertain (i.e. the Selkirk Fan may have been farther west or east relative to Thor-Odin, than is shown). Folds are only indicated on the front faces of the diagrams. Structures that formed during Middle Jurassic are overprinted by structures that formed during the mid- to Late Cretaceous (e.g. in the SD2, NE-verging folds are overprinted by SW-verging folds; see also Fig. 2). Also shown in (b) is granitoid emplacement. Middle Jurassic granitoids occur in the Selkirk Fan and Cariboo Mountains (and other areas of Domain 3), whereas mid- to Late Cretaceous granitoids were emplaced in Domain 3 and the Shuswap Metamorphic Complex (Domain 2). Melting occurred at deep structural levels (Domains 1 and 2 in the Middle Jurassic and Domain 1 in the mid- to Late Cretaceous).

of the reversal zone to the west of Cusson C r e e k / Vidler Ridge is unknown, but it presumably swings to the north, since east of Joss Mountain (cf. Fig. 1) it coincides with the transition from the MC to SA (cf. Williams & Jiang 2005); F3 structures in the MC are NE-vergent and in the SA they are SW-vergent (Fig. 4). The transition zone is cut by several north-trending faults, including the Victor Creek and Joss Pass faults east of Joss Mountain. M o v e m e n t on most of these faults is unknown, but the Victor Creek Fault has at least 1350 m displacement if a dip-slip fault, and more if, as we believe, it is a transcurrent fault (Kruse & Williams 2006). Displacement on the Victor Creek Fault m a y

be sufficient to explain the convergence of the reversal zone with the MC boundary east of Joss Mountain. In the high-grade nappe association (e.g. the MC), m e t a m o r p h i s m is high-grade (amphibolite facies and higher) and typically outlasts transposition (Williams & Jiang 2005). This is what is observed in Domain 1, except for the Valhalla Complex where F2 and F3 folds post-date peak amphibolite-facies mineral growth (Schaubs et al. 2002; cf. Cart et al. 1987; Simony & Carr 1997). M e t a m o r p h i s m is at upper amphibolite facies throughout most of the M C (e.g. Journeay 1986; Johnston et al. 2000; Norlander et al. 2002) and

CHANNEL FLOW IN THE CANADIAN CORDILLERA ,--

=

"9= Z

~

"~

~

~ .m

=

-~-"~ ~ ' ~

597

-"

>~

,-~

~

r~,~

~

~,.~

~

~_~

C '=~" ~

~,..=

"~

~

~,~, ~

~

.~-"~

~

_

....,

0'~

>~

~.

~-'~

~

"-~"~

9

~-~

,-, ~ .,.~

9 o,

o

o -

o

,.= = o 0

s ,,--,, ,-0

598

Y . D . K U I P E R ET AL.

r~ 0 .,,,~

~

~o

~

0

0

~ ~. . ~ ' ~

~

"

~

~ o ~ ~~;

:

~

~

..~

,~

0

~I.~

9" o ,.-t,

o',

,,,,,,-i

,,a2

E~u

;>-,

o ~o, ~v 9 0

0

0 .. .,.,,,,~ o

o

,,.o

O~

.o r~

@

r.~

0 r~

CHANNEL FLOW IN THE CANADIAN CORDILLERA

"~ ~ "9

~'~

~[a~

~"

i~ ~

='~

~

~

"

_ ~ ~

o

.

