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

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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~
~

~

~.~

~

~

~-.~ = ~ o.~ ~ ,~ ~ ,.~ ,.~ v

~

.,~

,-~

. ~

o~.~ ~,~

c.~

",~

E- ~

~

"'~

~

~

.,~

o

= ~:~ ~ ~ Or.~

. . ~ ~

~.-~ . ~

,-~

~ ~

.,~

10

L. GODIN E T A L .

Table 1. Compilation o f interpreted ages o f f a u l t motion on the M a i n Central thrust system and South Tibetan d e t a c h m e n t system, based on geochronological data Structure 1

Location 2

Age 3

Minerals 4

System s

Reference 6

Rb-Sr Th-Pb, Ar Ar, U-Pb U-Pb, Ar U-Pb, Ar Ar Rb-Sr Th-Pb U-Pb Th-Pb, Ar Th-Pb Ar At, FT Ar

4. 5. 6. 7. 8. 9. 10. 11. 12. 13. 14.

Dbzes et al. (1999) Searle et al. (1999) Vance et al. (1998) Inger (1998) Harrison et al. (1997a) Noble & Searle (1995) Vannay et al. (2004) Catlos et at. (2002) DeCelles et al. (2001) Vannay et al. (2004) DeCelles et al. (2001) Godin et al. (2006) Searle & Godin (2003) Hurtado et al. (2001) Searle et al. (1997) Hodges et al. (1996) Vannay & Hodges (1996) Guillot et al. (1994) Godin et al. (2001) Harrison et al. (1999b) Coleman (1998), Coleman & Hodges (1998) Hodges et al. (1996) Guillot et al. (1994) Copeland et al. (1990) Kohn et al. (2004) Johnson et al. (2001) Coleman (1998) Johnson & Rogers (1997) Hodges et al. (1996) Vannay & Hodges (1996) Nazarchuk (1993) Macfarlane (1993) Kohn et al. (2004) Catlos et al. (2001) Johnson & Rogers (1997) Harrison et al. (1997b) Macfarlane et al. ( 1992, Macfarlane (1993)

Western H i m a l a y a

STDu STD~ STDI

Zanskar Sutlej Zanskar

< 18 Ma to 16 Ma 23 Ma to 17 Ma 23 Ma to 20 Ma

STDI STDI STD1 STDI STD1 STD1 MCTu MCTu MCTu MCT1 MCT1

Zanskar Garhwal Zanskar Zanskar Garhwal Zanskar Sutlej Garhwal West Nepal Sutlej West Nepal

c. 22.2 Ma to 19.8 Ma 23 to 21 Ma c. 23 Ma to 20 Ma c. 26 Ma to 18 Ma <21.9 Ma 21 Ma to 19.5 Ma 23 Ma to 17 Ma c. 5.9 Ma 22 Ma to 15 Ma c. 6 to 0.7 Ma 15 Ma to 10 Ma

Ms, Bt Mz, Ms Ms, Bt, Xe, Mz, Zr Mz, Ms Mz, Ms Ms, Bt Ms, Bt Mz Mz Mz, Ms Mz Ms Ms, Ap, Zr Ms

1. Inger (1998) 2. Vannay et al. (2004) 3. Walker et al. (1999)

Central- W e s t H i m a l a y a

STDu STDu STDu STDu STDu STDu

Nar Nar Kali Gandaki Langtang Annapurna Kali Gandaki

19 Ma to 16 Ma < 19 Ma < 17.2 ka < 17.3 Ma c. 18.5 Ma 15 Ma to 13 Ma

Ms, Bt, Hbl Mz Terr Mz, Xe Zr Ms

Ar Th-Pb 14C U-Pb U-Pb Ar

15. 16. 17. 18. 19. 20.

STDL, STD1 STD1 STD1

Manaslu Kali Gandaki Manaslu Marsyandi

19 Ma c. 22.5 >22.9 22 Ma

Bt, Ms Mz Mz Mz

Ar U-Pb Th-Pb U-Pb

21. 22. 23. 24.

STD~ STD1 STD1 MCTu MCTu MCTu MCTu

Annapurna Manaslu Manaslu Langtang Kathmandu Marsyandi Kathmandu

22.5 Ma to 18.5 Ma > 2 2 Ma > 2 0 Ma 16 Ma to 13 Ma 22 Ma to 14 Ma 22 to 18 Ma 21 Ma to 14 Ma

Zr, Mz, Xy Hbl Ms Mz Mz, Zr Mz Ms, Bt

U-Pb Ar Ar Th-Pb U-Pb U-Pb Rb-Sr

25. 26. 27. 28. 29. 30. 31.

MCTu MCTu

Annapurna Kali Gandaki

c. 22.5 Ma > 15 Ma

Mz, Zr Ms

U-Pb Ar

32. 33.

MCTu MCTu MCTl MCT1 MCT1

Kali Gandaki Langtang Langtang Marsyandi Kathmandu

c. 22 Ma >5.8 Ma c. 9 Ma c. 13.3 Ma c. 17.5 Ma

Mz, Th Ms Ms Mz Ms, Bt

U-Pb Ar Ar U-Pb Rb-Sr

34. 35. 36. 37. 38.

MCT1 MCT1

Marsyandi Langtang

c. 16 Ma < 9 to 7 Ma; c. 2.3 Ma

Mz Ms

Th-Pb Ar

39. 40.

41. Searle et al. (2003) 42. Murphy & Harrison (1999) 43. Hodges et al. (1998) 44. Hodges et al. (1992)

to 16 Ma Ma Ma to 18 Ma

Central-East Himalaya STDu STDu

Everest Everest

< 16 Ma c. 17 Ma

Mz Mz

U-Pb Th-Pb

STDu STDu

Everest Everest

c. 16 Ma 22 Ma to 19 Ma

Xe, Mz, Zr Ti, Xe, Hbl

U-Pb U-Pb, Ar

(Continued)

INTRODUCTION

11

Table 1. C o n t i n u e d Structure 1

Location 2

Age 3

Minerals 4

System 5

STDu STD1 STDI STD1 STDu MCTu MCTu MCTu

Nyalam Everest Everest Everest Makalu Everest Dudh Kosi Everest

< 16.8 Ma c. 21u2 Ma 18 Ma to 17 Ma <20.5 Ma <21.9 Ma c. 21u2 Ma c. 25 to 23 Ma c. 21 Ma

Mz Mz, Mz, Mz, Mz Mz, Mz Hbl

MCTu

Everest

23 Ma to 20 Ma

Hbl, Bt

Ar

Mz, Zr Mz

Th-Pb Th-Pb

Xe Xe Xe, U Xe

U-Pb U-Pb U-Pb U-Pb U-Pb U-Pb Tb-Pb Ar

Reference 6 45. 46. 47. 48. 49. 50. 51. 52.

Sch~irer et al. (1986) Viskupic et al. (2005) Searle et al. (2003) Simpson et al. (2000) Sch~irer (1984) Viskupic et aL (2005) Catlos et al. (2002) Hubbard & Harrison (1989) 53. Hubbard (1989)

Eastern Himalaya STDu STDu

Sikkim c. 1415 Ma Khula Kangri < 12.5 Ma

STD1 STD1 STD1 MCTu MCTu MCTu MCTu MCT, MCT1 MCT1

Wagye La Sikkim Sikkim Sikkim Sikkim Bhutan Bhutan Bhutan Sikkim Bhutan

c. 12 Ma Mz 23 Ma to 16 Ma Grt c. 17 Ma Mz, Zr c. 22 Ma Mz 23 Ma to 16 Ma Grt c. 22 Ma; 18 Ma to 13 Ma Mz, Xe c. 113.5 Ma Mz > 14 Ma Ms 15 Ma to 10 Ma Mz < 11 Ma Ms

U-Pb Sm-Nd Th-Pb Th-Pb Sm-Nd U-Pb U-Pb Ar Th-Pb Ar

54. Catlos et al. (2004) 55. Edwards & Harrison (1997) 56. Wu et al. (1998) 57. Harris et al. (2004) 58. Catlos et al. (2004) 59. Carlos et al. (2004) 60. Harris et al. (2004) 61. Daniel et al. (2003) 62. Grujic et al. (2002) 63. StiJwe & Foster (2001) 64. Catlos et al. (2004) 65. Sttiwe & Foster (2001)

1MCT1,Lower MCT; (and/or) mostlybrittle, post-metamorphic;local names includeMCT-1, Ramgarh,Munsiari. MCTu,Upper MCT; (and/or) ductile, symnetamorphic,synmagmatic;local names include MCT-2, Vaikrita, Mahabharat, Chomrong. STDj, Lower STD; (and/or) ductile, synmetamorphic,synmagmatic;local names include Zanskar, Sangla, Annapurna, Deurali, Chame, Lhotse, Zherger La. STDu,Upper STD; (and/or) mostly brittle, post-metamorphic;local names include Jhala, Macchupuchare,Phu, Qomolangma 2See Figure 1 for location. 3Compilation of direct geochronologicalresults only; includes age constraints based on cross-cutting structures/intrusionrelationships. 4Ap, apatite; Bt, biotite; Grt, garnet; Hbl, hornblende;Ms, muscovite;Mz, monazite;Terr, terraces; Th, thorite; Ti, titanite; U, uraninite; Xe, xenotime; Zr, zircon. 5Ar, 4~ thermochronology;14C,carbon 14; FT; fission track geochronology;Rb-Sr, Sm-Nd,Th-Pb: thorium-leadion microprobe (2~ age); U-Pb, U-(Th)-Pb geochronology. 6Reference numbers refer to respective 'age range bar' on Figure 4.

GHS, with top-to-the-north shearing only appearing in the top-most part of the GHS, near and within the STD system. The location of this transition in shear sense has yet to be documented. Field and microstructural data indicate that this pervasive ductile deformation is characterized by heterogeneous general non-coaxial flow (components of both simple and pure shear) rather than by ideal simple shear. For example, quartz petrofabric data (Boullier & Bouchez 1978; Brunel 1980, 1983; Bouchez & P~cher 1981; Burg e t al. 1984; Greco 1989; Grujic et al. 1996; Grasemann et al. 1999; Bhattacharya & Weber 2004; Law et al. 2004) consistently indicate a component of pure shear. Williams et al. (2006), however, present an opposite interpretation of these data based on strain compatibility and mechanics theory. Quantitative vorticity analyses within the GHS document a progressively increasing component of simple shear traced upward towards the STD and overlying sheared Tethyan sedimentary rocks

(Law et al. 2004; Jessup e t al. 2006), a general shear deformation within the core of the GHS (Carosi et aL 1999a, b, 2006; Grujic e t al. 2002; Law et al. 2004; Vannay et al. 2004; see also Fig. 3C), and an increasing pure shear component traced downward towards the underlying MCT zone (Grasemann et al. 1999; Jessup et al. 2006; but cf. Bhattacharya & Weber 2004). Macro- and microstructural fabric data (especially conjugate shear bands, porphyroclast inclusion trails, and crenulation cleavage at various stages of development) also suggest a strong component of shortening across the foliation in addition to foliation-parallel shearing (e.g. Carosi e t a l . 1999a, b, 2006; Grujic et al. 2002; Law et al. 2004; Vannay et al. 2004). The structural data indicate that ductile deformation is pervasively distributed through the entire GHS, in the top part of the Lesser Himalayan sequence, and at the base of the Tethyan sedimentary sequence. A direct implication of the general flow model is that the bounding surfaces of the