0

.~,.a ~

~~~.

~

.. ~0

-~

~ ~

".~

~

~7~

~3"~"

,-~ r

~o ~'~

NN

~,

N N [-= ~ TM

< "a ,~ ~0. o

A ~

A

p-..

.,.-t

~'~

N

~

.~=

~o~

~ = ~ o . ~

~.~

~ ~ =

~,

~ =

~ . . . .

=

9~

~o~.~=.

9~

9 ~r~

~L~

~:

.'~

,~ ~

=

o

599

600

Y.D. KUIPER ET AL.

Fig. 7. Transposition foliation, in calc-silicate gneiss, NE Thor-Odin dome, looking down towards the NW. F2 folds are dismembered by local shears. Pencil for scale.

sillimanite was still stable during late crustal-scale extensional deformation (cf. Johnston et al. 2000). South of the MC, peak metamorphic amphibolitefacies conditions prevailed during F 2 and at least part of F3 (Carr 1991). Metamorphism during F2 and F3 reached kyanite-staurolite grade in the Malton Gneiss and sillimanite grade, up to 7.5 kbar and 700~ in the cover rocks above the Malton Gneiss (Morrison 1982). In Frenchman Cap (Fig. 1) the lowest structural level contains granitic intrusions of Palaeoproterozoic age which cut an earlier transposition foliation (Crowley et al. 2005; Gervais et al. 2005). Upwards through the structural pile the granitic intrusions become progressively more deformed until completely transposed by Cordilleran deformation (Crowley et al. 2005; Gervais et al. 2005). No obvious break in the fabric has been reported except for a change in lineation orientation. What looks like a single foliation is in fact two, a Palaeoproterozoic foliation underlying a Cordilleran foliation. Parallelism of the two foliations strongly suggests reactivation of the earlier foliation during Cordilleran deformation. The deepest levels of Thor-Odin are less deep than Frenchman Cap and the deformation is still Cordilleran. This Cordilleran front has not been recognized elsewhere, but it apparently represents the base of a crustal shear zone and suggests that the shear zone bottomedout in strong basement rocks rather than the mantle. However, the possibility remains that the front is simply a strong boudin that preserves the earlier history and that Cordilleran deformation exists below it, since no break is apparent in the seismic fabric (Carr 1995; Cook 1995). Domain

1: t i m i n g o f d e f o r m a t i o n

Brown et al. (1986) inferred deformation and deep burial of the MC in the Jurassic, based on regional

correlations and using balanced cross-sections. Journeay (1986) recognized two pulses of deformation and metamorphism related to Cordilleran orogenesis in Frenchman Cap. Metamorphism during F2 (M1) was at upper amphibolite-lower granulite facies in the core of Frenchman Cap and at middle amphibolite facies along the flanks of the dome. Journeay (1986) interpreted M~ as being related to Middle Jurassic regional deformation. Low-pressure metamorphism during F3 (M2) was at middle greenschist facies in the core of the complex, and at upper amphibolite-lower granulite facies at higher structural levels. M 2 and F3 were interpreted as being related to mid-Cretaceous regional deformation (Journeay 1986). The inverted isograds during M2 (Journeay 1986) are consistent with ongoing channel flow, as isotherms can become folded and locally overturned (cf. Beaumont et al. 2001, 2004; Jamieson et al. 2004) due to the non-coaxial flow. More recently, Cordilleran deformation and metamorphism in the MC have been interpreted as occurring diachronously between 78 Ma at high structural levels and 49 Ma at deep structural levels (Parrish 1995, and references therein; Crowley & Parrish 1999; Gibson et al. 1999). However, Kuiper (2003) reinterpreted these ages as representing the end of a long period of Cordilleran metamorphism and deformation (see below). Detachment or channel flow (Cordilleran deformation) and NE-verging F2 folds could have started in the Middle Jurassic, an age previously proposed by Brown et al. (1986) and Journeay (1986), consistent with ages of deformation recorded east and north of the Shuswap Metamorphic Complex (Pigage 1977; Archibald et al. 1983, 1984; Ferguson & Simony 1991; Crowley & Brown 1994; Colpron et al. 1996; Reid 2003). This would also be consistent with the timing of accretion of the Intermontane Superterrane to western North America (Monger et al. 1982; Murphy et al. 1995), which possibly provided the driving force for the deformation. Accretion of the Insular Superterrane in the mid-Cretaceous (Monger et al. 1982; Journeay 1986) may have sustained the driving force at that time. Deformation continued locally until at least c. 73 Ma, because F3 folding at Joss Mountain (Fig. 1) is bracketed between c. 73 and c. 70 Ma (Kuiper 2003; cf. Johnston et al. 2000). Middle Jurassic ages are not recorded within the MC, and mid-Cretaceous ages (i.e. older than the onset of extension at c. 75-70 Ma, cf. Kuiper 2003) are rare. The only published ages are a 78 Ma monazite (Gibson et al. 1999), which is hornslightly discordant, and c. 89 Ma 4~ blende inverse isochron ages (Spark 2001; Kuiper 2003). Foster et al. (2002) report LA-MC-ICPMS

CHANNEL FLOW IN THE CANADIAN CORDILLERA monazite ages in northern Frenchman Cap older than 100 Ma. It has been argued (Kuiper 2003) that the reason for the lack of Middle Jurassic to c. 7 5 - 7 0 Ma ages is that metamorphic conditions were such that zircon and other minerals used for U - P b dating were undergoing dissolution and there was no growth during this time interval. Growth started again as the rocks cooled, so that the U - P b ages represent the end of the high-grade metamorphism associated with channel flow. The U - P b ages get younger downwards as the system was unroofed by crustal extension and/or erosion. During high-grade metamorphism and zircon dissolution, the dissolved material migrated with the melts from the MC (and also from unexposed rocks that were at a similar structural level to the MC) and crystallized in the Middle Jurassic and mid-Cretaceous granitoids above the MC, similar to granitoids which are now present in rocks surrounding the complex (e.g. Archibald et al. 1983, 1984; Colpron et al. 1996). Because the various domains have moved laterally with respect to one another (see below), melts from the MC may have crystallized in areas that are now eroded or buried, and Middle Jurassic and mid-Cretaceous granitoids in the rocks that now surround the MC may have originated in Domain 1 levels that are currently unexposed. Dissolution in the MC was not complete, because Proterozoic zircon and monazite are present (e.g. Armstrong et al. 1991; Parkinson 1991; Crowley 1997, 1999; Crowley & Ghent 1999; Vanderhaeghe et al. 1999). This interpretation is discussed in more detail below. Movement on the Gwillim Creek shear zone (Fig. 1) in the Valhalla Complex, at some time between 110 and 51 Ma (Schaubs & Carr 1998), was synchronous with, and perhaps outlasted, F2 and F3 at higher structural levels in the complex (Schaubs & Cart 1998; Schaubs et al. 2002). Therefore Fa and F3 are at least the same age as the Gwillim Creek shear zone. Schaubs et al. (2002) interpret deformation as having taken place at c. 70 Ma, based on monazite and zircon ages of Spear & Parrish (1996). Although their interpretation is reasonable, we think that the possibility that c. 70 Ma zircon and monazite grew during extension cannot be ruled out (see Kuiper (2003) for an extensive discussion). Movement on the Gwillim Creek shear zone and F2 and F3 deformation could possibly be related to channel flow. The rocks above the Valkyr Shear Zone preserve Middle Jurassic and older structures (Carr et al. 1987). Also, they have been deformed and are intruded by Middle Jurassic, mid-Cretaceous and Middle Eocene plutons (Cart et al. 1987; Simony & Cart 1997, and references therein). They could be part of Domain 4. Domains 2 and 3 are not exposed in or around the Valhalla Complex,

601

perhaps because the Valkyr Shear Zone and the Slocan Lake Fault (Figs 1 & 3) cut through and expose deeper structural levels (Domain 1). The absence of exposure of Domains 2 and 3 may alternatively indicate that the channel flow zone in the north did not extend as far south as the latitude of the Valhalla Complex (cf. Cart & Simony 2006). Another distinctive aspect of the Valhalla Complex with respect to the MC is the presence of Late Cretaceous and Palaeocene-Eocene plutons (e.g. Simony & Carr 1997). This may indicate that, during channel flow and/or during detachment flow, the Valhalla Complex was at higher structural levels within Domain 1 than the MC. Domain

2: s t r u c t u r e s a n d m e t a m o r p h i s m

Domain 2 is an area currently exposed within the Shuswap Metamorphic Complex (Fig. 3). Northern Domain 2 is defined here as the region north of the MC and Southern Domain 2 is located west and south of it (Fig. 1). F 1 and F2 folds throughout Domain 2 are comparable in geometry to folds in Domain 1. Tight to isoclinal F1 folds including large-scale fold nappes, such as the SW-verging Scrip Nappe (Figs 1 & 5; Simony et al. 1980; Raeside & Simony 1982; Sevigny & Simony 1989; Carr 1991), may predate, but are transposed by, crustal-scale shear. NE-verging tight to isoclinal F2 folds (e.g. Fyson 1970; Simony et al. 1980; Raeside & Simony 1982; Johnson 1994; Carr 1991; Fig. 5) could be DFFs, or CFFs in the lower half of the channel flow zone. In southern Domain 2, south and west of the vergence reversal zone (Fig. 3), F3 folds verge towards the south. Elsewhere in southern Domain 2, F3 folds are WSW-verging, moderately to steeply inclined and commonly overturned (Fig. 4). F3 folds are interpreted as CFFs, formed in the upper half of the channel flow zone. They are overprinted by later discontinuous NW- and north-trending upright folds (late F3 and F4; Fyson 1970). Upperamphibolite-facies metamorphism was ongoing during NE-directed flow (Johnson 1994; Johnson & Brown 1996). Sillimanite grew before F3 and was stable during F3 (Johnson 1994). The area of northern Domain 2 is included in Domain 2, because the geometry of Ft and Fa folds is consistent with those of Domain 2. However, post-metamorphic F3 folds in northern Domain 2 are not consistent with the above description; they seem to have formed at a higher structural level than F3 folds elsewhere in Domain 3. In northern Domain 2, F3 axial surfaces are upright or steeply SW-dipping. Thus rocks in southern Domain 2 with SW-verging F3 folds have graded northward (and structurally upward) into rocks

602

Y.D. KUIPER ET AL.

with upright to steeply SW-dipping F3 axial surfaces, and lower metamorphic facies, in northern Domain 2. The Mica Creek Antiform (Fig. 5) and the Porcupine Creek Anticlinorium to the east (Fig. 1) are interpreted as an F3 fold (Simony et al. 1980). Metamorphism in northern Domain 2 varies from sillimanite-K-feldspar grade (c. 720~ and >7.34 kbar; Scammell 1993) near the MC to kyanite grade close to the Malton Gneiss (Digel et al. 1998; Crowley et al. 2000). F3 folds deform the peak metamorphic isograds, whereas the F2 folds are cross-cut by these isograds (Simony et al. 1980; Raeside & Simony 1982; Sevigny & Simony 1989; Digel et al. 1998). F2 folding therefore pre-dated the peak of metamorphism and F3 largely post-dated the peak of metamorphism. Post-metamorphic F3 structures such as the Mica Creek Antiform (and the Porcupine Creek Anticlinorium and adjacent structures in the Rocky Mountain Trench; see Simony et al. 1980; Fig. 1), with upright or steeply dipping axial surfaces, probably formed above the channel flow zone (and/or possibly after channel flow during a separate event), because there is no distinct vergence and because they formed after the peak of metamorphism (similar to F3 structures in Domain 3). They indicate crustal shortening at high structural levels and may be coeval with channel flow at lower levels. D o m a i n 2: t i m i n g o f d e f o r m a t i o n

In the Mica Creek area (Fig. 1) of northern Domain 2, metamorphic U - P b ages range between 175 and 50 Ma (Sevigny et al. 1989, 1990; Scammell 1993; Crowley et al. 2000), and have been subdivided into as many as five different deformation events: 175-160Ma, 140-120Ma, l l 0 M a , 1 0 0 - 9 0 M a and 7 5 - 5 0 Ma (Crowley et al. 2000). However Crowley et al. (2000) also present ages outside those ranges that are still between 175 and 50 Ma. Scammell (1993) interprets crustal shortening, metamorphism and anatexis as occurring from c. 135 until after c. 97 Ma. In his model, based on thermochronology of monazite, titanite, hornblende, rutile, muscovite and apatite, thrust-parallel extension occurred between 100 and 94 Ma; c. 10 km of overburden was removed and cooling occurred at around that time interval, during convergence. We suggest that metamorphism and mineral growth occurred roughly continuously between 175 and 50 Ma, with the youngest ages (Late Cretaceous-Eocene) indicating crustal extension and the end of metamorphism. If there was a gradual and continuous transition between Middle Jurassic F2 deformation and Cretaceous F3 deformation and subsequent extension, then this wide

age range is not surprising. Cooling and exhumation starting in the Cretaceous during crustal shortening (Scammell 1993) is consistent with F3 folds forming at high structural level above the channel flow zone (see above). Exhumation during crustal shortening may be a result of erosion enhanced by high relief, or of material sliding off the thickened pile (e.g. towards the present-day Rocky Mountains). The c. 77-51 Ma hornblende, biotite and muscovite K - A r ages of Sevigny et al. (1990) may be a result of crustal extension. West of the MC, 4~ hornblende ages are interpreted as Late Palaeocene to Middle Eocene hornblende cooling ages (Johnson 1994). No Jurassic or Cretaceous hornblende cooling ages exist in Domain 2 west and south of the MC. This may indicate that southern Domain 2 was at too deep a level to record these ages, similar to rocks of Domain 1 (e.g. the MC, see above). D o m a i n 3: s t r u c t u r e s a n d m e t a m o r p h i s m

Large-scale recumbent FI nappes, such as the Goldstream Nappe in the Selkirk Fan (Brown & Tippett 1978), are recognized in most parts of Domain 3 (Brown & Tippett 1978, and references therein; Murphy 1987; Ferguson & Simony 1991; Colpron et al. 1998, and references therein; Reid et al. 2002). These F1 nappes have been interpreted as being both east-verging and west-verging, mostly based on reversals in stratigraphy (e.g. Brown & Tippett 1978). They may have developed as upright folds that pre-date the crustal-scale shear. Alternatively, it is possible that the FI nappes are early NE-vergent DFFs and later SW-vergent CFFs with no overprinting evidence preserved. F2 folds verge to the SW and their axial planes dip to the NE (Brown & Tippett 1978; Crowley & Brown 1994; Colpron et al. 1996; Gibson et al. 2005; Reid et al. 2002; Fig. 4). They are similar in style to F3 folds in, and SW of, the MC and are interpreted as CFFs. F3 folds in the Clachnacudainn Complex and Illecillewaet Synclinorium are upright, open and ESEtrending (Crowley & Brown 1994; cf. Colpron et al. 1996; Fig. 4) and in the Cariboo Mountains they are NW-trending and upright to NE-vergent (Murphy 1987). They are post-metamorphic (see below). We interpret these F3 folds in Domain 3 as being a result of late shortening, at a high structural level, above the channel flow zone or after channel flow. F3 folds on the east side of the Selkirk Fan are NE-verging (Brown & Tippett 1978; Gibson et al. 2005) and formed during or before exhumation (Gibson et al. 2005). These may be: (1) folds that formed above the channel flow zone; (2) folds that formed during late crustal shear, when channel flow reverted to

CHANNEL FLOW IN THE CANADIAN CORDILLERA detachment flow (this may happen due to collapse of the orogen; Williams & Jiang, 2005); or (3) folds that formed in front of the channel, where the shear sense is in the direction of the channel flow (cf. Beaumont et al. 2004). The latter two possibilities explain why F3 folds on the east flank of the Selkirk Fan are syn- to post-metamorphic (Brown & Tippett 1978; Gibson et al. 2005), but they require uplift of deeper structural levels within Domain 3. Uplift may be explained by Cretaceous reverse movement along faults such as the Purcell Thrust (Fig. 1; Gibson et al. 2005). Metamorphism in the Clachnacudainn Complex was at greenschist facies during F1 and middle amphibolite facies during F2 (and slightly lower along the eastern margin of the complex). Between 5 and 10 km of exhumation occurred between F2 and F3 (Crowley & Brown 1994) and F3 folding occurred after cooling and exhumation. Regional metamorphism in the Illecillewaet Synclinorium ranges from the chlorite zone at high structural levels to biotite and garnet zones at deeper structural levels (Colpron et al. 1996). Cooling and exhumation occurred during F2 folding. The highest grade of metamorphism in the Cariboo Mountains (at the deepest structural level) is in the garnet zone (Murphy 1987; Reid 2003). F3 folds were generally late or post-metamorphic (except at the deepest structural levels, where a Cretaceous metamorphic overprint exists; Reid 2003). The highest metamorphic grade in the Selkirk Fan is sillimanite-K-feldspar (Gibson et al. 2005). Exhumation of the west flank of the fan occurred in the Middle Jurassic and exhumation of the east flank in the Late Cretaceous to Early Tertiary, during or after F3 (Gibson et al. 2005). Price (1986, 2000) interpreted the Clachnacudainn terrane as being a tectonic wedge (which extends to the west and includes rocks of the SA west of the MC as well as QuesneUia; Fig. 1), from which the overlying Illecillewaet Synclinorium and Kootenay Arc delaminated (western flank of the Purcell Anticlinorium; Fig. 1). In his model, the NE-verging thrust zone in the east eventually evolves into the basal drcollement of the Rocky Mountain foreland fold and thrust belt. SW-verging thrusts, folds and fold nappes are interpreted as having developed above the wedge in the Illecillewaet Synclinorium and the Clachnacudainn Complex. The centre of the wedge is below the present-day erosion surface (see cross-section by Price 1986, p. 250). The model implies a relatively rigid 'terrane' bounded by a top-to-the-west shear zone above, and a top-to-the-east shear zone below. The channel flow model is similar, but emphasizes ductile deformation, which is evident in the rocks. Channel flow does not require, but

603

may be compatible with, two discrete shear zones. Channel flow does require the same domainal distribution of shear senses as the wedge model, and is consistent with a zone of SW-verging folds representing the upper part of the channel flow zone and a zone of NE-verging folds representing the lower part. We suggest that the channel flow zone extended farther NE than the Clachnacudainn complex, as suggested by Price (1986, 2000) for the wedge model, and included at least part of the rocks in and below the Selkirk Fan, but it is unclear where the channel flow zone ended. Murphy (1987) proposed a similar model to explain the transition from upright biotite-grade F2 folds to near-recumbent kyanite-staurolitegrade NE-dipping SW-verging F2 folds in the Cariboo Mountains. Upright folds in this model result from boundary-parallel shortening at higher structural levels above the high strain (channel flow) zone, whereas at lower structural levels, shear strain becomes important. In Murphy's (1987) model, the SW-verging folds in the Cariboo Mountains resulted from northeastward underthrusting of the North American continental plate margin, comparable to the wedge model of Price (1986, 2000). The deep structural levels, where the SW-verging near-recumbent folds are present, are at the same level as the Illecillewaet Synclinorium, and are interpreted here as representing the top part of a channel. Towards higher structural levels, SW-verging folds grade into upright folds, which are interpreted as resulting from lateral shortening above the channel flow zone. These highest structural levels did not experience detachment or channel flow and are referred to as Domain 4. Domain

3: t i m i n g o f d e f o r m a t i o n

The bulk of SW-verging structures in the Illecillewaet Synclinorium formed between c. 173 and 168 Ma (based on relationships between folding and the Fang and Tangier stocks; see Colpron et al. 1996), and cooling occurred at approximately the same time (based on biotite and muscovite 4~ cooling ages). Formation of SWverging fold nappes (F2) and cooling in the Clachnacudainn Complex are inferred to have occurred in the Middle Jurassic, based on correlation with the Illecillewaet Synclinorium (Crowley & Brown 1994; Colpron et al. 1996). SW-verging structures in the Selkirk Fan developed between 172 and 167 Ma (Gibson et al. 2005). In the Cariboo Mountains, deformation has been interpreted as being mainly of Jurassic age, based on regional correlations (Brown & Tippett 1978; Murphy 1987; Ferguson & Simony 1991). Furthermore, F 2 folding in the Cariboo Mountains is older than

604

Y.D. KUIPER ET AL.

174 Ma, the age of the Hobson Lake Pluton (Ferguson & Simony 1991; Reid et al. 2002; Reid 2003). Pigage (1977) constrained F2 folding to between the Upper Triassic and the Upper Jurassic by R b - S r dating of granodiorite intrusions. In the Illecillewaet Synclinorium and western flank of the Selkirk Fan, exhumation, fast cooling and SW-verging (F2) folding at deeper structural levels took place in the Jurassic (Colpron et al. 1996; Gibson et al. 2005). Similarly, exhumation of the Cariboo Mountains (and also the Kootenay Arc) could have been related to Middle Jurassic F2 folding. For the metamorphic grade to be generally lower in Domain 3 than in the MC, and for cooling/ exhumation to have occurred in the Middle Jurassic, Domain 3 must have been exhumed while detachment or channel flow in the MC was still ongoing (see geological model below). F3 folds in the Clachnacudainn Complex and Illecillewaet Synclinorium (and in the Selkirk Fan) formed after the rocks cooled during ongoing exhumation, during or after channel flow, and indicate shortening at high structural level. Gibson et al. (2005) interpreted F3 folds in the eastern flank of the Selkirk Fan as having formed between 104 and 84 Ma, and a thermal overprint as having occurred between 100 and 92 Ma. Metamorphism in the Clachnacudainn Complex in the middle and Late Cretaceous was related to intrusion of granitoid plutons. Hornblende and mica 4~ cooling ages are consistent with ages of Cretaceous plutonism and later extension (Colpron et al. 1999). Colpron et al. (1996) report that 110-90 Ma plutons cross-cut all structures in the Illecillewaet Synclinorium. The presence of Cretaceous plutons in Domain 3 (cf. Fig. 1) suggests that Cretaceous melting, probably related to deformation (channel flow), occurred at deeper structural levels (similar to the MC) in those areas (see below). This interpretation is consistent with interpretations of Archibald et al. (1983, 1984), who concluded that in the Kootenay Arc, late-synkinematic to post-kinematic granodioritic plutons were emplaced in Middle Jurassic time (170-165 Ma) during amphibolite-facies regional metamorphism. Mid-Cretaceous (c. 100 Ma) plutons were emplaced in a 'tectonically dormant suprastructure' (Purcell Anticlinorium; Fig. 1), accompanying renewed heating, deformation and metamorphism in the 'deepest levels of an evolving infrastructure' (Kootenay Arc; Archibald et al. 1984, p. 567). The Kootenay Arc forms the western flank of the Selkirk Fan/ Purcell Anticlinorium. The Porcupine Creek Anticlinorium, farther to the east, is separated from the Selkirk Fan/Purcell Anticlinorium by the Purcell Thrust.

Domain

4

Domain 4 is the uppermost structural level. It remained above the crustal shear zone throughout its tectonic history. In the study area, it is exposed in the northern part of the Cariboo Mountains. The shallowly NE-dipping axial planes of SWverging F2 folds in Domain 3, exposed in the SE, curve towards steeper, near-upright orientations at higher structural levels, exposed in the NW (Murphy 1987). The transition from recumbent to upright structures occurs near the biotite and garnet isograds, and is accompanied by a decrease in metamorphic grade (Murphy 1987). We interpret the area of upright structures in the NW as belonging to our Domain 4.

Granitoid migration As noted above, the deepest structural levels (Domain 1), below the vergence reversal zone, are devoid of Middle Jurassic and mid-Cretaceous plutons (with the exception of the Valhalla Complex, which has mid- to Late Cretaceous plutons; Carr et al. 1987). We suggest that the reason for this is that rocks of Domain 1 were still too deep at the time of deformation and/or granite formation for granitoid crystallization (the exception to the interpretation presented here is the deepest level of Frenchman Cap, where rocks were below the base of the crustal shear zone; see above). During channel flow, melts may have formed within these deep areas (Domains 1 and 2 in the Middle Jurassic and Domain 1 in the midto Late Cretaceous), but then migrated upwards and crystallized in the overlying rocks (Domain 3 in the Middle Jurassic and Domains 2 and 3 in the mid- to Late Cretaceous; Fig. 6b), leaving the deep rocks relatively dry but migmatitic (e.g. the MC; cf. Kuiper 2003). Minerals (e.g. zircon, monazite, titanite, xenotime) used for U - P b dating that grew at earlier times in Domain 1 (e.g. Proterozoic in the MC; Armstrong et al. 1991; Parkinson 1991; Crowley 1997, 1999; Crowley & Ghent 1999; Vanderhaeghe et al. 1999; Kuiper 2003) were partially or completely dissolved, and the dissolved material migrated upwards with these melts to crystallize in the overlying rocks. Some undissolved crystal remnants are now present as old (e.g. Proterozoic) cores in minerals at the deeper structural levels. During crustal extension, granitoids such as the 62-52 Ma Ladybird leucogranite suite (Carr et al. 1987; Carr 1992; Simony & Carr 1997) and the 51 Ma Coryell syenite (Parrish et al. 1988; Ghosh 1995b) probably formed by decompressional melting and crystallized at levels that are at the present-day erosional surface or higher.