12

L. GODIN ET AL.

crystalline core (i.e. MCT and STD shear zones) must therefore be 'stretching faults' (Means 1989) accommodating transport-parallel pervasive stretching of the crystalline core during internal flow (Grasemann et al. 1999; Vannay & Grasemann 2001; Law et al. 2004). Metamorphic

characteristics

One of the most intriguing phenomena of the Himalaya is the inverted metamorphic sequence present in both the Lesser Himalayan sequence and GHS (see reviews by Hodges 2000). At the top of the GHS and at the base of the Tethyan sedimentary sequence, a strongly attenuated, rightway-up decrease in metamorphic grade is present. Models for inverted metamorphism include: (1) overthrusting of hot material ('hot iron effect'; Le Fort 1975); (2) imbricate thrusting (Brunel & Kienast 1986; Harrison et al. 1997b, 1998, 1999a); (3) folding of isograds (Searle & Rex 1989); (4) transposition of a normally zoned metamorphic sequence due to either localized simple shear along the base of the GHS (Jain & Manickavasagam 1993; Hubbard 1996), heterogeneous simple shear distributed across the Lesser Himalayan sequence and GHS (Grujic et al. 1996; Jamieson et al. 1996; Searle et al. 1999) or general shear of previously foreland-dipping isograds (Vannay & Grasemann 2001); and (5) shear heating (England et al. 1992; Harrison et al. 1998; Catlos et al. 2004). The metamorphic isograds can be deformed passively according to various kinematic models that are compatible with either extrusion or channel flow, or both (e.g. Searle et al. 1988, 1999; Searle & Rex 1989; Jain & Manickavasagam 1993; Grujic et al. 1996, 2002; Hubbard 1996; Jamieson et al. 1996; Davidson et al. 1997; Daniel et al. 2003). Coupled thermal mechanical finite element modelling (Jamieson et al. 2004) has been successful in replicating the distribution of the metamorphic isograds and P-T-t data obtained through field and laboratory studies, although it failed to predict the timing of the low temperature metamorphic overprint. Other models propose a specific style of thrusting along the base of the GHS as an alternative model to explain both the distribution of metamorphic zones and the timing of metamorphism (e.g. Harrison et al. 1998; Catlos et al. 2004). Lateral versus vertical transport of material

Integration of geobarometry and thermochronology can deduce the amount and timing of exhumation: more specifically, the rate of vertical displacement of rocks within the crust. Only the vertical

component of exhumation can be estimated using these techniques. Along low-angle shear zones like the STD and MCT, however, the horizontal component of displacement is predominant. Some investigations use the jump in pressures, estimated by metamorphic assemblages across the STD, to estimate the horizontal component of displacement (Searle et al. 2002, 2003). Displacement estimates based on temperatures inferred from metamorphic assemblages, however, involve assumptions about the shape of the isotherms, which may change during the exhumation process. Simplified restoration of the GHS (e.g. INDEPTH data; Nelson et al. 1996; Hauck et al. 1998) indicate that the GHS may extend down-dip "for at least 200 kin, and possibly up to 400 km (Grujic et al. 2002). Exhumation from mid-crustal levels at 35-40 km (as suggested by pressures at peak T; see Hodges (2000) for summary of data, and Hollister & Grujic (2006) for interpretation) indicates that lateral displacement rates in the GHS are five to ten times larger than the vertical displacement rates. These values ought to be compared with inferred surface denudation rates (e.g. Thiede et al. 2004; Vannay et al. 2004; Grujic et al. 2005), and estimation of shortening or displacements across the MCT and STD. Conventional cross-section (usually line-length) restoration techniques are used to estimate these values (e.g. Schelling & Arita 1991; DeCelles et al. 2002; Searle et al. 2003). However, if deformation is pervasive through the GHS and there is an inversion of the displacement along the STD, no single value can fully describe the displacement along the shear zone. Displacements across the GHS relative to the Lesser Himalayan sequence are also expected to progressively increase towards the core, and progressively decrease upward towards the STD, which is compatible with the calculations of particle displacement paths for various points within a model GHS (Jamieson et al. 2004, 2006). Discontinuity of protoliths across the c h a n n e l

According to the above discussion, the largest rate of particle displacement change occurs across the STD and MCT (e.g. Davidson et al. 1997). Most detrital zircon and isotopic studies suggest that the Lesser Himalayan sequence and GHS metasediments may have different protolith ages. Zircon and Nd model ages and the eNd values suggest a Late Archean to Palaeoproterozoic source for the metasediments of the Lesser Himalayan sequence versus a Meso- to Neoproterozoic source for

INTRODUCTION the GHS (Parrish & Hodges 1996; Whittington et al. 1999; Ahmad et al. 2000; DeCelles et al. 2000, 2004; Miller et al. 2001; Robinson et al. 2001; Argles et al. 2003; Martin et al. 2005; Richards et al. 2005; but cf. Myrow et al. 2003), although structural restoration suggests otherwise (e.g. Walker et al. 2001). The lithotectonic units, separated by the first-order shear zones, may have distinct palaeo-geographic origins; however, this does not necessarily mean they belong to different tectonic plates. Similar results are obtained by numerical modelling and particle tracking (Jamieson et al. 2006), which suggest that from the base to the top of the GHS, the protoliths should have a progressively more distal origin (with respect to the pre-collision plate margin), while the opposite situation is predicted for the Lesser Himalayan sequence. Although different protolith origins for the GHS and the Lesser Himalayan sequence might exist, a similar interpretation cannot be applied to the GHS and the Tethyan sedimentary sequence. Channel flow models predict that the STD should be the locus for large relative particle displacement, implying a different origin for the GHS and the Tethyan sedimentary sequence (Jamieson et al. 2006). Recent structural restorations and isotopic studies, however, propose the lower Tethyan sedimentary sequence as a potential protolith for some of the GHS (Vannay & Grasemann 2001; Argles et al. 2003; Gehrels et al. 2003; Searle & Godin 2003; Gleeson & Godin 2006; Richards et al. 2005). The STD is generally interpreted as either a d6collement surface (stretching fault), where the thick pile of continental margin rocks (Tethyan sedimentary sequence) has been decoupled without much internal disturbance to the stratigraphy, or a passive roof thrust within the MCT system, with a hanging-wall flat-footwall flat geometry (Searle et al. 1988; Yin 2002). The GHS is dominated by three lithologic units, which maintain their respective structural positions for over a thousand kilometres along-strike (Gansser 1964; Le Fort 1975). Recent detailed mapping across the GHS locally reveals a more complex distribution of, and variation within, these units (Searle & Godin 2003; Searle et al. 2003; Gleeson & Godin 2006). Nonetheless, the firstorder lateral continuity of the GHS units indicates an apparent lack of internal stratigraphic disturbance. This has been highlighted as a possible pitfall for the channel flow model (Harrison 2006). Model results indicate, however, that the channel may very well maintain internal 'stratigraphy', as long as the deformation is concentrated along the boundaries and flow is planar along the length of the channel (Jamieson et al. 2006).

Timing of melting

13 and shortening

structures

The channel flow model assumes that melts (leucosomes and granites) will substantially reduce the viscosity of a crustal layer (i.e. channel). It also predicts that these melts should be younger than shortening structures found in the upper plate-shortening structures that would have created the necessary crustal thickening and ensuing heating to partially melt and lower the viscosity of the underlying mid-crust. The Tethyan sedimentary sequence is the upper plate in the Himalaya. Leucosome and leucogranite bodies occur within all units of the GHS (Dietrich & Gansser 1981" Le Fort et al. 1987; Burchfiel et al. 1992; Guillot et al. 1993; Hodges et al. 1996; Hollister & Grujic 2006). Most U - T h - P b ages for the melts in the central Himalaya range from 23-22 Ma (Harrison et al. 1995; Hodges et al. 1996; Coleman 1998; Searle et al. 1999; Godin et al. 2001; Daniel et al. 2003; Harris et al. 2004) to 13-12 Ma (Edwards & Harrison 1997; Wu et al. 1998; Zhang et al. 2004). However, evidence for leucosome melt production during the Oligocene also exists (Coleman 1998; Thimm et al. 1999; Godin et al. 2001). North Himalayan granites found in southern Tibet range in crystallization age between 28 Ma and 9 Ma (Sch~irer et al. 1986; Harrison et al. 1997a; Zhang et al. 2004; Aoya et al. 2005). Syntectonic (synchannel?) granites yield ages of 23.1 _+ 0.8 (Lee et al. 2006). Some North Hinaalayan granites, however, yield zircon and monazite crystallization ages of 14.2 + 0.2 Ma and 14.5 + 0.1 Ma, respectively, indicating that vertical thinning and subhorizontal stretching had ceased by the middle Miocene (Aoya et al. 2005; Lee et al. 2006). Several phases of deformation are recorded by the overlying Tethyan sedimentary sequence (Steck et al. 1993; Wiesmayr & Grasemann 2002; Godin 2003). Although the absolute age(s) of the dominant shortening structures is disputed, most authors agree that significant thickening of the Tethyan sedimentary sequence occurred prior to the Miocene, most likely in the Oligocene or even before (Hodges et al. 1996; Vannay & Hodges 1996; Godin et al. 1999b, 2001; Wiesmayr & Grasemann 2002; Godin 2003; Searle & Godin 2003). Some of these shortening features are interpreted to be coeval with high-pressure metamorphism in the GHS (Eohimalayan phase; Hodges 2000), associated with early burial of the GHS beneath a thickening overlying Tethyan sedimentary sequence (Godin et al. 1999b, 2001; Godin 2003).

14 Channel

L. GODIN ETAL. thickness and late-stage

modifications

During periods of active channel flow, models predict that the channel should be 10 to 20 km thick (Royden et al. 1997; Clark & Royden 2000; Beaumont et al. 2004; Jamieson et al. 2004, 2006). The structural thickness of the GHS varies considerably, from 2 - 3 km in the Annapurna area (Searle & Godin 2003; Godin et al. 2006), up to 30 km in the Everest area (Searle et al. 2003, 2006; Jessup et al. 2006), and even more in the Bhutan Himalaya (Grujic et al. 2002). Substantial post-channel, post-extrusion modifications have altered the original geometry of the channel. Outof-sequence thrusts such as the Kakhtang thrust or the Kalopani shear zone (Grujic et al. 1996, 2002; Vannay & Hodges 1996), and large amplitude folding of the GHS (Johnson et al. 2001; Gleeson & Godin 2006; Godin et al. 2006) may account for some of the observed thickness variation. Alternatively, various instabilities and failure of the upper crust may induce local accretion of channel material and sequential development of domes (i.e. spatio-temporal variations of channel thickness), both along-strike and down-dip (e.g. Beaumont et al. 2004). Some out-of-sequence thrusting may be the result of such pulsed channel flow and related doming and extrusion (Grujic et al. 2004; Hollister & Grujic 2006). In the Kali Gandaki (Annapurna area) and eastern Bhutan, the GHS is as thin as 3 kin, while preserving its typical apparent internal 'stratigraphy' and metamorphic zoning. Whether this is a reflection of a lateral variation in the component of coaxial (pure shear) deformation and thinning during the channelling and/or extrusion phase remains unclear. Thin segments of the GHS may represent the most proximal parts of the channel, while the thicker segments are more distal parts of the palaeo-channel. In this alternate interpretation, variation in thickness of the GHS at the presentday topographic front could reflect along-strike variation in the foreland-directed advance of a channel flow regime.

Challenges and unresolved issues The challenge to testing the applicability of the channel flow model in the Himalaya-Tibet system lies within the Earth scientist's ability to accurately interpret deformation paths and palaeo-isothermal structures recorded by exhumed metamorphic rocks that exhibit finite strain and metamorphic field gradients. Limited subsurface geophysical coverage of the Himalaya-Tibet system makes correlation of surficial data with a putative channel at mid-crustal depths tentative.

The INDEPTH program laid the foundation for imaging critical mid- to lower crustal features that the channel flow hypothesis relies on. Unfortunately, data for the Himalaya are limited to a single transect. Higher resolution and more extensive seismic surveys may help resolve longstanding criticism of the channel flow models. For example, Harrison (2006) highlights the risk of generalizing the 'bright spots' low velocity zones to the entire southern Tibetan Plateau, because the seismic line was run within an active graben, where intrusions might be locally controlled by extension, and not crustal-scale melting. Assessing the applicability of the model to older, more deeply exhumed orogens will also prove to be challenging because of overprinting deformation and thermal events common to these systems. In older orogenic systems, probably the most important limitation to the application of the channel flow concept is the absence of control on palaeo-horizontal and palaeo-vertical directions. The channel flow model in the Himalaya-Tibet system is based on the concept of a lateral lithostatic pressure gradient acting as the driver for flow. Lateral variations in crustal thicknesses at the time of orogenesis in older systems are simply unknown. Furthermore, as pointed out by Jones et al. (2006), many of the structural/kinematic indicators of a channelized flow could also be compatible with tectonically rather than gravitationally driven systems (e.g. transpressional strike-slip tectonics). Potentially the most important limitation to testing the channel flow model is the fact that many of the initial parameters used in numerical models such as those presented by Beaumont et al. (2001, 2004) are taken directly from field and geochronological data collected in the Himalaya-Tibet system. Some participants at the Burlington House conference argued that it therefore follows that assessing the channel flow model, and testing its applicability against field constraints, becomes an inherently circular argument. In contrast, other participants (e.g. D. Grujic pers. comm. 2004) argued that one of the major strengths of the rigorously constructed thermal-mechanical generic models is their ability to test a wide range of potentially important parameters, and thereby identify the combinations of parameters that produce results which most closely resemble a given orogenic system. Yet other participants argued that where model results are sensitive to a wide range of potential boundary conditions and input parameters, it is inherently difficult to determine which are the most important parameters. In contrast, Jones et al. (2006) have argued that choosing geologically realistic boundary conditions and input parameters for

INTRODUCTION thermal-mechanical models is a critically important first step in ensuring that models are wellcalibrated to a specific orogen. They further argue that 'tuning' a model to match a specific orogen should not be regarded as a weakness of the modelling method, but is the basis for a better understanding of which factors are likely to have the most influence on orogenic processes such as channel flow and crustal extrusion.