C H A N N E L F L O W IN THE C A N A D I A N C O R D I L L E R A

The model of melt migration is similar to models proposed for the Himalayan orogen (e.g. Beaumont et al. 2001; Grujic et al. 2002; Scaillet & Searle 2006). A channel flow model has been proposed by Beaumont et al. (2001) and Grujic et al. (2002) for the Greater Himalayan Sequence, with channel flow and southward extrusion occurring between the Main Central Thrust and the South Tibetan Detachment. Migmatites formed within the channel flow zone, whereas leucogranites are more abundant in the top half of the channel flow zone (e.g. Coleman 1998; Searle 1999; Beaumont et al. 2001; Grujic et al. 2002; Scaillet & Searle 2006), with dykes and sills intruding the upper Greater Himalayan Sequence and above the channel flow zone boundary or South Tibetan Detachment (Coleman 1998; Grujic et al. 2002). Coleman (1998) proposed that the leucogranites are related to the migmatites below, based on proximity and timing. For our model of the southern Canadian Cordillera to be viable, the chemical signature of Jurassic and Cretaceous granitoids needs to be consistent with their inferred origin (i.e. Proterozoic North American crustal rocks, with possible input from the mantle). Eocene leucogranite and syenite suites may have a larger mantle component, and/ or they may have a crustal source. Sr and Nd isotopes are useful for tracing the origin of the granitoids. Additional information can be derived from inherited material (e.g. zircon) and the petrological composition of the granitoids. Origin o f M i d d l e - L a t e

Jurassic granitoids

In the Middle Jurassic, the westernmost crustally contaminated granitoids in the southern Canadian Cordillera with negative eNa values (-9.1 to -1.1) and initial St/ 86 Sr ratios greater than 0.704 (0.704-0.712; alkalic and calc-alkalic intrusives) occur within the Omineca Belt and east of Okanagan Lake in Quesnellia, (Fig. 1; Armstrong 1988; Ghosh 1995a). Quesnellia is a terrane within the Intermontane Superterrane (Monger et al. 1991; cf. Fig. 1). Early Proterozoic zircon inheritance and the presence of muscovite crystals in the crustally contaminated granitoids (Armstrong 1988; Ghosh 1995a) confirm their (partial) crustal derivation. Intrusive rocks around and west of Okanagan Lake do not show evidence for crustal contamination, suggesting absence of old crustal basement in that area. Based on geochemical and Nd isotopic data from Triassic metasedimentary rocks in southern Quesnellia, Unterschutz et al. (2002) concluded that crustal contamination from old North American continental rocks occurred earlier (Triassic), and farther west (west of Okanagan Lake).

605

Origin o f C r e t a c e o u s granitoids Cretaceous granites within Quesnellia east of Okanagan Lake are mantle-derived and have high to moderate degrees of crustal contamination (eNa = - 8 . 5 to 2.4 and 87Sr/86Sr = 0.704 to 0.7094; Brandon & Lambert 1993; Ghosh 1995a). 87Sr/S6Sr ratios are up to 0.730 in the Omineca Belt, indicating that these magmas involved large percentages of older crust, and a mantle component may not be involved (Armstrong 1988). Many of the granites exhibit S-type characteristics and enrichment in A1 and K (Armstrong 1988). Furthermore, the presence of biotite, muscovite and garnet in the granites, and to a lesser extent hornblende, indicates crustal contamination (Ghosh 1995a). Based on Sr and Nd isotopic data, Brandon & Lambert (1993) interpreted mid-Cretaceous granitoids as having formed largely due to melting of Precambrian crustal basement gneisses and Proterozoic metapelites in the Kootenay Arc. Peraluminous c. 100 Ma granites in the northern Monashee Mountains have high 87Sr/86Sr ratios (0.71492-0.72786) and inherited Precambrian zircons (Sevigny et al. 1989; Kyser et al. 1994), providing further evidence for a Precambrian crustal source. Origin o f P a l a e o c e n e - E o c e n e

granitoids

The 62-52 Ma Ladybird leucogranite suite (Carr et al. 1987; Carr 1992; Simony & Carr 1997) and the 51 Ma Coryell syenite (Parrish et al. 1988; Ghosh 1995b) have very low negative eNa values (-13.7 to -6.8) and high 87 Sr/ 86~Sr ratios (0.7071 to 0.7095). These values suggest that their magma sources were associated with crustal melting of a model 1.8 Ga old sialic crust in the case of the Ladybird suite, or incorporation of old crustal materials within more primitive (possibly mantlederived) magma in the case of the Coryell syenite (Ghosh 1995a, b). Alkalic magmatism and biotite-hornblende granite intrusion in the Palaeocene and Eocene suggest that more primitive magma was also present. Peraluminous c. 63 Ma granites in the northern Monashee Mountains have high 87Sr/86Sr ratios (0.71533-0.74181) and inherited Precambrian zircons (Sevigny et al. 1989; Kyser et al. 1994), supporting a Precambrian crustal source. C o n s i s t e n c y with the m o d e l p r e s e n t e d in this p a p e r The incorporation of progressively greater amounts of crustal material and less primitive mantle material from the Early Jurassic until the Late Cretaceous supports models incorporating thickening and melting of the crust (during detachment

606

Y.D. KUIPER E T AL.

and/or channel flow). The zone of thickened crust became geographically wider from the Jurassic to the Cretaceous (cf. Armstrong 1988; Ghosh 1995a). Early Proterozoic crust (e.g. the MC) being a source for Middle Jurassic and Cretaceous plutons in the Omineca Belt is consistent with the model presented here, in which melts migrated up from the MC at those times, while detachment/ channel flow was ongoing. Renewed input of primitive melts (including lamprophyres; e.g. Lane 1984; Adams et al. 2005) in the Palaeocene and Eocene is consistent with extension (and erosion) at that time, rather than continued crustal thickening, which would have resulted in further increase in crustal contamination. Extension is a reasonable cause for decompressional melting of (Early Proterozoic) crust, with simultaneous input from the mantle.

Geological model Four pre-extensional structural levels (Domains 14) can be recognized in the southern Omineca Belt (Fig. 3) and the following generalizations can be made. Domain 1 is the deepest structural level below the F3 vergence-reversal zone, where both DFFs and CFFs are NE-verging, and where the metamorphism was at upper amphibolite facies during both folding events. Currently exposed parts of Domain 1 include the MC, Valhalla Complex and Malton Gneiss (Figs 2 & 6). Domain 2 is the structural level located immediately above the F3 vergence-reversal zone (Figs 2 & 6), where DFFs or early CFFs are NE-verging and later CFFs are SW-verging, and where both deformation events occurred during upperamphibolite-facies metamorphism. Domain 3 is the structural level that contains SW-verging channel flow structures developed at (lower) amphibolite-facies conditions (Selkirk Fan and related areas; the Illecillewaet Synclinorium, Clachnacudainn Complex and Cariboo Mountains; Figs 2 & 6). Domain 4 is the highest structural level, where folds are upright and developed above the channel flow zone and above the biotite isograd (Fig. 3). Similar post-metamorphic upright or steeply inclined folds are present in the other three domains where they are the youngest folds. The northernmost part of Domain 2 contains NEverging DFFs or CFFs overprinted by postmetamorphic late upright folds (Fig. 6). Figure 2 summarizes our interpretation of fold overprinting relationships. During detachment flow (pre-Middle Jurassic?), NE-verging folds may have formed at some level in Domains 1 and 2. During channel flow, (Middle Jurassic) SW-verging folds developed in Domain 3 (this is the geologically oldest evidence for channel flow), while

NE-verging folding continued in Domains 1 and 2. Later during channel flow, in the mid- to Late Cretaceous, NE-verging folds were overprinted by NE-verging folds in Domain 1, and by SW-verging folds in Domain 2. To explain SW-verging folds in Domain 3, and the absence of earlier NE-verging folds in this domain, the channel flow zone in the Middle Jurassic has been drawn wider than the detachment zone before that time (Fig. 2a & b). Widening of the channel flow zone is consistent with an increase in temperature due to radioactive heating and consequent weakening of the crust, as is expected during thickening of the tectonic pile (Beaumont et al. 2004). Lower structural levels (Domains 1 and 2) continued flowing towards the NE in the Middle Jurassic. This implies that, in the Middle Jurassic, the parts of Domains 1 and 2 that are now exposed (e.g. the MC) were situated farther west of the exposed parts of Domain 3 (e.g. the Selkirk Fan) than in the Cretaceous or at present. The relative geographic positions of the various domains now exposed at the topographic surface may thus have changed during progressive deformation. In the Cretaceous, the channel flow zone moved down through the rock mass as the crust eroded and cooled from above while the crust continued to heat. The thickness of the crust need not have been decreasing since horizontal shortening continued, explaining the development of upright or steeply inclined, post-metamorphic folds overprinting SW-verging older folds in Domain 3. As the rocks of Domain 3 moved out of the channel flow zone, the rocks of Domain 2 moved from the lower part of the zone to the upper part and Domain 1 remained in the lower part of the zone (Fig. 2c).

Shortening in the Omineca belt and the Rocky Mountain fold and thrust belt The connection between channel flow in the southern Omineca Belt and thrusting within the Rocky Mountains may be indirect. For example, if shortening of the crust produced a mountain range, a critical slope could be achieved at the steepening edge of the edifice, so that thrusting in the Rocky Mountains could occur in response to local conditions (Fig. 8) rather than in response to a direct push via thrust faults such as the proposed Monashee D6collement. Similar interpretations have been proposed previously. Price & Mountjoy (1970) inferred that shortening due to thrusting in the Rocky Mountains should be associated with lateral spreading in the upwelling infrastructure to the west. Vanderhaeghe & Teyssier (1997) suggested that the latest

CHANNEL FLOW IN THE CANADIAN CORDILLERA

607

We wish to thank P. McNeill and S. Carr for their useful comments. Constructive reviews by P. Simony, R. Law (editor) and an anonymous reviewer significantly improved the manuscript. This work was part of Y.D.K.'s PhD research at the University of New Brunswick and was supported by an NSERC grant to P.F. Williams.

References

Fig. 8. Cartoon representing formation of the Rocky Mountain foreland fold and thrust belt. Thrusts develop in response to the steepening slope, and the resulting sliding of material, at the interface between thickened and normal crust. As the interface moves progressively to the right (a-c) deeper thrusts develop. Shortening of the thickened crust is achieved by folding. The Monashee Complex lies deep in the thickened crust where the upright folds are modified by shear (not shown in the cartoon).

movement in the Rocky Mountain Foreland Belt is related to gravity-driven lateral spreading of the upper crust in the hinterland, or Omineca Belt.

Conclusions Most of the structural, metamorphic and (reinterpreted) geochronological data in the southern Omineca Belt are consistent with a channel flow model. The channel flow model explains why: (a) in the deepest structural level NE-verging structures (possibly resulting from detachment flow, or from deformation in the lower part of the channel flow zone) are not overprinted by SW-verging structures; (b) intermediate structural levels contain the same NE-verging structures overprinted by SW-verging structures (from deformation in the upper half of the channel flow zone); (c) the highest levels contain SW-verging structures (formed in the highest part of the channel); (d) SW-verging structures are never overprinted by NE-verging recumbent structures; (e) high structural levels cooled while deeper structural levels continued to heat. Channel flow probably occurred mainly between the Middle Jurassic and the mid- to Late Cretaceous. Progressively younger stages of channel flow occurred towards deeper structural levels (from Middle Jurassic to Late Cretaceous), probably because the rocks moved upwards relative to the channel flow zone boundary through time. Detachment flow, and possibly channel flow, may have started before the Middle Jurassic.

ADAMS, M. G., LENTZ, D. R., SHAW, C. S. J., WILLIAMS, P. F., ARCHIBALD, D. A. & COUSENS, B. 2005. Eocene shoshonitic mafic dykes intruding the Monashee Complex, British Columbia: a petrogenetic relationship with the Kamloops Group volcanic sequence? Canadian Journal of Earth Sciences, 42, 11-24. ARCHIBALD, D. A., GLOVER, J. K., PRICE, R. A., FARRAR, E. 8s CARM[CHAEL,D. M. 1983. Geochronology and tectonic implications of magmatism and metamorphism, southern Kootenay Arc and neighboufing regions, southeastern British Columbia. Part I: Jurassic to mid-Cretaceous. Canadian Journal of Earth Sciences, 20, 1891-1913. ARCHIBALD, D. A., KROGH, T. E., ARMSTRONG, R. L. 8r FARRAR, E. 1984. Geochronology and tectonic implications of magmatism and metamorphism, southern Kootenay Arc and neighbouring regions, southeastern British Columbia. Part II: MidCretaceous to Eocene. Canadian Journal of Earth Sciences, 21, 567-583. ARMSTRONG, R. L. 1988. Mesozoic and early Cenozoic magmatic evolution of the Canadian Cordillera. In: CLARK, S. P., BURCHFIEL, B. C. c~r SUPPE, J. (eds) Processes in Continental Lithospheric Processes. Geological Society of America Special Paper, 218, 55-91. ARMSTRONG, R. L., PARRISH, R. R., VAN DER HEYDEN, P., SCOTT, K. &; RUNKLE, D. 1991. Early Proterozoic basement exposures in the southern Canadian Cordillera: core gneiss of Frenchman Cap, Unit I of the Grand Forks Gneiss, and the Vaseaux Formation. Canadian Journal of Earth Sciences, 28, 1169-1201. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2001. Himalayan tectonics explained by extrusion of a low-viscosity channel coupled to focused surface denudation. Nature, 414, 738-742. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & Medvedev S. 2004. Crustal channel flows: 1. Numerical models with applications to the tectonics of the Himalayan-Tibetan Orogen. Journal of Geophysical Research, 109, B06406. DOI: 10.1029/2003JB002809. BIRD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10,275-10,286. BRANDON, A. D. & LAMBERT, R. St. J. 1993. Geochemical characterization of mid-Cretaceous granitoids of the Kootenay Arc in the southern Canadian Cordillera. Canadian Journal of Earth Sciences, 30, 1076-1090.

608

Y.D. KUIPER ET AL.

BROWN, R. L. 1980. Frenchman Cap dome, Shuswap complex, British Columbia: a progress report. In: Current Research, Part A. Geological Survey of Canada, Paper 80-1A, 47-51. BROWN, R. L. & CARR, S. D. 1990. Lithospheric thickening and orogenic collapse within the Canadian Cordillera. Proceedings of the Pacific Rim '90 Congress, Brisbane, Australia, Australasian Institute of Mining and Metallurgy, 2, l - 1 0 . BROWN, R. L. & GIBSON, H. D. 2006. An argument for channel flow in the southern Canadian Cordillera and comparison with Himalayan tectonics. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 543 -559. BROWN, R. L. & JOURNEAY, J. M. 1987. Tectonic denudation of the Shuswap metamorphic terrane of southeastern British Columbia. Geology, 15, 142-146. BROWN, R. L. & TIPPETT, C. R. 1978. The Selkirk fan structure of the southeastem Canadian Cordillera. Geological Society of America Bulletin, 89, 548-558. BROWN, R. L., JOURNEAY, J. M., LANE, L. S., MURPHY, D. C. & REES, C. J. 1986. Obduction, backfolding, and piggyback thrusting in the metamorphic hinterland of the southeastern Canadian Cordillera. Journal of Structural Geology, 8, 255 -268. BROWN, R. L., CARR, S. D., JOHNSON, B. J., COLEMAN, V. J., COOK, F. A. & VARSEK, J. L. 1992. The Monashee d6collement of the southern Canadian Cordillera: a crustal-scale shear zone linking the Rocky Mountain foreland belt to lower crust beneath accreted terranes. In: MCCLAY, K. R. (ed.) Thrust Tectonics. Chapman and Hall, London, 357-364. CARR, S. D. 1991. Three crustal zones in the ThorOdin - Pinnacles area, southern Omineca Belt, British Columbia. Canadian Journal of Earth Sciences, 28, 2003-2023. CARR, S. D. 1992. Tectonic setting and U - P b geochronology of the early Tertiary Ladybird leucogranite suite, T h o r - O d i n - Pinnacles area, southern Omineca Belt, British Columbia. Tectonics, l l , 258-278. CARR, S. D. 1995. The southern Omineca Belt, British Columbia: new perspectives from the Lithoprobe Geoscience Program. Canadian Journal of Earth Sciences, 32, 1720-1739. CARR, S. D. & SIMONY, P. 2006. Ductile thrusting versus channel flow in the southeastern Canadian Cordillera" evolution of a coherent crystalline thrust sheet. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 559-578. CARR, S. D., PARRISH, R. R. ~z BROWN, R. L. 1987. Eocene structural development of the Valhalla complex, southeastern British Columbia. Tectonics, 6, 175-196.

CASEY, M. & DIETRICH, D. 1997. Overthrust shear in mountain building. In: SENGUPTA, S. (ed.) Evolution of Geological Structures in Micro- to Macro -scales. Chapman and Hall, London, 119-142. CLARK, M. K. ~z ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. COLEMAN, M. E. 1998. U - P b constraints on Oligocene-Miocene deformation and anatexis within the central Himalaya, Marsyandi Valley, Nepal. American Journal of Science, 298, 553-571. COLPRON, M., PRICE, R. A., ARCHIBALD, D. A. CARMICHAEL, D. M. 1996. Middle Jurassic exhumation along the western flank of the Selkirk fan structure: Thermobarometric and thermochronometric constraints from the Illecillewaet synclinorium, southeastern British Columbia. Geological Society of America Bulletin, 108, 1372-1392. COLPRON, M., WARREN, M. J. & PRICE, R. A. 1998. Selkirk fan structure, southeastern Canadian Cordillera: Tectonic wedging against an inherited basement ramp. Geological Society of America Bulletin, 110, 1060-1074. COLPRON, M., PRICE, R. A. & ARCHIBALD, D. A. 1999. 4~ thermochronometric constraints on the tectonic evolution of the Clachnacudainn complex, southeastern British Columbia. Canadian Journal of Earth Sciences, 36, 1989-2006. COOK, F. A. 1995. Lithospheric processes and products in the southern Canadian Cordillera: a Lithoprobe perspective. Canadian Journal of Earth Sciences, 32, 1803-1824. CROWLEY, J. L. 1997. U - P b geochronologic constraints on the cover sequence of the Monashee complex, Canadian Cordillera: Paleoproterozoic deposition on basement. Canadian Journal of Earth Sciences, 34, 1008-1022. CROWLEY, J. L. 1999. U - P b geochronologic constraints on Paleoproterozoic tectonism in the Monashee complex, Canadian Cordillera: elucidating an overprinted geologic history. Geological Society of America Bulletin, 111, 560-577. CROWLEY, J. L. 8z BROWN, R. L. 1994. Tectonic links between the Clachnacudainn terrane and Selkirk allochthon, southern Omineca Belt, Canadian Cordillera. Tectonics, 13, 1035-1051. CROWLEY, J. L. & GHENT, E. D. 1999. An electron microprobe study of the U - T h - P b systematics of metamorphosed monazite: the role of Pb diffusion versus overgrowth and recrystallization. Chemical Geology, 157, 285-302. CROWLEY, J. L. & PARRISH, R. R. 1999. U - P b isotopic constraints on diachronous metamorphism in the northern Monashee complex, southern Canadian Cordillera. Journal of Metamorphic Geology, 17, 483-502. CROWLEY, J. L., GHENT, E. D., CARR, S. D., SIMONY, P. S. & HAMILTON, M. A. 2000. Multiple thermotectonic events in a continuous metamorphic sequence, Mica Creek area, southeastern Canadian Cordillera. Geological Materials Research, 2, 1-45.

CHANNEL FLOW IN THE CANADIAN CORDILLERA CROWLEY, J. L., BROWN, R. L. & PARRISH, R. R. 2001. Diachronous deformation and a strain gradient beneath the Selkirk allochthon, northern Monashee complex, southeastern Canadian Cordillera. Journal of Structural Geology, 23, 1103-1121. CROWLEY, J. L., BROWN, R. L., GIBSON, H. D. & TUBRETT, M., 2005. Weakly deformed c. 1.85 Ga granites in the Monashee complex, British Columbia: evidence for a rigid basement during Tertiary orogenesis in the hinterland of the Canadian Cordillera. Cordilleran Tectonics Workshop Meeting ( 18 - 20 February 2005), Whitehorse, Yukon, 8 - 9. DAVIDSON, C., GRUJIC, D. E., HOLLISTER, L. S. & SCHMID, S. M., 1997. Metamorphic reactions related to decompression and synkinematic intrusion of leucogranite, High Himalayan Crystallines, Bhutan. Journal of Metamorphic Geology, 15, 593-612. DIGEL, S. G., GHENT, E. D., CARR, S. D. t~ SIMONY, P. S. 1998. Early Cretaceous kyanite-sillimanite metamorphism and Paleocene sillimanite overprint near Mount Cheadle, southeastern British Columbia: geometry, geochronology, and metamorphic implications. Canadian Journal of Earth Sciences, 35, 1070-1087. FERGUSON, C. A. & SIMONY, P. S. 1991. Preliminary report on structural evolution and stratigraphic correlations, northern Cariboo Mountains, British Columbia. Geological Survey of Canada, Paper 91-1A, 103-110. FOSTER, G., GIBSON, H. D., PARRISH, R. R., HORSTWOOD, M., FRASER, J. & TINDLE, A. 2002. Textural, chemical and isotopic insights into the nature and behaviour of metamorphic monazite. Chemical Geology, 191, 183-207. FYSON, W. K. 1970. Structural relations in metamorphic rocks, Shuswap Lake area, British Columbia. In: WHEELER, J. O. (ed) Structure of the Southern Canadian Cordillera. Geological Association of Canada, Special Paper, 6, 107-122. GERVAIS, F., BROWN, R. L. & CROWLEY, J. L. 2005. Proterozoic strain in the core of the Frenchman Cap dome, Monashee complex; tectonic implications of a strong basement. Cordilleran Tectonics Workshop Meeting (18- 20 February 2005), Whitehorse, Yukon, 10-11. GHOSH, D. K. 1995a. N d - S r isotopic constraints on the interactions of the Intermontane Superterrane with the western edge of North America in the southern Canadian Cordillera. Canadian Journal of Earth Sciences, 32, 1740-1758. GHOSH, D. K. 1995b. U - P b geochronology of Jurassic to early Tertiary granitic intrusives from the Nelson-Castlegar area, southeastern British Columbia, Canada. Canadian Journal of Earth Sciences, 32, 1668-1680. GIBSON, H. D., BROWN, R. L. 8z PARRISH, R. R. 1999. Deformation-induced inverted metamorphic field gradients: an example from the southeastern Canadian Cordillera. Journal of Structural Geology, 21, 751-767. GIBSON n. D., BROWN, R. L. 8z CARR, S. D. 2005. U T H - P b geochronologic constraints on the structural evolution of the Selkirk fan, northern

609

Selkirk Mountains, southern Canadian Cordillera. Journal of Structural Geology, 27, 1899-1924. GRUJIC, D., CASEY, M., DAVIDSON, C., HOLLISTER, L. S., KUNDIG, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalayan Crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-44. GRUJIC, D., HOLLISTER, L. S. & PARRISH, R. R. 2002. Himalayan metamorphic sequence as an orogenic channel: insight from Bhutan. Earth and Planetary Science Letters, 198, 177-191. JAMIESON, R. A., BEAUMONT, C., NGUYEN, M. H. & LEE, B. 2002. Interaction of metamorphism, deformation and exhumation in large convergent orogens. Journal of Metamorphic Geology, 20, 9-24. JAMIESON, R. A., BEAUMONT, C., ~EDVEDEV, S. & NGUYEN M. H. 2004. Crustal channel flows: 2. Numerical models with implications for metamorphism in the Himalayan-Tibetan orogen. Journal of Geophysical Research, 109, B06407. DOI: 10.1029/2003JB002811. JOHNSON, B. J. 1994. Structure and tectonic setting of the Okanagan Valley fault system in the Shuswap Lake area, southern British Columbia. PhD Thesis, Carleton University, Ottawa. JOHNSON, B. J. & BROWN, R. L. 1996. Crustal structure and early Tertiary extensional tectonics of the Omineca belt at 51 ~ latitude, southern Canadian Cordillera. Canadian Journal of Earth Sciences, 33, 1596-1611. JOHNSTON, D. H., WILLIAMS, P. F., BROWN, R. L., CROWLEY, J. L. & CARR, S. D. 2000. Northeastward extrusion and extensional exhumation of crystalline rocks of the Monashee Complex, southeastern Canadian Cordillera. Journal of Structural Geology, 22, 603-625. JOURNEAY, J. M. 1986. Stratigraphy, internal strain, and thermo-tectonic evolution of northern Frenchman Cap dome, an exhumed basement duplex structure, Omineca hinterland, southeastern Canadian Cordillera. PhD Thesis, Queen's University, Kingston, Ontario. JOURNEAY, J. M. & WILLIAMS, S. P. 1995. GIS map Library: a window on Cordilleran geology (v. 1.0). Geological Survey of Canada, Open File 2948. KRUSE, S., McNuTT, M., PHIPPS-MORGAN, J., ROYDEN, L. & WERNICKE, B. 1991. Lithospheric extension near Lake Mead, Nevada: A model for ductile flow in the lower crust. Journal of Geophysical Research, 96, 4435-4456. KRUSE, S. • WILLIAMS, P. F. 2006. The Victor Creek Fanlt of southern British Columbia: constraints on geometry and kinematics. Canadian Journal of Earth Sciences, 42, 2141-2160. KRUSE, S., McNEILL, P. & WILLIAMS, P. F. 2003. A new compilation map of the Thor-Odin culmination, Monashee complex, BC. Geological Association of Canada Mineralogical Association of Canada and Society of Economic Geologists Annual Meeting (25-29 May 2003), Vancouver. KRUSE, S., McNEILL, P. & WILLIAMS, P. F. 2004. Geological map of the Thor-Odin dome, southern