Concluding remarks The proposed channel flow model explains many features pertaining to the geodynamic evolution of the Himalaya-Tibetan Plateau system, as well as other older orogenic systems. It reconciles the apparent coeval nature of the MCT and STD faults and kinematic inversions at the top of the GHS, leading to southward extrusion and exhumation of the crystalline core of the Himalaya from beneath the Tibetan Plateau. In addition, it provides an alternative and quantitative explanation of the inverted metamorphic sequence at the orogen scale, and effectively couples the tectonic and surface processes. The proposal that the middle or lower crust acts as a ductile, partially molten channel flowing out from beneath areas of overthickened crust (such as the Tibetan Plateau) towards the topographic surface at the plateau margins remains controversial, however, both with respect to the Himalaya-Tibet system and particularly older, less well documented orogenic systems. The channel flow model nonetheless presents an exciting new conceptual framework for understanding the geodynamic evolution of crystalline cores of orogenic belts, and may become the source for a paradigm shift in continental tectonics studies. The following 26 papers in this Special Publication are arranged into four main groups. In the first group of papers this brief introduction to channel flow and ductile extrusion processes is paired with a more in-depth review by Grujic of channel flow processes associated with continental collisional tectonics. In the second group of papers detailed overviews are given by Klemperer and l-lodges of the geophysical and geological databases from which the concepts of channel flow and ductile extrusion in the HimalayaTibetan Plateau system originally developed. Different aspects of the modelling of channel flow and ductile extrusion processes are covered in the third group of papers. Coupled thermalmechanical finite element models are presented in papers by Beaumont et al., Medvedev & Beaumont and Jamieson et al., while the effects of volume change on orogenic extrusion are considered by Grasemann et aL In the last two papers in this group, problems associated with

15

identifying channel flow and ductile extrusion in older orogens are discussed by Jones et al., while linkages between flow at different crustal levels (infrastructure and suprastructure) and constraints on the efficiency of ductile extrusion processes are explored by Williams et al. The fourth and largest group of papers is composed of a series of predominantly field-based case studies providing geological constraints on channel flow and ductile extrusion as an orogenic process. This last group of papers is divided into subsections on the Himalaya-Tibetan Plateau system, the Hellenic and Appalachian orogenic belts, and the Canadian Cordillera. The Himalaya subsection begins with a wide-ranging critique by Harrison of the applicability of channel flow models to the Himalaya-Tibetan Plateau system. Subsequent papers in the Himalaya subsection focus dominantly on specific field areas within the Lesser and Greater Himalaya and are arranged in order of geographic location starting with western Nepal (Robinson & Pearson) and then progressing eastwards through the Annapurna region of central Nepal (Godin et al., Seaillet & Searle, Annen & Scaillet) and the Nyalam-Everest regions of Tibet and eastern Nepal (Wang et aL, Searle et al., Jessup et al.) to the Bhutan Himalaya (Hollister & Grujie, Carosi et aL). The Himalaya-Tibetan Plateau subsection concludes with papers by Lee et aL and Aoya et al. on gneiss domes exposed to the north of the Himalaya and their implications for mid-crustal flow beneath southern Tibet. Geological evidence for and against channel flow and ductile extrusion in older orogenic systems is discussed in the remaining two subsections of this volume. Xypolias & Kokkalas present integrated strain and vorticity data indicating ductile extrusion of mid-crustal quartz-rich units in the Hellenides of Greece, while Hatcher & Mersehat present field evidence in support of channel flow operating parallel to orogenic strike in the Appalachian Inner Piedmont, USA. Arguments for (Brown & Gibson, Kuiper et aL) and against (Carr & Simony) channel flow in the crystalline interior of the Canadian Cordillera are presented in the final three papers of the volume.

The authors wish to warmly thank all participants of the 2004 Burlington House conference for fruitful discussions during and after the meeting. We thank C. Beaumont, R. L. Brown, R. A. Jamieson, M. Jessup, K. Larson, S. Medvedev, R. A. Price and P. F. Williams for discussion, and K. Larson, M. Jessup, and Chief Editor J. P. Turner for their detailed reviews of earlier versions of this manuscript. D.G. acknowledges support from the Canadian Institute for Advanced Research (CIAR). L.G., R.D.L. and M.S. are funded by the Natural Sciences and

16

L. GODIN ET AL.

Engineering Research Council of Canada (NSERC), the National Science Foundation of the United States (NSF), and the National Environment Research Council of the United Kingdom (NERC), respectively.

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Channel flow and continental collision tectonics: an overview D. GRUJIC Department

of Earth Sciences, Dalhousie

University,

Halifax, Canada (e-mail: [email protected]) Abstract: The principle of channel flow as definedin fluid dynamics has been used in continental

geodynamics since the 1980s. The basic equations for one-dimensional flow introduced to geologists by Turcotte and Schubert were further developed by several research groups to meet the needs of specific studies. The most substantive differences among numerical models are results of different solutions for flow in crust, developed for different boundary conditions. The concept of channel flow has met with strong opposition and criticism from geophysicists and modellers. Although it is difficult to prove unambiguously that there is an active weak channel, it is still the most successful model to explain and predict the tectonics, metamorphism and exhumation of high-grade terranes in some orogens. Moreover, the concept of channel flow has stimulated novel approaches to the study of both the tectonics and metamorphism of large, hot orogens and the interaction between tectonic and surface processes.

The concept of channelized flow of a weak crustal layer has been applied to various tectonic settings (Godin et al. 2006): (a) asthenospheric counterflow; (b) lower crustal channels; (c) intra-crustal channels; (d) subduction channels; (e) salt tectonics. In this overview only the concept of intra-crustal channels in collision orogens will be discussed, while detailed discussion of the range of proposed geologically relevant solutions to Stoke's equation will be presented elsewhere. In active collision orogens the channel flow model has been used to explain the coupling between crust and mantle, strain in the crust, metamorphism, synorogenic exhumation of high-grade terranes and landscape evolution. Some numerical modelling strongly supports existence of a weak crustal layer (e.g. Royden 1996; Clark & Royden 2000; Beaumont et al. 2001a, b; 2004, 2006; Clark et al. 2005) while others oppose it (e.g. Toussaint & Burov 2004; Toussaint et al. 2004). Similarly, geophysical data can be interpreted both in favour of an active lower- or mid-crustal channel (Nelson et al. 1996; Klemperer 2006) or against it (e.g. Flesch et al. 2005; Hilley et al. 2005). In the models the details of channel flow are sensitive to the flow laws assumed to operate in the crust, the mechanical properties of bounding crustal layers and density distribution in the crust. A number of analytical solutions to Stoke's equation have been derived and experiments were developed for fundamentally different boundary conditions: conclusions will thus be model-dependent. The Himalaya-Tibet orogen is the most studied putative example of both active- and palaeochannel flow. The Greater Himalayan Sequence (GHS), also referred to as Higher Himalayan Crystallines, is a sequence of amphibolite- to

granulite-facies ortho- and paragneisses, migmatites and syntectonic leucogranites. The GHS is both underlain and overlain by greenschist and lower grade metasediments and forms the metamorphic core of the Himalaya (e.g. Hodges 2000, 2006; Jamieson et al. 2004, 2006). The GHS is bounded by the Main Central thrust (MCT) at the base and by the South Tibetan detachment (STD) at the top. The two shear zones are subparallel, have opposite senses of shear and have operated coevally over an extended period of time (Hodges 2000; Godin et al. 2006). The change from burial to exhumation of the GHS is reflected in the change in shear sense along the upper bounding shear zone. The earlier phase of deformation is dominated by thrust-sense shear zones while the return channel flow and ductile extrusion are characterized by normal-sense shear along the roof of the exhumed channel. The presence of melt during both south- and north-directed shearing at the base and top of the GHS respectively (Fig. 1; e.g. Grujic et al. 1996, 2002; Davidson et al. 1997; Daniel et al. 2003; Harris et al. 2004) shows that high-temperature metamorphism and anatexis were an integral part of the exhumation process of the GHS (Godin et al. 2006; Hollister & Grujic 2006). Two closely related terms, ductile extrusion and channel flow, are used in the literature in regard to tectonics of a weak crustal layer. Kinematic models for both are based on the concept of a pair of coeval subparallel dip-slip shear zones with opposite senses of shear bounding a crustal layer with significantly lower strength or viscosity than the bounding layers. The concept of channel flow is based on the physical laws of fluid dynamics with a particular set of boundary conditions. Ductile extrusion is more difficult to define

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, 25-37. 0305-8719/06/$15.00 9 The Geological Society of London 2006.

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D. GRUJIC

Fig. 1. Outcrop photographs of deformed migmatites from the metamorphic core of the Himalaya in central Bhutan (i.e. the Greater Himalayan Sequence; e.g. Grujic et al. 1996), suggesting an extended period of coeval partial melting and deformation. (a) A finely foliated migmatite has been folded into chevron-type folds. The axial planes are associated with a new generation of little-deformed, yet sheared leucosome (for discussion see Davidson et al. 1997). The shear planes are parallel to the dominant foliation associated with the top-to-the south thrusting. View to the east. (b) Folded stromatic migmatite with leucosome intruding along axial surfaces. Sample is c. 6 cm long. (c) Normal-sense, top-to-the north shear band with leucosome in the core of the shear band. (d) Leucosome in the pressure shadows around a garnet.

precisely and uniquely and will probably be of different forms in different orogens.

Ductile extrusion Two types of extrusion along dip-slip fault or shear zones have been proposed in the literature. First is the extrusion of a rigid crustal sliver or wedge between discrete faults (e.g. Chemenda et al. 1995; Northrup 1996). This model is probably equivalent to the 'expulsion' model of Hodges (1998, his fig. 3). The second type is ductile extrusion of a deformable crustal layer between ductile shear zones (e.g. ductile extrusion of the metamorphic core of the Himalaya; Grujic et al. 1996; Grasemann & Vannay 1999; Grasemann et al. 1999). A similar process has been described in a number of other orogens, e.g. the Norwegian Caledonides (Northrup 1996), the Monashe Complex of the Canadian

Rockies (Gibson et al. 1999; Williams & Jiang 2005) and the Hellenides (Xypolias & Koukouvelas 2001; Xypolias & Kokkalas 2006). In the former type strain is concentrated along the boundaries and there is little or no deformation in the core. In the latter, the deformation is pervasively distributed throughout the extruding wedge. In the case of pervasive deformation the strain may be simple shear with strain gradient toward the boundaries, or different degree of general shear (combination of simple and pure shear; e.g. Law et al. 2004; Jessup e t al. 2006). Even deformation of the wedge by pure shear only would require simple shear deformation along the boundaries (see also Williams et al. (2006) for an argument against wedge extrusion). When the extruding wedge is composed of materials with yield strength, the deformation of the wedge is probably best represented by the Prandtl cell model (Price 1972).

CHANNEL FLOW AND COLLISION TECTONICS Lateral extrusion of rigid crustal blocks between steep strike-slip faults as in SE Asia (e.g. Tapponnier et al. 1982, 2001) or the Eastern Alps (Ratschbacher et al. 1991) will not be discussed here. The driving forces of extrusion may be buoyancy forces (caused by different crustal densities), tectonic overpressure (caused by non-parallel boundaries of a weak crustal layer), pressure differences caused by varying crustal thickness (i.e. topography) or the combination of any of these. Both ductile extrusion and extrusion of a rigid wedge are coupled with surface erosion (e.g. Chemenda et al. 1995; Grujic et al. 1996; Hodges 1998). Extrusion may thus be regarded as an exhumation process. It is evident that the volume of the extruding wedge limits the extrusion process. If ductile extrusion is linked to a crustal channel, the viscous part of the channel that cools below a critical temperature is converted rheologically into an extruding wedge: extended flow of a crustal channel provides replenishment of the extruding wedge. In this context, extrusion may mean the pumping or forcing out of material at the end of the viscous channel. The extrusion zone is, then, the part of the channel near the open end through which channel material is exhumed toward the surface (Beaumont et al. 2004, 2006). This means that in active orogens the exhumed metamorphic core may represent a fossil channel while the active channel is in the interior of the orogen at mid- to lower crustal level and beneath the related continental plateau. An important point is that the margins of the cell/wedge now exposed at the topographic surface may be more related to the processes of late stage extrusion/exhumation, rather than original channel flow within the more distal parts of the orogenic system. The classic example here may be the Greater Himalayan Sequence of the Himalaya (Hodges 2000; Grujic et al. 2002; Searle et al. 2003, 2006; Searle & Szulc 2005; Godin et al. 2006) and the putative active channel beneath the south Tibetan Plateau (Nelson et al. 1996; Hodges et al. 2001; Klemperer 2006).