610

Y.D. KUIPER ET AL.

Monashee complex, BC. Website: http://www. unb.ca/fredericton/science/geology/monashee KUIPER, Y. D. 2003. Isotopic constraints on timing of deformation and metamorphism in Thor-Odin, Monashee Complex, southeastern British Columbia. PhD Thesis, University of New Brunswick, Fredericton. KYSER, T. K., GHENT, E. D. & SEVIGNY, J. H. 1994. Petrogenesis of two-mica granites, northern Monashee Mountains, British Columbia: Evidence from oxygen, hydrogen and strontium isotopic compositions. Geological Society of America Bulletin, 106, 1531-1537. LANE, L. S. 1984. Deformation history of the Monashee Decollement and Columbia River Fault zone. PhD Thesis, Carleton University, Ottawa. MCDONOUGH, M. R. & SIMONY, P. S. 1988. Structural evolution of basement gneisses and Hadrynian cover, Bulldog Creek area, Rocky Mountains, British Columbia. Canadian Journal of Earth Sciences, 25, 1687-1702. MCKENZIE, D., NIMMO, F., JACKSON, J. A., GANS, P. B. & MmLER, E. L. 2000. Characteristics and consequences of flow in the lower crust. Journal of Geophysical Research, 105, 11,029-11,046. MANCKTELOW, N. S. 1995. Nonlithostatic pressure during sediment subduction and the development and exhumation of high pressure metamorphic rocks. Journal of Geophysical Research, 100, 571-583. MEANS, W. D. 1995. Shear zones and rock history. Tectonophysics, 247, 157-160. MONGER, J. W. H., PRICE, R. A. & TEMPELMANKLUIT, D. J. 1982. Tectonic accretion and the origin of the two major metamorphic and plutonic welts in the Canadian Cordillera. Geology, 10, 70-75. MONGER, J. W. H., WHEELER, J. O., TIPPER, H. W., ET AL. 1991. Part B. Cordilleran terranes. In: GABR1ELSE, H. & YORATH, C. J. (eds) Upper Devonian to Middle Jurassic Assemblages, Geology of the Cordilleran Orogen in Canada, Geological Survey of Canada, Geology of Canada 4, 281-327. MORRISON, M. L. 1982. Structure and petrology of the Malton gneiss complex. PhD Thesis, University of Calgary. MURPHY, D. C. 1987. Suprastructure/infrastructure transition, east-central Cariboo Mountains, British Columbia: geometry, kinematics and tectonic implications. Journal of Structural Geology, 9, 13-29. MURPHY, D. C., VAN DER HEYDEN, P., PARRISH, R. R., KLEPACKI, D. W., McMILLAN, W., STRU1K, L. C. & GABITES, J. 1995. New geochronological constraints on Jurassic deformation of the western edge of North America, southeastern Canadian Cordillera. In: MILLER, D. M. & BUSBY, C. (eds) Jurassic Magmatism and Tectonics of the North American Cordillera. Geological Society of America Special Paper, 299, 159-171. NORLANDER, B. H., WHITNEY, D. L., TEYSSIER, C. & VANDERHAEGHE, O. 2002. Partial melting and decompression of the Thor-Odin dome, Shuswap

metamorphic core complex, Canadian Cordillera. Lithos, 61, 103-125. PARKINSON, D. L. 1991. Age and isotopic character of Early Proterozoic basement gneisses in the southern Monashee Complex, southeastern British Columbia. Canadian Journal of Earth Sciences, 28, 1159-1168. PARRISH, R. R. 1995. Thermal evolution of the southeastern Canadian Cordillera. Canadian Journal of Earth Sciences, 32, 1618-1642. PARRISH, R. R., CARR, S. D. & PARKINSON, D. L. 1988. Eocene extensional tectonics and geochronology of the southern Omineca Belt, British Columbia and Washington. Tectonics, 7, 181 - 212. PIGAGE, L. C. 1977. R b - S r dates for granodiorite intrusions on the northeast margin of the Shuswap Metamorphic Complex, Cariboo Mountains, British Columbia. Canadian Journal of Earth Sciences, 14, 1690-1695. PRICE, R. A. 1981. The Cordilleran Thrust and Fold Belt in the southern Canadian Rocky Mountains. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Geological Society, London, Special Publications, 9, 427-448. PRICE, R. A. 1986. The southeastern Canadian Cordillera: thrust faulting, tectonic wedging, and delamination of the lithosphere. Journal of Structural Geology, 8, 239-254. PRICE, R. A. 2000. The evolution of the southern Cordilleran foreland thrust and fold belt and the kinematics of Cordilleran orogenesis. GeoCanada2000, Abstract 435. PRICE, R. A. & MOUNTJOY, E. W. 1970. Geologic structure of the Canadian Rocky Mountains between Bow and Athabasca rivers - a progress report. Geological Survey of Canada, Special Paper, 6, 7-25. RAESIDE, R. P. & SIMONY, P. S. 1982. Stratigraphy and deformational history of the Scrip Nappe, Monashee Mountains, British Columbia. Canadian Journal of Earth Sciences, 20, 639-650. RAMSAY, J. G. 1997. The geometry of a deformed unconformity in the Caledonides of NW Scotland. In: SENGUPTA, S. (ed.) Evolution of Geological Structures in Micro- to Macro-scales. Chapman and Hall, London, 445-472. READ, P. B. & BROWN, R. L. 1981. Columbia River fault zone: southeastern margin of the Shuswap and Monashee complexes, southern British Columbia. Canadian Journal of Earth Sciences, 18, 1127-1145. REESOR, J. E. & MOORE, J. M. 1971. Petrology and structure of the Thor-Odin gneiss dome, Shuswap metamorphic complex. Geological Survey of Canada, Bulletin, 195. REID, L. F. 2003. Stratigraphy, structure, petrology, geochronology and geochemistry of the Hobson Lake area (Cariboo Mountains, British Columbia) in relation to the tectonic evolution of the south Canadian Cordillera. PhD Thesis, University of Calgary. REID, L. F., SIMONY, P. S. & Ross, G. M. 2002. Dextral strike-slip faulting in the Cariboo Mountains, British Columbia: a natural example of

CHANNEL FLOW IN THE CANADIAN CORDILLERA wrench tectonics in relation to Cordilleran tectonics. Canadian Journal of Earth Sciences, 39, 953 -970. ROYDEN, L. H. 1996. Coupling and decoupling of crust and mantle in convergent orogens: Implications for strain partitioning in the crust, Journal of Geophysical Research, 101, 17,679-17,705. ROYDEN, L. H., BURCHFIEL, B. C., KING, R. W., CHEN, Z., SHEN, F. & LIU Y. 1997. Surface deformation and lower crustal flow in eastern Tibet. Science, 276, 788-790. SCAILLET, B. & SEARLE, M. P. 2006. Mechanisms and timescales of felsic magma segregation, ascent and emplacement in the Himalayas. In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 293-308. SCAMMELL, R. J. 1993. Mid-Cretaceous to Tertiary thermotectonic history of former mid-crustal rocks, southern Omineca Belt, Canadian Cordillera. PhD Thesis, Queen's University, Kingston, Ontario. SCHAtJBS, P. M. & CARR, S. D. 1998. Geology of metasedimentary rocks and late Cretaceous deformation history, in the Northern Valhalla complex, British Columbia. Canadian Journal of Earth Sciences, 38, 1018-1036. SCHAUBS, P. M., CARR, S. D. & BERMAN, R. G. 2002. Structural and metamorphic constraints on ca. 70 Ma deformation of the northern Valhalla complex, British Columbia: implications for the tectonic evolution of the southern Omineca belt. Journal of Structural Geology, 24, 1195-1214. SEARLE, M. P. 1999. Emplacement of Himalayan leucogranites by magma injection along giant sill complexes: examples from the Cho Oyu, Gyachuang Kang and Everest leucogranites (Nepal Himalaya). Journal of Asian Earth Sciences, 17, 773-783. SEVIGNY, J. H. & SIMONY, P. S. 1989. Geometric relationship between the Scrip Nappe and metamorphic isograds in the northern Adams River area, Monashee Mountains, British Columbia. Canadian Journal of Earth Sciences, 26, 606-610. SEV1GNY, J. H., PARRISH, R. R. & GHENT, E. D. 1989. Petrogenesis of peraluminous granites, Monashee Mountains, southeastern Canadian Cordillera. Journal of Petrology, 30, 557-581. SEVIGNY, J. H., PARRISH, R. R., DONELICK, R. A. & GHENT, E. D. 1990. Northern Monashee Mountains, Omineca Crystalline Belt, British Columbia: Timing of metamorphism, anatexis, and tectonic denudation. Geology, 18, 103-106.

611

SHEN, F., ROYDEN, L. H. & BURCHFIEL, B. C. 2001. Large-scale crustal deformation of the Tibetan Plateau. Journal of Geophysical Research, 106, 6793-68t6. SIMONY, P. S. d~;CARR, S. D. 1997. Large lateral ramps in the Eocene Valkyr shear zone: extensional ductile faulting controlled by plutonism in southern British Columbia. Journal of Structural Geology, 19, 769-784. SIMONY, P. S., GHENT, g. D., CRAW, D., MITCHELL, W. & ROBBINS, D. B. 1980. Structural and metamorphic evolution of northeast flank of Shuswap complex, southern Canoe River area, British Columbia. In: CRITTENDEN, M. D., CONEY, P. J. & DAVIS, G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America Memoir, 153, 445-460. SPARK, R. N. 2001. Crustal thickening and tectonic denudation within the Thor-Odin culmination, Monashee Complex, southern Canadian Cordillera. PhD Thesis, University of New Brunswick, Fredericton, New Brunswick, Canada. SPEAR, F. S. & PARRISH, R. R. 1996. Petrology and cooling rates of the Valhalla Complex, British Columbia, Canada. Journal of Petrology, 37, 733-765. UNTERSCHUTZ, J. L. E., CREASER, R. A., ERDMER, P., THOMPSON, R. I. & DAUGHTRY, K. L. 2002. North American margin origin of Quesnel terrane strata in the southern Canadian Cordillera: Inferences from geochemical and Nd isotopic characteristics of Triassic metasedimentary rocks. Geological Society of America Bulletin, 114, 462-475. VANDERHAEGHE, O. & TEYSSIER, C. 1997. Formation of the Shuswap metamorphic core complex during late-orogenic collapse of the Canadian Cordillera: Role of ductile thinning and partial melting of the mid- to lower crust. Geodinamica Acta, 10, 41-58. VANDERHAEGHE,O., TEYSSIER,C. &WYSOCZANSKI,R. 1999. Structural and geochronological constraints on the role of partial melting during the formation of the Shuswap metamorphic core complex at the latitude of the Thor-Odin dome, British Columbia. Canadian Journal of Earth Sciences, 36, 917-943. WESTAWAY, R. 1995. Crustal volume balance during the India-Eurasia collision and altitude of the Tibetan Plateau: A working hypothesis. Journal of Geophysical Research, 100, 15,173-15,192. WILLIAMS, P. F. & JIANG, D. 2005. An investigation of lower crustal deformation: Evidence for channel flow and its implications for tectonics and structural studies. Journal of Structural Geology, 27, 1486-1504.

Index Page n u m b e r s in italic refer to figures. Page n u m b e r s in b o l d signify entries in tables. Adamant Pluton 542 Adams-North Thompson Fault 560 Altyn Tagh fault 40, 356 Ama Dablam 363 Appalachian Inner Piedmont 517- 518, 535 attributes 520-527 cross-sections 532 dominant foliation 522 geological map 529 mineral stretching lineation 527, 528 Neocadian metamorphic isograd map 530 stromatic migmatite 525 structure 523 tectonothermal time line 531 channel flow 527-535 flow model 526 tectonic setting 518-520, 519 Astor River 208 Bangong suture 356 Bangong-Nuijang suture 40, 356 lithospheric electrical conductivity 50 seismic anisotropy 52 seismic reflectivity 54 Bhutan geological cross-section 429 geological map 416, 426, 428 geological setting 426-427 normal-sense shear zones 425-426, 439-441 cathodoluminescence (CL) images 435 geometry and kinematics 427-429, 431, 434, 441 Tera-Wasserburg diagrams 439 U - P b geochronology 433-436, 437-438 pulsed channel flow 415-417, 420-421 data 418-420, 419, 420 T - t paths 417 Bonnington pluton 566, 570 Canadian Cordillera 543, 555-556, 561-567, 562, 582-583, 589-561,607 channel flow model applied to Ominecca Belt 593 Domain 1 - structures and metamorphism 594-560 Domain 1 - timing of deformation 600-601 Domain 2 - structures and metamorphism 601-602 Domain 2 - timing of deformation 602 Domain 3 - structures and metamorphism 602-603 Domain 3 - timing of deformation 603-604 Domain 4 604 general statements 593-595 overview 597-598 tectonic interpretations 596 coherent thrust sheet architecture 569-572 comparison with High Himalaya 555 continuity of mid-crustal ductile zone 552-553