Channel flow Channel flow is a process in which viscous fluid flows through a channel lying between two rigid sheets that deform the viscous fluid between them through induced shear stress and pressure gradients within the fluid channel (e.g. Batchelor 2000; Turcotte & Schubert 2002). In geologically relevant analyses, boundary conditions that lead to channel flow solutions are diverse. In most cases the crustal plates bounding a weak crustal layer are rigid (i.e. the horizontal component of velocity is zero) but deformable (i.e. both the upper and

27

lower surface of the layer can independently move vertically). The bounding plates may be also ductile and allowed to fail. The horizontal velocity profile in the channel depends on the geometry of the channel, but the simplest qualitative characteristic of these models in the case of rigid and undeformable bounding plates is that the onedimensional (1D) velocity field is a hybrid between two end-member components: (1) Couette flow where the induced shear across the channel produces a uniform vorticity across the channel (Fig. 2a); and (2) Poiseuille flow (also known as the 'pipe-flow' effect) in which the induced pressure gradient produces highest velocities in the centre and opposite vorticity for the top and bottom of the channel (Fig. 2b). In the first case the bounding plates move parallel to each other; in the second case they are stationary. In order to avoid misunderstandings it is recommended that these end-members be referred to by their names and the term 'channel flow' be reserved for hybrid flow with specified boundary conditions. For the general channel flow the velocity profile will be a hybrid of the two endmembers (Fig. 2c). The channel flow takes place in an infinitely long and wide channel (therefore a layer) whose lower boundary moves with velocity Uo relative to its stationary upper boundary. The velocity profile of the Couette flow is a function of the displacement rate between the two bounding plates. For the simple Couette flow we assume that the velocity of the lower plate is Uo ~ 0 and the velocity of the upper plate is u = 0. Non-slip boundary conditions apply; a viscous fluid in contact with a solid boundary must have the same velocity as the boundary. If the applied pressure gradient is zero, p~ = P2 or d p / d x = 0, the solution is a linear velocity profile:

u = u0(1--h@h)

(1)

If there is no relative displacement between the bounding plates, i.e. Uo = 0, the velocity of Poiseuille flow is governed by the pressure gradient d p / d x , viscosity tZch, depth z, and channel thickness hch:

U

--

1

dp(z2

hchZ)

(2)

21Xch d_r

(Turcotte & Schubert 2002, equation 6-14). Pressure in the crust may be specified in three ways with increasing level of approximation: (a) real pressure (estimations of it are not obvious; however, some analyses of crustal flow use this approach, e.g. McKenzie et al. 2000; McKenzie & Jackson 2002); (b) lithostatic pressure in crust with density

28

D. GRUJIC variations (e.g. Bertotti et aL 2000; Lehner 2000); (c) lithostatic pressure in crust with uniform density. This third level approximation is the simplest case and is assumed in the following discussions about channel flow in the crust. The flux U of material in the channel can be expressed by integrating the velocity u of the material over the channel thickness:

hch

U =

I

(3)

u(z) dz

0

Applied to Equation 2 this leads to the flux U given by:

U--

h3h dp 12/xch dx

(4)

(Turcotte & Schubert 2002). For a linear viscous fluid Turcotte & Schubert (2002, equation 6 - 1 7 ) give the following equation for the mean velocity of flow (equal to flux/thickness) in a parallelsided channel: _

h~h dp 12/Xch dx

Fig. 2. Idealized picture of channel flow. Here the bounding plates of a weak viscous channel are depicted as rigid and undeformable; in nature the boundaries of a channel show smooth gradients in mechanical properties and both the weak channel and bounding crustal levels above and below are deforming. (a) Couette flow component. (b) Poiseuille flow component. (e) Return channel flow for a particular set of parameters. Symbols are discussed in the text. The vorticity w values (rotational component of the flow profile) are schematically indicated by the width of the black bar, with highest vorticities (simple shear dominant) indicated by a wide bar segment (see also Jessup et al. 2006). Only the absolute value of the vorticity oJ is indicated regardless of whether it is positive (sinistral simple shear) or negative (dextral simple shear) (after England & Holland 1979; Mancktelow 1995; Grujic et al. 2002). ~oc:vorticity in pure Couette flow. Wp: vorticity in pure Poiseuille flow. oJ=:vorticity in a hybrid channel flow.

uo

+5-

(5)

Equation 5 accounts for both Couette flow and Poiseuille flow (for the latter if Uo = 0). In a crust with uniform densities the principal factors that influence channel flow are therefore the relative velocity of the bounding plates, the thickness of the channel, the viscosity of the channel material and the pressure gradient along the channel. Which of the end-members will be dominant in the hybrid velocity profile is a function of threshold value for a particular parameter (e.g. England & Holland 1979; Mancktelow 1995). For a constant plate convergence rate, pressure gradient and channel thickness, lowering of channel viscosity below a threshold value will shift the uniform Couette flow to heterogeneous flow where one part of the channel will be dominated by Poiseuille flow (Fig. 2c). A similar effect may be achieved with constant viscosity and channel thickness by slowing the plate convergence rate, or increasing the pressure difference (e.g. surface relief). The pressure gradient or pressure difference along the channel may be provided by non-parallel channel walls as a tectonic overpressure (e.g. Mancktelow 1995) or by a different lithostatic load (which in most active orogens is related to topography). It is important to note here that in the case of a non-horizontal channel, when pressures are close to lithostatic and crustal densities are uniform (e.g. in models by Beaumont et al. 2004; Jamieson et al. 2004), the pressure gradient

CHANNEL FLOW AND COLLISION TECTONICS is approximately proportional to variations in the burial depth of the channel (e.g. Gemmer et al. 2004); this does not correspond to topography (i.e. surface relief) Ah because the channel is not horizontal (for comparison see Hodges et al. 2001). However, in an inclined channel beneath flat topography and without other driving forces (i.e. tectonic overpressure or buoyancy in the crust with density variations) there will be no flow (S. Medvedev, pers. comm. 2005). The pressure gradient can be approximated as the difference in pressure beneath thick and thin crust. Considering Airy compensation the lateral pressure gradient can be written as: dp ~ 9 p~gAh dx L/2

Pcg Oh(x) 21,Xc Ox z(hch -- z)

(7)

The pressure gradient across a mountain range can increase due to build-up of surface topography, i.e. evolution of the continental plateaus (e.g. Medvedev & Beaumont 2006). Assuming lithostatic pressure, constant viscosity of the channel and uniform density of the crust, the material flux (Equation 4) becomes: U_

pcgh3ch aS

will influence the rate of flow in the channel. Using this approach and suggesting a viscosity of 1018 Pa s for the lower crust, Clark & Royden (2000) calculated that topography could be built by a constant flux of material into a crustal channel from beneath the thick part of the plateau. This may explain the eastward propagation of the Tibetan Plateau and the slope distribution along the eastern Plateau margin as a function of channel thickness and/or foreland crustal structure. Medvedev & Beaumont (2006) solved Equation 9 analytically for a constant boundary flux, U = cons, for the evolution of upper surface with time:

(6)

(after Kruse et al. 1991), where 9 is the isostatic amplification factor (Pm- Pc)/Pm (Turcotte & Schubert 2002), Pm is the density of the mantle, Pc the uniform density of crust, g is the acceleration due to gravity, 2~h are the variations in crustal thickness, and L is the horizontal length scale of transport (characteristic horizontal scale of topographic perturbations). This equation shows that the rates of flow depend on the characteristic horizontal length of pressure distribution. The velocity in the viscous layer subject to variations of lithostatic pressure is given by Gemmer et al. (2004) as:

Up --

29

(8)

121.t,ch OX

(Medvedev & Beaumont 2006) where S is the topographic surface elevation. Using Airy isostatic equilibrium the topographic elevation S over time, as a result of flux of crustal material in the lower crust, can be written as: (9)

(after Clark & Royden 2000). In other words a weak crustal channel will influence the topographic wavelength and in the same time a topographic gradient

where S(x, t = 0) = 0 is the initial position of the surface, a=(~pcgh3ch)/31~ch, erfc is the complementary error function (Turcotte & Schubert 2002), and t is time. Testing for range of model parameters values, Medvedev & Beaumont (2006) propose three end-members for growth of continental plateaus in the presence of a weak crustal layer. However, the existence of a large-amplitude, shortwavelength topographic relief along the eastern margin of Tibet led McKenzie & Jackson (2002) to argue (assuming different boundary conditions from Clark & Royden 2000) that it is unlikely that extensive lower crustal flow has occurred there because of a short response time at short topographic wavelengths. Alternatively, Zhong (1997) used Maxwell theology and assumed a minimum viscosity for the lower crust of 1021 Pa s to show that short-wavelength topography could remain uncompensated for extensive periods of time. Most recent landscape analyses of the Tibetan Plateau have indicated the existence of areas of anomalously high topography (Burchfiel 2004; Clark et al. 2005). At least along the eastern plateau margin, around the Sichuan basin, such high topography could be interpreted as dynamic (Clark et al. 2005). The authors apply the geometry of a thin viscous channel with a rigid cylindrical obstacle (simulating the crust of the Sichuan basin) perpendicular to the bounding plates that causes dynamic pressure to the base of the overlying elastic crust. Relative to the background channel flow, such a dynamic loading can cause surface uplift upstream, and subsidence downstream of the rigid obstacle. If this concept is correct, then dynamic topography may explain the short-wavelength topography observed along the eastern Tibetan margin, despite predictions that the short response time would prevent such landscape development in the case of an active crustal channel (McKenzie & Jackson 2002).

30

D. GRUJIC

The shape of the velocity profile for Poiseuille and Couette flow is a function of the rheology of the channel material and of the temperature distribution across the channel, respectively. The velocity profile of channel flow of viscous fluids features a velocity gradient that decreases towards the centre of the channel. Hence the shear stress Os transmitted by a fluid layer also decreases toward the channel centre. Turcotte & Schubert (2002, equation 6 - 8 ) show the shear stress gradient in the channel.

(a) - 0.5

0.4 0.3 0.2 0.1

]

o

dos dp dz--dx

(11)

0.1

For a linear viscous fluid the Poiseuille flow velocity profile is a parabola that is symmetric about the centreline of the channel (Fig. 3a). In a power-law fluid, however, for increasing values of n (the exponent of power-law creep) the gradients in the velocity become large near the walls where the shear stress is a maximum (Fig. 3a). As a consequence, a nearly rigid core flow develops where the shear stress is low. A similar velocity profile is present in Bingham fluids (viscous-plastic materials that are characterized by a yield stress). As Bingham fluids become solid when the applied shear stress falls below the yield stress, the deforming material may become solid in the centre of the channel. There a solid 'plug' will be moving within the flow. The effect of the power-law exponent is especially clear for the change between n = 1 and n = 2, (Fig. 3a). However, the question remains whether it is actually possible to distinguish between linear and power-law flows from the finite strain data in nature (C. Beaumont, pers. comm. 2004). The effective viscosity for channel flow of a power-law fluid is given by Turcotte & Schubert (2002, equation 7-127):

0.2 0.3 0.4 0.5 (b)

1.o I 0.8

0.6

,

I/l~

:

0.4

o

0

.

0

1.0

0.8

2

0.6

0.4

0.2

0

u~o lx~ff - - d u / d z

--

2z ~ l-, 4(n + 2)~ \hch/

By assuming a crustal structure, flow laws for crustal material, and the geometry of the putative viscous channel, the viscosity of the channel material can be estimated. Studies aimed at constraining physical parameters of the crust have yielded estimates of effective viscosity of the lower crust of 1017-102o Pa s for assumed channel thicknesses of 10-25 km (e.g. Kruse e t al. 1991; Wdowinski & Axen 1992; Kanfman & Royden 1994; Clark & Royden 2000; Medvedev & Beaumont 2006). However, the values of involved parameters (Equations 10 and 12) cannot be estimated precisely. The viscosity of the channel

Fig. 3. Velocity profiles U/U l v e r s u s z / h in a channel flow with different rheologies. (a) Velocity profiles for Poiseuille flow showing the effect of power-law exponent: n = 1 (linear viscous), n = 3 and 5. Adapted from Turcotte & Schubert (2002, their fig. 7-13). The plug-flow appearance of the velocity profile is a consequence of the stress dependence of the effective viscosity. (b) Velocity profiles for Couette flow for different temperature dependences of viscosity (i.e. different values of the dimensionless activation energy parameter). Shear stress o-, is a constant in the absenceof a horizontal pressure gradient. The upper wall is maintained at temperature To while the lower wall is kept at temperature Tl (T1 > To), assuming that the temperature difference Tl - To is small compared with To,. Adapted from Turcotte & Schubert (2002, their fig. 7-14).

CHANNEL FLOW AND COLLISION TECTONICS can be estimated with an accuracy not better than two orders of magnitude, while hc~ cannot be estimated with an accuracy better than one order of magnitude (S. Medvedev, pets. comm. 2005). The same effect can be produced by a thick channel with a high viscosity as by a thin channel with very low viscosity. Because the parameter c~ (Equation 10) scales as h3h/l~, an uncertainty in hch of a factor slightly more than 2 causes an order of magnitude uncertainty in/x even when the flux U is known (Medvedev & Beaumont 2006). The results are also critically dependent on whether the channel thickness, flux and viscosity are constant or not (Clark & Royden 2000; Medvedev & Beaumont 2006). Moreover, if the channel absorbed all the deformation, the viscosity of the rest of the crust cannot be estimated. The rheological structure, with Coulomb-plastic and power-law viscous regions, leads to a more complex flow structure in general. An important point about channel flow is that the Poiseuille flow produces shear traction on the base of the overburden resulting in the horizontal force Fp (Fig. 3b), which, in special cases, does not depend on viscosity: hch Fv=Tpcg(hl

-- h2)

(13)

(Fig. 2; Gemmer et al. 2004). This force, depending on the ratio between overburden thickness and channel thickness, and on overburden strength, may lead to failure of the overburden, i.e. failure of the upper crust (Beaumont et al. 2004; Gemmer et al. 2004). If the stresses are above the yield strength in Coulomb materials, near-surface extensional flow may occur giving rise to various structures produced by interaction between the viscous channel and failure of the overburden. For example 'plateau collapse' may be the effect of the channel flow, rather than of gravitational instability of a thickened crust. In numerical models (e.g. Beaumont et al. 2004, 2006) the flux and the channel thickness change and may lead to inflation or deflation of the channel where the flow rate changes along the channel. For example, the formation of North Himalayan Gneiss domes (Hodges 2000, 2006), or out-of-sequence thrusts within the GHS (Grujic et al. 2002, 2004) can be interpreted in terms of channel flow in an inflating channel (Beaumont et al. 2004). The Couette flow velocity field is also affected by the temperature distribution in the channel which influences the viscosity with an exponential dependence on the inverse absolute temperature (Equation 14). When account is taken of heating by viscous dissipation in the shear flow, the temperature dependence of the viscosity couples the

31

temperature T(z) and velocity u(z) profiles in the channel. Both quantities T(z) and u(z) must be determined simultaneously, since one depends on the other. The velocity u depends on T through the dependence of /x on T, and T depends on u because frictional heating depends on the amount of shear in the velocity profile. Turcotte & Schubert (2002, equation 7-248) give the equation for the effective viscosity /-Left of a power-law material:

/Zeff -- 2Co_,,_1 exp

(14)

where C is the pre-exponential term of power-law creep, o" is the flow stress, Ea is the activation energy, R is the Boltzmann constant and T is the absolute temperature (K). When fluid viscosity is independent of temperature (the dimensionless activation energy parameter, Ea/RTo = 0), the velocity profile is linear (Fig. 3b). As the viscosity becomes increasingly temperature-dependent (larger values of Ea/RTo), the shear is confined to progressively narrower regions where the fluid is hottest and the viscosity is smallest (Fig. 3b). Thus, frictional heating can have consequences on the shear flow of a fluid with a strongly temperature-dependent viscosity. The transition from Couette flow- to Poiseuille flow-dominated channel flow is highly sensitive to viscosity and lateral pressure gradients. A critical change of viscosity within the channel will trigger changes in the flow pattern. For high viscosity material, Couette flow dominates and channel material will be underplated; for low viscosity material, Poiseuille flow will be dominant (e.g. Mancktelow 1995; Fig. 2c) and will cause flow opposite to the movement of the bounding plates, i.e. return flow of previously buried rock mass. Return flow in the channel is, therefore, a transient phenomenon that is recorded in the changing kinematic character of the structures within it (Fig. 4). The viscosity of the rock in the channel is likely to change significantly over short intervals of geological time (Beaumont et al. 2001a, b). A change in viscosity could be due to a metamorphic reaction leading, for example, to melting (e.g. Hollister 1993). Although certain details are disputed, it is evident that partial melting decreases the strength and the effective viscosity of rocks over several orders of magnitude (e.g. Cruden 1990). The volume of partial melt, not the temperature, controls the strength of rocks, and various values of critical melt percentage needed to dramatically reduce strength have been listed in the literature (Rosenberg & Handy 2005). Recent analyses suggest that in the range between solidus and a

32

D. GRUJIC

Fig. 4. (a) Schematic diagram of the deformation sequence along the upper boundary of a channel. S 1, thrust shear zone formed by Couette flow (prograde metamorphism); $2, ductile normal sheafing formed during return flow in the channel or during ductile extrusion (or both) at thermal peak of metamorphism and/or isothermal decompression. (b) Sample of deformed granite from the 2 km wide zone of the top-to-the north shear at the top of the GHS, NW Bhutan. Sample is c. 10 cm long.

7% melt fraction, the volumes of partial melt are sufficient to lower the strength by almost one order of magnitude (Fig. 5; Rosenberg & Handy 2005). Experiments on the mechanical behaviour of partially molten synthetic granite (Rutter et al. 2005) provide the most recent estimates of the theological behaviour of partially molten rocks. The flow law for low strain data is of the form: m

==Aexp(Bf

F-H1

.

) exp [-R-~-J o-

Fig. 5. Schematic plot of aggregate strength Ao- versus melt fraction f for partially melted granite between the liquidus and solidus. The vertical scale of the lower part of the ordinate is exaggerated in order to make the LST (solid-to-liquid transition) visible. There are two steep segments of the strength curve corresponding to the MCT (melt connectivity transition) and to the LST, or RCMP (rheological critical melt percentage). Adapted from Rosenberg & Handy (2005).

(15)

where A and B are constants, m and n are exponents, H is the activation enthalpy, e is the strain rate (s-1), f is the melt fraction, o-is flow stress, and T is the temperature (Rutter et al. 2005). The extrapolation to geological strain rates using the above flow law suggests that in nature migmatites

containing granitic melt will be extremely weak much weaker than silicate rocks deforming by intra-crystalline plasticity (Rutter et al. 2005). The weakening of rocks with increasing melt fraction is neither linear (m = 2, Equation 15) nor gradual. As suggested by Rosenberg & Handy (2005), the mechanical response of a rock containing c. 8% melt may be very different from rock that contains

CHANNEL FLOW AND COLLISION TECTONICS 7% or less melt, but is not significantly different from rock that contains 50% melt. In rocks with a high melt percentage the stresses will be low and thus deformation-driven melt extraction processes hindered; weak rocks stay weak. This implies complex feedback mechanisms between rheological weakening by partial melting, melt extraction processes by deformation, and deformation localization. Moreover, the effective viscosity of the power-law fluid is proportional to o-~-" (Turcotte & Schubert 2002). For large n the viscosity is high where shear stress is small and low where shear stress is large. The principle of channel flow is based on uniform distribution of stresses in the crust resulting in high strains in weak parts. However, weak parts in the centre of the channel have low stress, and consequently high viscosity if they deform by power-law rheology. Based on this principle there would be no channel flow. Rutter et al. (2005) propose that the stress exponent n for migmatites is 1.8 - smaller than for most geological materials (Twiss & Moores 2000; Turcotte & Schubert 2002; Pollard & Fletcher 2005). Accordingly, the effective viscosity of partially molten rock depends less on strain rate and stress but is more a function of temperature and pressure. In addition, near the walls where shear stress is high, the effective viscosity is low in a power-law fluid, and the velocity gradient is large (Fig. 3a). It follows that strain weakening caused by strain partitioning into shear zones is less important than hardening of the slower deforming material surrounded by shear zones (N. Mancktelow, pers. comm. 2005). Consequently, in situ melting is the most likely strength-lowering mechanism needed to initiate and sustain channel flow, although other weakening processes have also been suggested (Beaumont et al. 2004).

Discussion and conclusions In summary, crustal extrusion can be defined as the forcing out of a crustal sliver or wedge above and below brittle or ductile shear zones, maintained by concurrent crustal driving forces and focused surface erosion. Channel flow is a lateral flow of a weak crustal layer between stronger bounding layers. In active orogens the viscous flow is driven by shear stresses induced by relative motion between moving bounding plates and by a pressure difference along the channel. In collisional orogens we may consider the upper plate stationary and the lower one underplating or subducting. However, the upper plate may fail and its extensional movement be influenced by the viscous channel. The movement in the upper plate may in turn influence the flow in the viscous channel.

33

The pressure gradient in a crust with uniform density is influenced by the variations in crustal thickness and therefore surface relief. In addition, non-parallel channel walls may cause tectonic overpressuring, and the channel material may have a lower density than the surrounding crust and thus buoyancy forces may drive the flow as well. The boundaries of a channel are more likely to be transitional than sharp. At depth the rheology of a channel changes as a function of temperature, pressure, lithology and melt content. Based on different numerical models of channels, and on field observations, at least three types of bounding conditions for a channel can be proposed: (a) rheological boundary - such a boundary may be due to lithologic contacts (e.g. salt-sediment boundary; Gemmer et al. 2004) or rheologic stratification within the crust (Lobkovsky & Kerchman 1991); (b) phase transitions - the most likely process is partial melting as briefly described here and assumed in some numerical models; (c) thermal boundary, e.g. exponential thermal weakening of crustal material (Bird 1991). In nature the first type of channel boundary may be marked by the sharpest gradient in mechanical properties, while the third type would be characterized by the smoothest transition. Yet the idealized thermal boundary condition is widely used in modelling (e.g. Clark & Royden 2000). The first two types of channel boundary may be mapped in the field as protolith boundaries (e.g. Davidson et al. 1997) while the thermal boundary may be mapped as a metamorphic isograd (e.g. Jamieson et al. 2004). If all these boundaries are present in an orogen they probably have parallel strikes but are differently inclined, and very likely located at different crustal levels. In the Himalaya the presence of very close, nearly parallel shear zones at both MCT and STD level (e.g. Searle & Godin 2003; Searle et al. 2003, 2006; Godin et al. 2006; Hollister & Grujic 2006) may indeed reflect multiple closely spaced boundary conditions. Changing boundary conditions allow for changes in the channel flow regime, along the channel and with time, and do not restrict the definition of channel flow to idealized 1D flows of the type discussed by Turcotte & Schubert (2002). The velocity profile of channel flow is controlled by the balance between the shear stress gradient across the channel and the pressure gradient along the channel, while the exhumation history of the channel is controlled by the pattern of surface erosion (e.g. Jamieson et aL 2002; Beaumont et al. 2004). In an active orogen both channel viscosity and surface relief or crustal thickness (i.e. pressure gradient) are likely to change with time. In addition, numerical modelling allowing deformable and ductile bounding plates (e.g. Beaumont et al. 2004) suggests that