Cretaceous mid-crustal ductile zone 545-546 Big Bend area 546-549 Big Bend area deformation and exhumation 549-552, 550 ductile zone west of Monashee Complex 552 cross-sections 547, 551,563, 567, 571 detachment flow 564-565 geological background 591-593 geological model 606 geological setting 5 4 3 - 5 4 5 , 5 4 4 gneiss dome formation 554-555 granitoid migration 604-605 consistency with other models 605-606 origin of Cretaceous granitoids 605 origin of Middle-Late Jurassic granitoids 605 origin of Palaeocene-Eocene granitoids 605 Gwillim Creek shear zone 573-574 geometry and depth 575-578 large displacement and link to Rocky Mountain thrust fault systems 578-580 northward continuation 580-581 time constraints 574-575 U - P b data 576, 577, 578, 583-584 interpretation of Cretaceous mid-crustal ductile zone 553 support for channel flow model 553 shortening in the Ominecca Belt 606-607 suprastructure-infrastructure association 568-569 tectonic evolution 568 tectonic map 590 tectonic model 554 Valhalla Complex origin 560 reconciliation of crystalline thrust sheet model 581-582 upper margin 572-573 Cariboo Mountains 590 deformation, metamorphism and age constraints 599 Carolina superterrane 519, 533,534 Cat Square terrane 519, 520 cross-section 532 Champion Lakes Fault 566 Changtse 363 channel flow 25-26, 3 3 - 3 4 see also pulsed channel flow challenges and unresolved issues 14-15 ductile extrusion 207-208 channel flow in mid- and upper crust 204 ductile flow in lower crust 204 field characteristics 205-206 requisite conditions 205 thermal-mechanical models 204 thermal-mechanical models, implications for 204-205 dynamics 2 - 4 exhumation 6 extrusion 4 - 6 , 5

614

INDEX

channel flow (Continued) dynamics (Continued) kinematic relationship between channel flow and extrusion 5 viscous channel flow 4 flow relationships 27-33, 28 aggregate strength versus melt fraction 32 velocity profiles 30 GHS models 165-166, 179-180 crustal-scale model results 166-174, 170-173 design 166, 167, 168, 169 requirements and characteristics of channel flow 6 - 7 channel thickness and late-stage modifications 14 coeval channel-bounding structures 8 discontinuity of protoliths across channel 12-13 internal deformation within channel 8 - 1 2 kinematic inversions 8 lateral versus vertical transport of material 12 metamorphic characteristics 12 plateau formation 7 timing of melting and shortening structures 13 viscosity 7 Tibet flow modes 41-43, 41 channel flow folds (CFFs) 593-595 Channel Flow-Extrusion hypothesis for Himalayan-Tibetan orogen 71, 72, 83 evolution of ideas 1 - partially molten crust beneath Tibet 71-73 2 - lateral flow of Tibetan fluid layer 7 3 - 7 4 3 - eastern growth of Tibetan Plateau 74-75, 74 4 - Miocene slip on Himalayan thrust faults coeval with normal faults 75-77, 76 5 - Greater Himalayan channel roots northward 77 6 - monsoon erosion and channel extrusion linked 7 7 - 7 8 Phase I - steady state configuration during Early-Middle Miocene 7 8 - 7 9 Phase II - second emergent channel established during Middle Miocene-Early Pliocene 79-81, 80 Phase III - intensive extrusion at Himalayan front from Late Pliocene-Recent 81-82 testing the hypothesis 82 applicability to other orogenic systems 83 theoretical studies 8 2 - 83 channel injection (CI) model of continental plateau growth 161-162 model comparison 148 theory 149-150, 149 model CI-1 - transition zone with uniformproperty channel 150-151, 151 model CI-2 - constant thickness, variable viscosity channel 151-153, 152 model CI-3 - temperature-dependent channel properties 153, 154, 155 thermal and rheological consistency 153 criteria for rheological consistency 157- ! 58, 157, 158 criteria for thermal and mechanical models 154-156, 156 upper crust 159 stability 159-161

viscous thickening model 159 channel-bounding structures 8 Chekha Formation 416, 418, 426, 428 China Creek Dome 566 U - P b data 576, 577, 578, 583-584 Chomolhari 428 Clachnacudainn Complex 590 deformation, metamorphism and age constraints 599 College Creek pluton 570 Columbia River fault 544, 546, 548, 551,562, 566, 590 continental plateau growth 147-149 channel injection (CI) modelling 161 - 162 theory 149-153, 149 thermal and rheological consistency 153-158 upper crust, behaviour of 159-161 possible models 148 thermal model 162-164 continuum mechanics model for dilatency effects dilatant plane strain 186-188 isochoric plane strain 185-186, 185, 187 Couette flow 27-33, 28, 41 deformation sequence 32 velocity profiles 30 Cowrock terrane 519 crustal extension Nanga Parbat-Haramosh Massif 208-210, 208 numerical modelling 202, 202 Wing Pond Shear Zone (WPSZ), New Foundland 210-214, 211,212, 213 crustal flow flow modes in hot orogens 91-93, 137-138 classification according to temperaturemagnitude plots 92 crustal-scale model results 99-126 effect of thermal relaxation and incubation time on crustal flow 135-136 flow modes in temperature-magnitude space 134-135 infrastructure and superstructure 136-137 numerical calculation of crustal- and upper-mantle-scale flows 93-99 upper-mantle-scale model results 126-134 Tibet flow modes 41-43, 41 inferred channel flow 59 lateral strain variations, vertical strain and strength profiles 39-41 crustal-scale (CS) modelling 93 density structure and isostatic compensation 98-99 design advantages and limitations 138 complexity and tuning 138-139 Himalayan-Tibetan (HT) models 140-141 model-data comparisons 142-143 philosophy of numerical approach 138 scaling of laboratory power-law creep flow laws 139-140 testing basal boundary conditions 139 mechanical models 96-98 melt-weakening 98 parameters 95-96, 99

INDEX results 99-116 model LHO-1 - homogeneous channel flow 100-107, 116, 137 model LHO-2 - heterogeneous channel flow 108-115, 116-117, 137 model LHO-3 - hot fold nappes 117-126, 118-125, 137 topographic evolution in LHO models 126, 127 surface processes 99 thermal model 99 velocity boundary conditions and reference frames 93, 94 Cusson Creek 590 Darondi River 245 detachment flow folds (DFFs) 593-595 dilatancy effects on extrusion 183-184, 195-196 considerations 193 kinematic vorticity changes 195 slab stretching and thinning 195 volume change 193-195, 194 continuum mechanics model dilatant plane strain 186-188 isochoric plane strain 185-186, 185, 187 homogeneous deformation constant thickness 190-192, 191 finite 188-190, 189 modelling approach 184-185 progressive deformation constant thickness 192-193, 192 dilatant plane strain 186-188 Dinarides 498 drag folds 229-231,230 Drucker-Prager yield criterion 96 ductile extrusion 26-27, 201,497, 513-514 channel flow 207-208 channel flow in mid- and upper crust 204 ductile flow in lower crust 204 field characteristics 205-206 requisite conditions 205 thermal-mechanical models 204 thermal-mechanical models, implications for 204-205 orogenic thickening and crustal extrusion 201-202, 203-204, 203, 214 evidence from central and eastern Himalayas 203 numerical modelling 202, 202 dynamics of channel flow 2 East Rongbuk glacier 369, 395 Everest, Mount 373-374 crustal structure 361-364, 362-363, 395-397,396 geological background 383-385, 384 geological map 395 regional setting 355-357, 356 Greater Himalayan sequence (GHS) 357-359, 358, 360 restoration 364-366, 365, 371-373, 372, 373 timescales of metamorphism, melting and channel flow 370-371,370 exhumation 6 External Greater Himalayan klippen 257-259, 259 extrusion 4 - 6 see also ductile extrusion

615 extrusive flow 225-227, 226 kinematic relationship between channel flow and extrusion 5

flow partitioning in the Himalaya 379-381,403-409 channel flow models 381 distribution of flow regimes 402-403 extrusion models 381-382, 381 kinematic models 382-383 thermal-mechanical models 383 geological background of Everest transect 383-385, 384 geological map 380 lithological, structural and temporal controls 401-402 petrography Everest summit-Kangshung Valley 395-397 Main Central Thrust (MCT) 397-399, 398, 399 Rongbuk valley transects 389-395, 391 tectonic setting 381 vorticity analysis 386, 388, 390, 393, 394 Everest transect 401 rigid grain data plots 385-389, 387, 404-408 techniques 385 Foreland Belt 544, 548 Frenchman Cap Dome 544, 551,562, 590 Gangdese batholith 360 Gaza Dzong 428 Ghat 362, 365 Gilgit River 208 Grandfather Mountain window 519 Greater Himalayan sequence (GHS) 2, 25, 357-359, 373 -374 buckling at Nar valley, central Nepal 269-270 exposed deformation features 285 geology 272-274 implications for cooling rate 285-287, 286 implications for southward extrusion 287-289, 288 structural constraints 274-279 timing constraints 279-285 channel flow models 165-166, 179-180, 359-361 crustal-scale model results 166-174, 170-173 design 166, 167, 168, 169 channel thickness and late-stage modifications 14 channel-bounding structures 8, 9 crustal extrusion 203 distribution of flow regimes 402-403 exhumation 6 geological map 358 INDEPTH profile 361 leucosome and leucogranite bodies 13 metamorphic characteristcs 12 pressure-temperature diagram 417 protoliths 12-13 provenance 174-176, 175, 177, 178 NHG domes and leucogranites 176-179 restoration 364-366, 365, 371-373, 372, 373 timescales of metamorphism, melting and channel flow 370-371,370

616

INDEX

Greater Himalayan sequence (Continued) Slab core 399-400, 400 transport of materials 12 viscosity 6 Gwillim Creek shear zone 562, 566, 572, 573-574, 590 geometry and depth 575-578 U-Pb data 572, 577, 578, 583-584 large displacement and link to Rocky Mountain thrust fault systems 578-580 northward continuation 580- 581 time constraints 574-575 Harramosh, Mount 208 Helenides, External 497 comparison with other ductile extrusion zones 513-514 geotectonic framework 498-499, 498 Phyllite-Quartzite (PQ) unit 498-499, 498 deformation temperatures 504-506 displacement and extrusion rate 510-513, 512 finite strain analysis 506 finite strain symmetry 504 interpretation and synthesis 509-510 kinematic vorticity analysis 508-509, 509 major fabric elements 499-500 patterns and shear sense 502-503 quartz c-axis fabrics 500-506 sampling 500, 501 stable-orientation analysis 508 thinning and dip-parallel elongation 510, 511 vorticity analysis 506-508, 507 Plattenkalk (PLK) unit 498-499 Hermit's Gorge 363, 368, 386, 391,394-395,406 Herzssprung-Russell diagram for classification of stars, used as template for orogen classification 92 heterogeneous deformation 497 High Himalayan leucogranites (HLL) 294, 296 final stage of building 302-303 petrogenesis 294-297, 295 thermal evolution 309-312, 310, 311 geological constraints 312- 313 implications for Miocene denudation of Himalaya 323-324 numerical model 313-315, 313 numerical model results 315-323 Himalayan orogen 3 see also Tethyan Himalaya flow partitioning 379-381,403-409 channel flow models 381 distribution of flow regimes 402-403 extrusion models 381-383, 381 geological background of Everest transect 383-385, 384 Greater Himalayan Slab 399-400, 400 lithological, structural and temporal controls 401-402 petrography 389-399 vorticity analysis 385-389, 386, 387, 388, 390, 393, 394, 401,404-408 geological map 380 interpreted ages of motion 9, 10-11 magmatic processes 293-294, 304-305

compaction 298-300 final stage of HHL building 302-303 geological setting 294-297, 295, 296 melt extraction 297-301 melt extraction, consequences for crustal flow 304 shearing 300-301 timescales of magma cooling 303-304 regional setting 447 regional tectonic map 446 tectonic setting 381 Himalaya-Tibetan plateau system 2, 14-15, 25 accretion/rapid denudation models focused denudation-induced channel flow 246-247 focused denudation-induced channel flow, limitations 247 thrust ramp models 246 uniqueness of predictions 248 unresolved features 247-248 channel flow 256-257 Channel Flow-Extrusion hypothesis 71, 72, 83 evolution of ideas 71-78 Phase I - steady state configuration during Early-Middle Miocene 78-79 Phase II - second emergent channel established during Middle Miocene-Early Pliocene 79-81 Phase II - second emergent channel established during Middle Miocene-Early Pliocene 80 Phase I I I - intensive extrusion at Himalayan front from Late Pliocene-Recent 81-82 testing the hypothesis 82-83 deformed migmatites 26 geological setting background 237-238, 239 inverted metamorphism 238-240 paired leucogranite belts 240 Tibetan rifts 240-241 Himalaya extruded from beneath Tibet theory 237, 248-249 timing 259-261,261 shallow Tibetan anatexis model INDEPTH survey 241 supporting evidence 241-242, 242 Zhao and Morgan hypothesis 241 shallow Tibetan anatexis model, critique of 246 consistency with 3He/4He data 243-244 consistency with absence of Gangdese zircons 244-245 consistency with cold southern Tibetan Moho 243 consistency with geology and geochemistry of Yangbajain rift 244 crustal thickening explanation 244 nature of 'bright spots' 243 representativeness of Yadong-Gulu rift 242-243 stratigraphy 245-246, 245 homogeneous deformation constant thickness 190-192, 191 finite 188-190, 189