34

D. GRUJIC

the channel thickness is not constant, leading to much more complex tectonics. Further complication may arise due to the coupling between tectonics and surface processes. The results of thermal-mechanical models by Beaumont and co-workers (Beaumont et al. 2001b, 2004) suggest that surface denudation and strength of the upper crust may have a strong influence on flow in lowviscosity regions of the crust. Strong coupling between a crustal channel and topography indicates that concepts of coupling between surface processes (i.e. denudation) and tectonics that only take isostatic compensation into account may provide misleading results if there is an active channel. Potential diagnostic techniques for determining the existence of active channelized crustal flow include geophysical and landscape analyses. Recent geophysical data indicate a weak crustal layer and the presence of fluids, perhaps melt, underneath southern Tibet (e.g. Nelson et al. 1996) and suggest channel flow may be active there (e.g. Klemperer 2006). However, some interpretations of GPS (e.g. Flesch et al. 2005) and seismic data (assuming that the mid- to lower crust deforms as a Maxwell visco-elastic rheology, e.g. Hilley et al. 2005) exclude the possibility of decoupling between mantle and crust in part of the Tibetan crust, and thus argue against a weak crustal layer there. Landscape analyses suggest that the topography of the Tibetan Plateau, its eastward propagation and steep southern boundary (i.e. the Himalaya) may be explained by the existence of a lower- or midcrustal channel flow (e.g. Fielding et al. 1994; Clark & Royden 2000; Hodges et al. 2001; Shen et al. 2001; McKenzie & Jackson 2002; Clark et al. 2005). However, if different boundary conditions are assumed, the same landscape features may argue against a weak crustal channel (e.g. Zhong 1997). In the case where a channel has been exhumed, geological criteria apply. (a) High grade rocks are bounded by subparallel, coeval ductile shear zones with opposite sense of shear. (b) The shear zone at the roof shows kinematic inversion. (c) In the estimated initial orientation, the dominant deformational/metamorphic foliation was subhorizontal (e.g. Williams & Jiang 2005). (d) Palaeo-channels should have unique structural and metamorphic histories; structures alone are not sufficient, structural histories are important. (e) The metamorphic histories of rocks should progressively change across the channel and show jumps across the boundaries of a channel (e.g. Jamieson et al. 2004, 2006). (f) The channel is represented by high-grade rocks, with syntectonic partial melt suggesting that during deformation these rocks were weaker relative to both under- and overlying crustal layers. (g) The high-grade rocks were extruded

between the two bounding shear zones from beneath a coeval plateau. (h) There is strong orographic, focused erosion along the metamorphic belt interpreted as the palaeo-channel. High-grade quartzo-feldspathic rocks and migmatites may thus be the most likely lithology in an exhumed crustal channel, and melt weakening is the most likely viscosity-lowering mechanism needed to promote channel flow. An outstanding question regarding channel dynamics therefore relates to the melting of rocks due to the pressure distribution in a channel (e.g. Harris et al. 2004), and encompasses processes such as partial melt segregation, magma coalescence, migration and emplacement. Melt weakening, however, is difficult to implement in numerical models (e.g. Beaumont et al. 2004) and at the moment there are no analytical models able to predict the style of deformation involving melt weakening and the development of channel flow. In nature the boundaries of a channel show smooth gradients in mechanical properties and both the weak channel and bounding crustal layers above and below are deforming; the deformation is pervasively distributed throughout the channel but is concentrated in the weakest parts of the deforming crust causing channelized flows. Both channel flow and extrusion are influenced by surface denudation, which is in turn influenced by these two tectonics processes because of their control on surface uplift and surface slope. Therefore, the most important questions are what is a crustal channel in nature, and which boundary conditions produce intra-crustal channel flow? This work has benefited from discussions with C. Beaumont,R. Jamieson and S. Medvedev.Thoroughand constructive reviews by R. Bendick and S. Medvedev supported by meticulous editorial work by R. Law have greatly helped to improve the manuscript. Support from the Canadian Institute for Advanced Research - CIAR, is greatly acknowledged.

References BATCHELOR, G. K. 2000. An Introduction to Fluid Dynamics. Cambridge University Press. BEAUMONT, C., JAMIESON, R. A., NGUYEN, M. H. & LEE, B. 2001a. 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. 200lb. Mid-crustal channel flow in large hot orogens: results from coupled thermal-mechanical models. In: Slave-Northern Cordillera Lithospheric Evolution (SNORCLE) and Cordilleran Tectonics Workshop Transect Meeting 79, 112-170.

CHANNEL FLOW AND COLLISION TECTONICS

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Crustal flow in Tibet: geophysical evidence for the physical state of Tibetan lithosphere, and inferred patterns of active flow S. L. K L E M P E R E R

D e p a r t m e n t o f Geophysics, Stanford University, Stanford, CA 94305-2215, USA (e-mail: [email protected]) Abstract: Many seismic and magnetotelluric experiments within Tibet provide proxies for litho-

spheric temperature and lithology, and hence rheology. Most data have been collected between c. 88~ and 95~ in a corridor around the Lhasa-Golmud highway, but newer experiments in western Tibet, and inversions of seismic data utilizing wave-paths transiting the Tibetan Plateau, support a substantial uniformity of properties broadly parallel to the principal Cenozoic and Mesozoic sutures, and perpendicular to the modern NNE convergence direction. These data require unusually weak zones in the crust at different depths throughout Tibet at the present day. In southern Tibet these weak zones are in the upper crust of the Tethyan Himalaya, the middle crust in the southern Lhasa terrane, and the middle and lower crust in the northern Lhasa terrane. In northern Tibet, north of the Banggong-Nujiang suture, the middle and probably the lower crust of both the Qiangtang and Songpan-Ganzi terranes are unusually weak. The Indian uppermost mantle is cold and seismogenic beneath the Tethyan Himalaya and the southernmost Lhasa terrane, but is probably overlain by a northward thickening zone of Asian mantle beneath the northern Lhasa terrane. Beneath northern Tibet the upper mantle has not been replaced by subducting Indian and Asian lithospheres, and is warmer than to the south. These inferred vertical strength profiles all have minima in the crust, thereby permitting, though not actually requiring, some form of channelized flow at the present day. Using the simplest parameterization of channel-flow models, I infer that a Poiseuille-type flow (flow between stationary boundaries) parallel to India-Asia convergence is occurring throughout much of southern Tibet, and a combination of Couette (top-driven, between moving boundaries) and Poiseuille lithospheric flow, perpendicular to lithospheric shortening, is active in northern Tibet. Explicit channel-flow models that successfully replicate much of the large-scale geophysical behaviour of Tibet need refinement and additional model complexity to capture the full details of the temporal and spatial variation of the India-Asia collision.

'Only fools, or unusually insightful individuals not yet recognized to be ahead of their time, would doubt that the warmth of the lower crust in regions of extension makes decoupling of upper crust and upper mantle more likely there than in other settings. Thus, the real challenge to understanding how such decoupling or coupling occurs will require study of regions of intracontinentalcrustal shortening' (Molnar 2000).

Lateral strain variations, and vertical strain and strength profiles in Tibet Tibet forms the largest and highest plateau on Earth today. Two extreme and opposed views of the mechanism(s) responsible for regional uplift and a c c o m m o d a t i o n of shortening are that: (1) discrete tectonic blocks, internally relatively undeformed, are being expelled eastward between lithospheric strike-slip faults (e.g. Tapponnier et al. 1982, 2001); and (2) deformation is essentially continuous, with diffuse deformation of the crust and upper mantle over broad areas (e.g. Dewey & Burke 1973; England & Houseman 1986; England & Molnar 1997; Shen et al. 2001), in addition to

the convergence taken up at the plateau margins. With the increasing availability of global-positioning data from Tibet and adjacent regions, both in n u m b e r of sites (now exceeding 500: Zhang et al. 2004) and length of their time series (up to nine years: Chen et al. 2004), the extreme model of coherent blocks separated by narrow fault zones along which most deformation is localized seems a less plausible description of the active deformation of Tibet. At the Earth's surface there is a continuously varying surface strain field across Tibet when observed at scales of >100 k m (Wang et al. 2001; Zhang et al. 2004). The major active strike-slip faults such as the Kunlun and Jiali faults (Fig. 1) have motions c. 10 m m a -a and separate internally deforming blocks in which about two-thirds of the total 25 m m a-1 eastward extrusion of Tibet is a c c o m m o d a t e d by smaller normal faults and conjugate strike-slip faults (Taylor et al. 2003; Chen et al. 2004; Zhang et al. 2004). In contrast, vertical strength and strain profiles in Tibetan lithosphere cannot be directly measured, and so remain less certain and more contentious.

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, 39-70. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

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Fig. 1. Location map of main geophysical experiments referred to in this paper. Heavy black lines and upper-case italicized names are generalized location of seismic experiments; heavy dotted lines and lower-case italicized names are magnetotelluric profiles; thin lines are major sutures and faults; thin dotted line is the 3000 m contour. Faults: KKF, Karakax fault; ATF, Altyn Tagh fault; QF, Qaidam Border fault; KF, Kunlun fault; JRS, Jinsha River suture; BNS, Banggong-Nujiang suture; JF, Jiali fault; IYS, Indus-Yarlung suture; MCT, Main Central thrust; MBT, Main Boundary thrust. Himalayan syntaxes: Np, Nanga Parbat; Nb, Namche Barwar. Representative geothermal areas: Ybj, Yangbajain graben in Yadong-Gulu rift; Pu, Puga. Representative north Himalayan gneiss domes: Kd, Kangmar dome; Tmc, Tso Morari complex. Profile names (and selected references): seismic, INDEPTH, IN-1 (Zhao et al. 1993; Makovsky et al. 1996), IN-2 (Nelson et al. 1996; Kind et al. 1996; Alsdorf et al. 1998b; Makovsky & Klemperer 1999), IN-3 (Zhao et al. 2001; Ross et al. 2004; Huang et al. 2000; Rapine et al. 2003; Tilmann et al. 2003); Sino-French, SF-1 (Him et al. 1984a), SF-2 (Zhang & Klemperer 2005), SF-3 (Him et al. 1995; Galv~ et al. 2002b), SF-4 (Herquel et al. 1995; Vergne et al. 2003), SF-5 (Wittlinger et al. 1998), SF-6 (Galv~ et al. 2002a; Vergne et al. 2003), SF-7 (Wittlinger et al. 2004); PASSCAL (McNamara et al. 1994, 1995, 1997; Rodgers & Schwartz 1998; Sherrington et al. 2004); WT, West Tibet (Kong et al. 1996); TB, Tarim basin (Gao et al. 2000; Kao et al. 2001); NP, Nanga Parbat (Meltzer et al. 2001); HIMPROBE (Rai et aL 2006); TQ, Tarim-Qaidam (Zhao et al. 2006); HIMNT, HimalayaNepal-Tibet experiment (Schulte-Pelkum et al. 2005); NE INDIA (Mitra et al. 2005). Near-vertical reflection profiles are IN-I, IN-2 and locally along IN-3, TB and HIMPROBE; controlled-source wide-angle profiles are SF-1, SF-2, SF-3, SF-6, IN-I, IN-2, IN-3, WT and TQ; teleseismic recording experiments (typically for tomography, receiverfunction, shear-wave-splitting, and waveform inversions) are PASSCAL, IN-2, IN-3, SF-3, SF-4, SF-5, SF-6, SF-7, TB, NP, HIMPROBE, HIMNT and NE INDIA. Magnetotelluric profiles: INDEPTH, in-100, in-200 (Chen et al. 1996), in-500 (Wei et al. 2001), in-600 (Unsworth et al. 2004), in-700 (Spratt et aL 2005), in-800 (Unsworth et al. 2005); himprobe (Gokarn et aL 2002); nf, Nepal-French (Lemonnier et al. 1999); wt, West Tibet (Kong et aL 1996).