INDEX Illecillewaet Synclinorium 590 Indus River 208 Indus-Tsangpo Suture Zone (ITSZ) 445 Indus-Yarlung suture 40 seismic reflectivity 54 infrastructure zones 221-222, 222, 232-233 drag folds 229-231,230 kinematic vorticity number and strain compatibility 224-225, 228-229 distributed pure shear 228 extrusive flow 225-227, 226, 227 volume loss 227-228 model characterisation and kinematics 222-223 channel flow 223 objectives 223-224, 224 transport flow 223 strain localization 232 tectonic models 231 isochoric plane strain 185-186, 185, 187 Jiali fault 40, 240, 356 Jinsha River suture 40, 356 lithospheric electrical conductivity 50 Joss Mountain 588 Kahtang thrust 416, 419-420, 426, 428 Kailas, Mount 356 Kampa Dome 446 Kangmar Dome 40, 446, 473 geological map 473 stereographic plot 478 Kangmar Thrust 360 Kangshung Valley 395 Karakax fault 40 Karakoram Fault 356 Karakoram Range 208 Karakoram-Jiali fault system seismic anisotropy 52 Karikhola 362 Khang Bum 428, 429 Khumbu glacier 395 Khumbu Thrust 363, 365 kinematic inversions 8 kinematic vorticity 195, 224-225, 228-229 distributed pure shear 228 extrusive flow 225-227, 226, 227 volume loss 227-228 Kinnaird pluton 570 Kohistan 208 Kootenay Arc 562, 563, 590 Kootenay River 566 Kun Lun suture 356 Kunlun fault 40, 356 lithospheric electrical conductivity 50 Kunlun Shan lithospheric electrical conductivity 50 Kushma Formation 263 Ladakh 208 Ladybird granite 566, 570, 571,572-573 Laya 428 Lesser Himalayan sequence (LHS) pressure-temperature diagram 417

617

provenance 174-176, 175, 177 structure 261-263,263 Lewis Thrust 562, 563 Lhasa Terrane 40 channel flow 59 lithospheric electrical conductivity 50 seismic reflectivity 54 strength-depth profiles 57 surface-wave velocity models 47 Lhotse detachment 363, 365, 367, 392, 395 Lingshi Dzong 428 Lobuche 362 Lukla 362, 365 Lumpola Valley seismic anisotropy 52 Lunpola Basin 356 Mabja Dome, Tibet 445-447,465 a~Aratitefission track analysis 455 /39mr thermochronology 451-456, 455, 456, 457-458, 460 cathodoluminescence (CL) images 454 cross-section 449 formation 461-464, 462 geological map 448 geology metamorphic history 450-451 rock units 447-450 structural history 450 Miocene crustal flow 464-465 regional tectonic map 446 significance of results 456-461,459, 461 U - P b geochronology 451,452, 453-454 Mackie pluton 571 Main Boundary Thrust (MBT) 40, 360 Main Central Thrust (MCT) 2, 25, 40, 362, 365 channel-bounding structures 8, 9, 10-11 crustal structure 397-399, 398, 399 exhumation 6 transport of materials 12 Main Himalayan Thrust (MHT) 360 Malashan Dome, Tibet 446, 471-473, 493 contradictory interpretations 491-492 general features 491 geological map 474, 477 granite bodies 481 bulk chemistry 485, 487, 488 Cuocu and Paika granites 485-490, 489, 490 origins 490-491 similarity to Kangmar granite 482-485, 484, 486 implications for channel flow models 492-493 metasedimentary schists D2 deformation kinematics 476-478, 479, 480 D2 structure orientation 476 deformation stages 473-476, 475 metamorphic zonation 473 sedimentary age 478-481,482 stereographic plot 478 tectonic map 472 Malton Complex deformation, metamorphism and age constraints 597 Malton gneiss 590 Marsyandi River 245, 271

618 Masang Kang 428, 429 Mica Creek 590 Milton Belt 519 Monashee Complex 544, 546, 548, 551,562, 590 deformation, metamorphism and age constraints 597 Monashee D~collement 551 Namche 362, 365 Namche Barwar 40 Nanga Parbat, Mount 40, 208 Nanga Parbat-Haramosh Massif crustal extrusion 208-210, 208 Nar valley, central Nepal 269-270, 289 exposed deformation features 285 geology GHS 272-274 Tethyan sedimentary sequence 274 structural constraints 274 crustal-scale buckling 277-278 Phu detachment 275-276 structural evolution 278-279 structural level-1 274, 276 structural level-2 275,276 structural level-3 277 timing constraints 4~ thermochronology 279-282, 282 b i o t i t e 4~ ages 284-285 h o r n b l e n d e 4~ ages 282, 283 muscovite 4~ 39Ar ages 282-284 U - P b geochronology 279, 280, 281 Nelson batholith 566 Nepal 269-270 geology of north-central region 270-272, 271 Newton window 519 normal-sense shear zones 439-441 geometry and kinematics 427-429 vorticity analysis 431-433,434, 441 north Himalayan gneiss (NHG) domes 176-179, 360 Northern Rongbuk valley 386, 391,392, 404, 405 Nuptse 363 Nyainqentanghla (NQTL) massif 240 Nyalam detachment, southern Tibet 351-352 analytical methods 334-335 40 Ar/-39Ar thermochronology calculated cooling rates and tectonic interpretations 345-348 fault footwall 340-342, 341-342, 349 fault surface 335-340, 336-337, 343-346, 347-348, 349 south-central GHS 342-345, 349 geological setting and structural framework 330, 331 deformational features 329 deformational features in footwall 329-331 Nyalam detachment, southern Tibet geological setting and structural framework tectonic setting 329 mechanical model 350-351,351 sampling characteristics and distribution 331 detachment fault footwall 331-333, 333-334 detachment fault surface 331,332, 3 3 3 - 3 3 4 south-central GMS 334 timing of detachment 348-350

INDEX Oknagan Fault System 544, 551 Oknagan-Eagle River Fault 590 Ominecca Belt 544, 607 channel flow model 593 Domain 1 - structures and metamorphism 595-600 Domain 1 - timing of deformation 600-601 Domain 2 - structures and metamorphism 601-602 Domain 2 - timing of deformation 602 Domain 3 - structures and metamorphism 602-603 Domain 3 - timing of deformation 603 -604 Domain 4 604 general statements 593-595 overview 597-599 tectonic interpretations 592 geological model 606 granitoid migration 604-605 consistency with other models 605-606 origin of Cretaceous granitoids 605 origin of Middle-Late Jurassic granitoids 605 origin of Palaeocene-Eocene granitoids 605 shortening 606-607 Paiku Lake 474, 477 Parmon Window 498, 501, 502 micrographs 505 quartz c-axis fragments 504 vorticity analysis 507 kinematic vorticity analysis 509 Paro 428 Passmore Dome 562, 570 Peloponese 498 Phaplu augen gneiss 365 Pheriche 362 Phu detachment 271, 273, 275-276, 277 Pine Mountain window 519 Pinnacles Area 590 deformation, metamorphism and age constraints 598 plateau formation 7 Poiseuille flow 27-33, 28, 41 velocity profiles 30 Porcupine Creek Anticlinorium 590 progressive deformation constant thickness 192-193, 192 Puga 40 pulsed channel flow 415-417, 420-421 Bhutan data 418-420, 420 Pumori 362 Punaka 428 Purcell Anticlinorium 562, 563, 590 Purcell Thrust 542, 544, 588 Purcell Thrust Fault 551 Qaidam Basin 40, 356 Qaidam Border fault 40 Qiangtang Terrane 40, 240 channel flow 59 lithospheric electrical conductivity 50

INDEX seismic reflectivity 54 strength-depth profiles 57 surface-wave velocity models 47 Qomolangma detachment 363, 365, 367, 369, 392, 395 Ram Tso 428 Rayleigh number 314 Red River fault 356 Renbu Zedong Thrust 360 Rocky Mountain Thrust 562 Rocky Mountain Trench 590 Rodophu Valley 429 Rongbuk glacier 395 Rongbuk Monastery 386, 391,392-394, 405, 406 Rongbuk valley 389-395, 391 Rundle Thrust 562 Sauratown mountains window 519 Scrip Nappe 590 Selkirk allochthon 542, 590 Selkirk Detachment Fault 546, 551 Selkirk Fan 590 deformation, metamorphism and age constraints 599 Selkirk Fan Axis 551,590 Selkirk Mountains 546 Shuswap Metamorphic Complex 590 Sichuan Basin 40 Siwalik hills 360 Slocan Lake Fault 566, 590 Smith River allochthon 519 Songpan Ganze 356 Songpan-Ganzi Terrane 40 channel flow 59 lithospheric electrical conductivity 50 South Tibetan detachment (STD) 2, 25, 310, 351-352, 360, 365, 367 analytical methods 334-335 4~ thermochronology calculated cooling rates and tectonic interpretations 345-348 fault footwall 340-342, 341-342, 349 fault surface 335-340, 336-337, 343-346, 347-348, 349 south-central GHS 342-345, 349 as passive roof fault to extruding channel 366-370 channel-bounding structures 8, 9, 10-11 exhumation 6 geological constraints 312- 313 geological setting and structural framework of Nyalam detachment 330, 331 deformational features 329 reformational features in footwall 329-331 tectonic setting 329 implications for Miocene denudation of Himalaya 323-324 kinematic inversions 8 lithospheric electrical conductivity 50 mechanical model 350-351,351 numerical model 313-315,313 numerical model results geothermal gradients 319-323,319, 320, 322 slip and erosion rates 315-319, 315, 316, 317

619

protoliths 12-13 sampling characteristics and distribution 331 detachment fault footwall 331-333, 333-334, 331,332, 333-334 south-central GMS 334 tectonic diagram 328 thermochronological constraints on cooling and exhumation 327-329 timescales of metamorphism, melting and channel flow 370-371,370 timing of Nyalam detachment 348-350 transport of materials 12 Southern Rocky Mountain Trench 546 Tamji 428, 429 Tarmim Basin 40, 356 Taurides 498 Taygetos Window 498, 500, 501 micrographs 505 quartz c-axis fragments 503 vorticity analysis 507 kinematic vorticity analysis 509 Tethyan Himalaya 40 see also Himalayan orogen channel flow 59 lithospheric electrical conductivity 50 seismic reflectivity 54 strength-depth profiles 57 surface-wave velocity models 47 Theri Kang 428 Thimphu 428, 429 Thor Odin Dome 544, 562, 590 thrust ramp models 246 Thyanboche 362 Tibet, crustal flow 62 evaluation of existing channel-flow models northern/eastern Tibet 61-62 southern Tibet 60-61 geophysical data data sources 43-44 geothermal measurements 44-45 high conductivity zones 49-51, 50 seismic anisotropy 51-53, 52 seismic attenuation 48-49 seismic reflectivity 53-55, 54 seismic velocity 45-48, 47 seismicity cut-off 45 temperature measurements 55-56 topography and gravity 44 lateral strain variations, vertical strain and strength profiles 39-41 physical properties inferred from geophysical observations vertical strength profiles 56-60, 57, 59 Tibet exhumation of Greater Himalayan rock 255-256, 256, 263-264, 289 tinting of extrusion 259-261,261 shallow anatexis model INDEPTH survey 241 supporting evidence 241-242, 242 Zhao and Morgan hypothesis 241 shallow anatexis model, critique of 246 consistency with 3He/4He data 243-244

620 Tibet (Continued) Shallow anatexis model, critique of (Continued) consistency with absence of Gangdese zircons 244-245 consistency with cold southern Tibetan Moho 243 consistency with geology and geochemistry of Yangbajain rift 244 crustal thickening explanation 244 nature of 'bright spots' 243 representativeness of Yadong-Gulu rift 242-243 stratigraphy 245-246, 245 Toma La 428, 429 Trail pluton 570 Tso Morari complex 40 Tugaloo terrane 519, 520 Ulug Muztagh 356 upper-mantle-scale (UMS) modelling 93, 126-128 density structure and isostatic compensation 98-99 design advantages and limitations 138 complexity and tuning 138-139 Himalayan-Tibetan (HT) models 140-141 model-data comparisons 142-143 philosophy of numerical approach 138 scaling of laboratory power-law creep flow laws 139-140 experiment description 128-129, 128 mechanical models 96-98 melt-weakening 98 parameters 95-96, 97-98, 99

INDEX results models LHO-LS1 and LHO-LS2 129-134, 130-133 thermal model 99 velocity boundary conditions and reference frames 93-96 Valhalla Complex 562, 590 cross-sections 563, 567 deformation, metamorphism and age constraints 597 geological map 566 origin 580 reconciliation of crystalline thrust sheet model 581-582 upper margin 572-573 Valkyr shear zone 566, 572-573, 590 Vidler Ridge 590 viscosity 7 viscous channel flow 4 vorticity, kinematic see kinematic vorticity Wagye La 428 Wing Pond Shear Zone (WPSZ), New Foundland crustal extrusion 210-214, 211,212, 213 Xianshui-he fault 356 Yangbajain graben 40 lithospheric electrical conductivity 50 Yarlung Tsangpo suture 360 Zanskar shear zone (ZSZ) 359

Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones Edited by R. D. Law, M. P. Searle and L. Godin

This collection of 27 review and research papers provides an overview of the geodynamic concepts of channel flow and ductile extrusion in continental collision zones. The focal point for this volume is the proposal that the middle or lower crust acts as a ductile, partially molten channel flowing out .~;;~ from beneath areas of over-thickened crust, such as the .-~'~..;~ ~, ..."~.~ Tibetan plateau, towards the topographic surface at plateau margins. This controversial proposal explains many features related to the geodynamic evolution of the plateau and, for example, extrusion and exhumation of the crystalline core of the Himalayan mountain chain to the south. In this volume -'.~ thermal-mechanical models for channel flow, extrusion and exhumation are presented, and geological and geophysical evidence both for and against the applicability of such models to the Himalayan-Tibetan Plateau system, as well as older continental collision zones Such as the Hellenides, the Appalachians and the Canadian Cordillera, are discussed.

_,~

Visit our online bookshop: http://wvvw.geolsoc.org.ul(Jbookshop Geological Society w e b site: http:llvvvvvv.geolsoc.org.uk

ISBN 1-86E39-E09-9

Cover illustration: Massive pale-coloured Miocene leucogranite intruded into dark sillimanite gneisses,south face of Makatu (8485 m) from the Barun glacier, Nepal Photograph by Mike Seafle

UU[I[![IIIU!!!![III