Opposing views are that the entire lithosphere deforms homogeneously ('vertical coherent deformation') (e.g. England & Houseman 1986; England & Molnar 1997; Flesch e t al. 2005), or that deformation is dominated by a more rapid ductile flow in the middle and/or lower crust above a stronger upper mantle ('channel flow') (e.g. Zhao & Morgan 1987; Bird 1991; Shen e t al. 2001; Beaumont e t al. 2004). The depth extent of such channelized flow - middle crust, lower crust, or both - depends critically on the actual

strength profile. These opposing views have recently been conflated with a new controversy about vertical strength profiles for the continental lithosphere. Laboratory experiments have long been used to infer a combined frictional plus ductile behaviour of the lithosphere, with a quartz-dominated strength maximum in the uppermiddle crust, and an olivine-dominated strength maximum at the top of the upper mantle (Brace & Kohlstedt 1980). It is also widely held that the greater part of lithospheric strength is in the upper

CRUSTAL FLOW IN TIBET crust for typical continental thicknesses and heat flow (e.g. Zoback et al. 2003). For 25 years discussions of continental rheology have been dominated by the belief that although earthquakes are largely localized in the upper crust and essentially absent from the lower crust, they do occur, though rarely, in the upper mantle (Meissner & Strehlau 1982; C h e n & Molnar 1983), leading to the strong upper crust-weak lower crust-strong upper mantle 'jelly-sandwich' model of continental lithosphere. A strength minimum in the middle-lower crust permits channel flow and complete decoupling of the upper crust and upper mantle, but does not require it: boundary forces arising from plate motions could be applied equally at all depths in the lithosphere, leading to vertically coherent deformation. However, basal drag due to asthenospheric flow is applied to the mantle lithosphere, and body forces due to topographic loads are applied to the crust, and so these will be transmitted upwards or downwards only to the extent that the upper crust is mechanically coupled to the mantle (Flesch et al. 2005). An opposing view of continental strength profiles arises from a re-evaluation of the focal depths of supposed upper-mantle earthquakes that places many of these in the lower crust (Maggi et al. 2000) leading some to argue that the lower crust is typically stronger than the upper mantle (Jackson 2003; though see Burov & Watts (2006) for counter-arguments). In this case the lack of a broad strength minimum throughout the lower crust must greatly inhibit channel flow, the timescale of which varies inversely with the cube of the layer thickness in which such flow is occurring (for a Newtonian fluid) (Kusznir & Matthews 1988). Even with a weak upper mantle, intracrustal channel flow may still occur as a transient phenomenon due to the introduction of melts or hydrous fluids (McKenzie & Jackson 2002), and widespread lower-crustal flow may occur in conjunction with upper-mantle flow if the viscosities of both are low (Jackson 2003). Clearly, the strong dependence of rock strength on both composition and temperature (e.g. Afonso & Ranalli 2004) as well as hydration state and presence of partial melt (e.g. Kirby 1984; Rosenberg & Handy 2005) means the relative strength of lower crust and upper mantle will vary in time and space, and that an evaluation of this relative strength requires detailed knowledge of the lithosphere. M o d e s o f crustal f l o w a n d c h a n n e l f l o w

Any viscous or plastic deformation may be regarded as flow. Such flow need not be channelized within the crust or lithosphere - it may affect the entire outer shell, as in the 'thin viscous sheet' model applied to Tibet by England & Houseman

41

(1986) and England & Molnar (1997), or in the 'vertically coherent deformation' style inferred for Tibet from seismic and geodetic data by Silver (1996) and Flesch et al. (2005). In these models, channelized flow is precluded. In contrast, channel flow as used in this paper refers to any flow in which a viscosity minimum at some depth strongly localizes horizontal material flow and partially or totally decouples flow at different depths (e.g. for Tibet, models of Zhao & Morgan 1987; Bird 1991; Royden 1996; Royden et al. 1997; Beaumont et al. 2004). Although it is possible to model channel flow with free or partly free boundary conditions (McKenzie et al. 2000; McKenzie & Jackson 2002) such models are only now being applied to Tibet (Bendick 2004). The simplest numerical models have explicit channel boundaries with imposed velocity and position conditions. Two end-members are Poiseuille flow (or pipe flow), an extrusion between stationary boundaries in which the induced pressure gradient produces the highest velocities in the centre of the channel (Fig. 2a), and Couette flow, in which shear across the channel produces the highest velocities at the top or base of the channel (Fig. 2b).

Fig. 2. Modes of channel flow with imposed rigid boundary conditions. Columns of arrows represent relative horizontal velocities. Shaded zones are lowviscosity and undergoing channel flow. No vertical scale is implied; the flowing layer may be all or part of the middle and/or lower crust and/or upper mantle. (a) Poiseuille flow (channel flow between stationary rigid boundaries) driven by a pressure gradient. (b) Couette flow (channel flow between moving rigid boundaries) driven by upper-crustal extrusion (above) or lithospheric underthrusting (below). (c) Combined Poiseuille and Couette flow.

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These conceptually useful mathematical simplifications allow approximation of important parameters such as viscosity and channel thickness, but cannot be exact descriptions for Tibetan (or any other midlower crustal) flow. Pure Poiseuille flow has been used to model the gentle slopes of the eastern plateau margin (Clark & Royden 2000), and an analogous form of flow emerges naturally in numerous Tibetan deformational models (e.g. Bird 1991; Royden 1996; Royden et al. 1997; Shen et al. 2001) as a low-viscosity channel within or at the base of the crust. Such models incorporate depthdependent lower-crustal strength (Royden et al. 1997; Shen et al. 2001) or thermal weakening (Beaumont et al. 2004) that effectively confines flow between boundaries that do not move substantially. In contrast, Couette flow refers to a channelized flow with moving rigid boundaries, for example in which gravitational collapse with outward translation of the upper crust, or underthrusting of cold lithosphere, drives flow in the lower crust by viscous coupling (Beaumont et al. 2004) (Fig. 2b). If the entire lower lithosphere is weak, then in this form of flow the entire lithosphere shares the same sense of motion, with velocity decreasing downwards. In recent discussions of Tibetan geodynamics, 'channel flow' has typically been used to describe deformation with Poiseuille (Fig. 2a) or hybrid Poiseuille-Couette characteristics (Fig. 2c). Independent of numerical models, the concepts of channel flow have been applied to both the Himalaya and Tibet at multiple scales. At a subregional scale, in the NW syntaxis of the Himalaya (Nanga Parbat, Np in Fig. 1), observations of simultaneous highgrade metamorphism, anatexis and rapid denudation led to the recognition of local feedbacks between tectonic and surface processes, and to the 'tectonic aneurysm' model (Zeitler et al. 2001) in which large-magnitude fiver incision focuses deformation of weak crust, leading to crustal flow into the region (Koons et al. 2002). The very rapid denudation leading to exposure of Quaternary metamorphic and magmatic rocks, coupled with low seismic velocities, high seismic attenuation and shallow microseismicity (Meltzer et al. 2001) leaves little doubt that strongly localized crustal flow exists at present beneath Nanga Parbat (Zeitler et al. 2001). Along the main Himalayan front, seminal observations of extrusion of the crystalline Greater Himalayan Sequence (GHS) (southward with respect to the lower-grade Tethyan Himalaya above and Lesser Himalayan Sequence below) were made first in the central (Burchfiel & Royden 1985), then in the western (Searle & Rex 1989) Himalaya, albeit lacking a channel-flow terminology. Grujic et al. (1996), working in the

eastern Himalaya, were the first to apply the mechanics of Poiseuille and Couette flow to ductile extrusion of the GHS. However, these and the earlier observations may be as compatible with transient ductile extrusion of a single wedge during growth of a critical-taper orogenic thrust belt, as with a long-lasting channel flow requiring continuous material replenishment into the channel at some unknown location north of the exposures of ductilely deformed GHS rocks. As applied to southern Tibet, the 'channel-flow model' typically refers to the concept, first clearly articulated by Nelson et al. (1996), that the GHS represents the ongoing extrusion of a fluid middle crust formed as Indian crust subducted northwards beneath the Lhasa terrane heats up, partially melts, and forms a southward-directed return flow driven by the gravitational potential energy of the Tibetan Plateau. In this view the leucogranites exposed within the GHS represent progressively younger, frozen, snapshots of the partially molten mid-crustal layer interpreted from INDEPTH seismic and magnetotelluric (MT) data north of the Indus-Yarlung suture (Brown et al. 1996; Makovsky & Klemperer 1999; Li et al. 2003). Wu et al. (1998) explicitly added to this model the importance of orographic precipitation (the Indian monsoon) to denudationally exhume the return flow to the surface between coaeval slip (Hodges 2006) on the normal-sense South Tibetan Detachment (STD) above and the thrust-sense Main Himalayan Thrust (MHT) or Main Central Thrust (MCT) below. This concept and a broad range of observed Himalayan metamorphic and geochronologic data have been successfully modelled in two dimensions with time-varying, plane-strain, coupled thermal -mechanical finite-element numerical models (Beaumont et al. 2004; Jamieson et al. 2004) in which sufficient radiogenic self-heating and mantle heat flux exist to create melt-weakening to reduce the channel viscosity to 1019 Pa s. The preferred model of Beaumont et al. (2004) and Jamieson et al. (2004) is successfully tuned to provide dramatically good fits to a wide range of geological observations, but requires some arbitrary forcing functions (e.g. an arbitrary time-variable surface denudation rate to drive the channel exhumation), and grossly simplifies some aspects of geological evolution (e.g. assumes a continuous evolution towards the modern low-angle subduction, in contrast to the widely recognized twostage evolution of the Himalaya (e.g. Leech et al. 2005) in which early steep subduction is terminated by a slab break-off event prior to establishment of the modern low-angle subduction). Though it seems remarkable that the Beaumont et al. (2004) model can fit current observations so well given the elements of geological history not contained

CRUSTAL FLOW IN TIBET in their model, Beaumont et al. (2006) do show that their modelled channel flows are not strongly perturbed by mantle processes, thereby mitigating at least one cautionary note. Although arbitrarily complex initial parameters and model specifications allow arbitrarily complex fits, the good fit thus far obtained provides encouragement for the future that more complex models based on a more complete understanding of initial conditions will continue to fit the complete geological history. At the largest scale, simpler models (lacking thermal-mechanical coupling but tractable in three dimensions) have explored how gravitational potential can drive outward channel flow from the Tibetan Plateau with wavelengths > 1000 km to produce marginal surface uplifts of c. 2 km, in a tunnelling Poiseuille mode (Royden et al. 1997; Clark & Royden 2000; Hodges 2006). In these models the flow is channelized in the vertical domain by a specified and strongly depth-dependent theology (Shen et al. 2001) and confined laterally by resistant (high-viscosity) crustal blocks along Tibet's southern and northern margins (India, Tarim, Qaidam) and locally along the eastern margin (Sichuan Basin) (Clark et al. 2005). The absence in eastern Tibet of exposed middle or lower crust (the lack of an analogue to the GHS) is explained by insufficient denudation to exhume the channel. Such 'channel injection' models are further developed by Medvedev & Beaumont (2006). Whether at the scale of the individual gneiss domes of the Himalayan syntaxes, the GHS along the length of the Himalaya, or the entire Tibetan Plateau, these proposed models may all be termed channel flow. Independent of terminology, a key task for geophysics is to define the spatial variation of relative vertical strength profiles in Tibet, controlled by lithology, temperature and fluid content; and hence to define the relative vertical velocity profiles, controlled by plate-tectonic boundary forces, topographic (gravity) forces, and asthenospheric basal drag forces.

Geophysical data bearing on crustal flow in Tibet Data sources

It is the aim of this paper to summarize the now large number of geophysical measurements in Tibet that bear on the composition, temperature and fluid content of Tibetan lithosphere. Referencing is of necessity selective: the Georef database contains 4000 articles published since 1990 with 'Tibet' or 'Himalaya' in the title. Twenty years ago a review of geophysical constraints on the

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deep structure of Tibet (Molnar 1988) was still able to comprehensively reference the entire preceding literature. Many seminal studies discussed by Molnar (1988) are not further reviewed here: it was by then already possible to conclude that Tibetan crust averages 65-70 km in thickness, and that there is marked lateral heterogeneity in the mantle beneath Tibet, with lower shear-wave velocities and higher attenuation defining a wanner mantle beneath northern Tibet that compensate a somewhat thinner crust there than in southern Tibet. The availability of portable digital recorders enabled a new generation of geophysical experiments in Tibet in the 1990s that continue to the present day, yielding data that in both quality and density exceed that possible in earlier projects. Figure 1 shows the locations of projects discussed in this paper. Acronyms in Figure 1, upper-case for seismic measurements, lower-case for magnetotelluric (MT) profiles, are used in the text to indicate from which data-set particular results were derived. Three collaborations have provided many crucial data reviewed here: (1) the pioneering multidisciplinary 'Sino-French' transect that recorded controlled-source seismic, teleseismic, MT and geothermal data in southern Tibet in 1980-1982, and continued across the entire plateau with a series of teleseismic and wide-angle seismic experiments through the 1990s and 2000s (sometimes utilizing the name 'Lithoscope') (transect locations SF-1 to SF-7 in Fig. 1); (2) the 1991-1992 SinoUS 'PASSCAL' transect that recorded the first broad-band seismic profile across Tibet on 11 seismographs at c. 150 km spacing; and (3) the 19921998 'INDEPTH' (International Deep Profiling of Tibet and the Himalaya) experiments that recorded the first deep near-vertical reflection profiles on the plateau (IN-1), and have since recorded controlledsource, passive-source (IN-l, -2, -3) and MT data (in-100, -200, -500, -600) in a corridor across the plateau (this transect will continue in 2007 as INDEPTH-4 across the northem boundary of Tibet). Without minimizing the contributions of many other groups, these three international programmes collected data that continue to be reanalysed today, a decade or more after acquisition. All three transects focused their efforts along the NNE-trending Lhasa-Golmud highway and areas accessible from it, at c. 88-95~ leading to a strong concentration of data-sets in eastern Tibet (Fig. 1). However, although reflection and refraction experiments typically obtain data only directly beneath the seismic array, shear-wave splitting and receiver-function observations pertain to a region several tens of kilometres around the recorders (depending on the depth of the anomaly), bodywave tomography and local-seismicity recordings effectively study a region several hundreds of

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kilometres around the array, and regional waveform modelling and surface-wave inversions provide measurements integrated over earthquake-to-receiver paths of 102-103 km that can average properties of the entire plateau, even if recorded on only a single station. Although MT measurements are largely sensitive to structure directly beneath the array, their relative low cost, speed and simplicity have allowed acquisition that spans the plateau from north to south along the INDEPTH transect, and from west to east in southern Tibet (in-100 to in-800, nf and himprobe in Fig. 1). The other fundamentally new datasets that became available in the 1990s were high-resolution topography and moderate-resolution gravity data for the entire plateau. Some newer geophysical transects along the southern margin of Tibet (HIMPROBE, HIMNT, NE India in Fig. 1) are sufficiently recent that their data are not yet fully published.

T o p o g r a p h y a n d gravity: e v i d e n c e f o r intra-crustal c o m p e n s a t i o n

The availability since the early 1990s of sub-100 m lateral-resolution topographic data for Tibet has allowed analysis of the wavelengths on which relief is developed. The low relief of the interior of the Tibetan plateau suggests that upper-crustal isostatic compensation acts to level the surface of the plateau, requiting a 'fluid lower-crustal process' the cmst is not strong enough to maintain large topography at the surface or corresponding stresses at depth (Fielding et al. 1994). Detailed analysis of short-wavelength flexural rift-flank uplift adjacent to Plio-Pleistocene grabens in Tibet implies a lowviscosity middle and lower crest (Masek et al. 1994) such that the crustal viscosity in a 15-kmthick channel would be c. 3 x 1020 Pa s. Braitenberg et al. (2003) and Jordan & Watts (2005) used a range of forward and inverse techniques to map the spatial variation of elastic thickness (Te) and to show the interior of the plateau has Te typically <20 kin, whereas the Himalayan foreland has Te >100 km (more than double older estimates of McKenzie & Fairhead 1997), and the Tarim, Qaidam and Sichuan basins north and east of the plateau have Te c. 50 km. Modelling of Te has also been used to define an area of lithosphere beneath southem Tibet that can be identified as Indian mantle sufficiently strong that it must subduct rather than undergo internal deformation (Jin et al. 1996). Hence Jin et al. map out the 'Indian mantle suture' (point at which top of subducting Indian crest intersects Tibetan Moho) and the 'Indian mantle front' (northemmost extent of Indian lithosphere beneath Tibet), at latitudes c. 29.5~ and c. 33~ along the LhasaGolmud highway.

In contrast to the southern and northern margins of Tibet, where Quatemary upper-crustal deformation is occurring as shown by modem seismicity and mapped fault pattems, the eastern margin of the plateau lacks large-scale Quatemary surface shortening (Royden et al. 1997). This can be modelled if the lower cmst is sufficiently weak (1018 Pa s in a 15-km-thick layer) to decouple the upper crust from the upper mantle (Royden et al. 1997; Shen et al. 2001). The topography of eastern Tibet and its time-evolution are both compatible with outward flow of a weak crest. Clark & Royden (2000) have shown that the more than order-of-magnitude contrast between the steep plateau margin adjacent to the Tarim basin and Sichuan basin, and the lowgradient slope of the NE and SE plateau margins, can be explained by a lower-crustal channel with lateral variations in viscosity. Assuming pure Poiseuille flow in a 15-km-thick channel, they inferred channel viscosities of 1021, 1018 and 1016 P a s beneath the steep margins gentle margins, and plateau interior, respectively. Evidence for sequential reorganization of drainage patterns in southeastem Tibet (Clark et al. 2004) and rapid late Cenozoic cooling in the southwestem Qinling (Enkelmann et al. 2006) as a result of Neogene uplift of the plateau margin is consistent with this model of long-wavelength intra-crustal channel flow. The relative viscosities of Clark & Royden (2000) are important if topography is a proxy for lower-crustal viscosity; their inferred viscosities are less important since they are strongly model-dependent and scale with the cube of the channel thickness. Indeed, the lowest viscosity inferred by Clark & Royden is inherently unlikely because it is hard to reconcile with plausible regional crustal viscosities even if the inferred channel occupies the entire crustal thickness (Hilley et al. 2005). Absolute viscosities < 1019 Pa s probably require a few per cent of in situ partial melt (cf. Beaumont et al. 2004; Rosenberg & Handy 2005), which should be detectable by magnetotelluric and seismic methods.

Geothermal measurements: evidence for recent m a g m a t i s m in s o u t h e r n Tibet

Geothermal exploration of the Puga region in NW India (Pu in Fig. 1) measured heat flow >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

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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 <4.5, recorded on a local network deployed for <18 months) have been located up to 40 km below the Moho (up to 100 km below sea-level) in a joint tomographic/receiver-function inversion (HIMNT: Schulte-Pelkum et al. 2005; Monsalve et al. 2005, 2006). This evidence for significant strength in the mantle beneath the Himalaya is unsurprising if elastic thicknesses there are indeed of the same order as crustal thickness (Jordan & Watts 2005).

Seismic velocity and crustal thickness: n o r t h - s o u t h dichotomy o f uppermost-mantle velocities but plateau-wide low crustal velocities Seismic velocities at any depth depend on lithology, temperature and fluid distribution. Low velocities

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tend to correlate with higher SiO2 content (Christensen & Mooney 1995), hence with lower melting point, and hence lower creep strength at any given temperature since the homologous temperature is lower at any given temperature (e.g. McKenzie & Jackson 2002). Seismic velocities decrease with increased temperature (Christensen & Mooney 1995), which weakens rocks. Velocities are also greatly reduced by addition of aqueous (Hyndman & Shearer 1989) or magmatic (Hammond & Humphreys 2000) fluids that also weaken their host rocks (Kirby 1984; Rosenberg and Handy 2005). Crust or mantle with unusually low seismic velocities is therefore weaker than average, and pronounced low-velocity zones relative to material above and below indicate lowstrength channels in the lithosphere. The recognition of marked lateral heterogeneity in the mantle beneath Tibet remains one of the most fundamental and reliable observations bearing on Tibetan geodynamics. Evidence for a warmer upper mantle beneath northern Tibet east of 85~ as compared to southern Tibet, was first found in strong attenuation of Sn (uppermost mantle shear-wave) (Ni & Barazangi 1983) and lower upper-mantle shear-velocity beneath a thinner crust inferred from Rayleigh-wave-phase velocities (Brandon & Romanowicz 1986). These early studies using seismic stations outside Tibet have been reinforced and extended to include P-wave observations from portable broadband deployments in eastern Tibet by the PASSCAL (McNamara et al. 1995, 1997), Sino-French Lithoscope (SF-4: Wittlinger et al. 1996) and INDEPTH (IN-3: Tilmann et al. 2003) groups. All these datasets show significant differences (McNamara: 3-4%; Tilmann: 2%) between a faster mantle south of approximately the Banggong-Nujiang suture and slower mantle further north, corresponding to a temperature increase of c. 200-300~ from south to north (McNamara et al. 1997), significantly lower than earlier estimates as high as 500~ (Molnar et al. 1993). Noting that all these temporary deployments were located east of 88~ Dricker & Roecker (2002) use teleseismic SS-S differential traveltimes (that are sensitive to structure to depths of c. 300 km) to show that the north-south mantle variability is only pronounced in eastern Tibet. In their model, Tibetan mantle west of 87~ is uniformly faster (colder) than to the east, providing an ongoing disagreement with the Pn tomography of McNamara et al. (1997) that shows low uppermost mantle velocities extending west to 80~ north of the Banggong-Nujiang suture. Unfortunately Pn has not been observed on controlledsource refraction profiles recorded over the Tibetan Plateau due to the great crustal thickness and strong crustal attenuation.

Despite the south-north change in temperature of the lithospheric mantle, lithospheric thickness appears relatively uniform. Kumar et al. (2006) use S-to-P receiver functions to map a plateau-wide discontinuity that they identify as the lithosphereasthenosphere boundary. The base of the Indian lithosphere dips north from c. 170km beneath the Himalaya to c. 210km beneath the BanggongNujiang suture; and the base of the Asian lithosphere is nearly horizontal at 170 to 190 km beneath central and northern Tibet (Kumar et al. 2006). Waveform modelling of earthquake phases at regional distances (e.g. so-called Pnl waveforms comprising the P arrival (Pn) followed by the crust-trapped P-SV multiple conversions (PL)) is an effective way to study crustal velocities averaged laterally over hundreds of kilometres. For Tibet, Rodgers & Schwartz (1998) demonstrate that average crustal P-wave velocity (Vp) in the Qiangtang and Lhasa terranes (6.2 and 6.0 km s-l) is significantly less than the global continental average of 6.45 km s -~ (Christensen & Mooney 1995). Controlled-source refraction experiments typically provide the best-resolved estimates of Vp (e.g. SF-I: Him et al. 1984a; IN-3: Zhao et al. 2001; SF-6:Galv6 et al. 2002a), and further demonstrate that Tibetan crust has unusually low seismic velocity at all depths compared to global averages (Haines et al. 2003). Despite the low average Vp, many studies - including receiver-function and related P-waveform modelling methods that detect the amplitude and polarity of Moho and intracrustal conversions - suggest a higher-velocity layer ( > 7 km s -1) at the base of the crust. This higher-velocity layer is c. 12km thick in the Tethyan Himalaya (SF-I: Him et al. 1984a), c. 20 km thick in the southern Lhasa terrane south of 31~ (IN-2: Kind et al. 2002), c. 10 km beneath the northern Lhasa terrane (PASSCAL: Zhao et al. 1996; Owens & Zandt 1997) and southern Qiangtang terrane (IN-3: Zhao et al. 2001), and c. 20 km thick in the Qiangtang and Songpan-Ganzi terranes (SF-4 and -6: Vergne et al. 2002). The presence of this higher-velocity lower-crustal layer means that the upper and middle crust must have even lower velocity for the crustal average to be so low. The high-velocity layer may represent Indian cratonic lower crust: DeCelles et al. (2002) offer the provocative hypothesis that 'Greater Indian lower crust' extends beneath Tibet to the Kunlun fault, even if Greater Indian mantle extends no further north than southern Tibet. Receiver functions provide independent estimates of crustal velocity structure beneath individual stations, e.g. along the INDEPTH-2 transect (Yuan et al. 1997) or across the entire plateau along the 1991-1992 Sino-US PASSCAL experiment (Owens & Zandt 1997). These studies

CRUSTAL FLOW IN TIBET identify a distinct low-velocity zone below 20 km depth in the Lhasa terrane which dies out southward approximately at the I n d u s - Y a r l u n g suture (IN-2: Kind et al. 1996), and confirm the unusually low average-crustal Vp across the entire plateau, as well as suggesting low lower-crustal S-wave velocity (Vs) and correspondingly high V p / V s in the northern plateau (PASSCAL: Owens & Zandt 1997; IN-3: Tian et al. 2005). A reinterpretation of the PASSCAL results in conjunction with transect SF-4 (Vergne et al. 2002) highlights the inherent non-uniqueness of receiver-function inversions (even in the absence of anisotropic or dipping structure) (Ammon et al. 1990), and suggests V p / V s ratios north of the B a n g g o n g - N u j i a n g suture are normal, with no evidence for widespread partial melting beneath the Qiangtang terrane. Surface-wave studies integrate over large regions of the crust and are particularly useful for picking out broad trends in crustal velocity (Fig. 3). Early teleseismic surface-wave observations used to

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document a crustal low-velocity zone throughout Tibet (Chun & Yoshii 1977) have been refined by more detailed observations from the INDEPTH transects that provide distinct averaged velocitydepth profiles for the Qiangtang and the northern Lhasa terranes (IN-3: Rapine et al. 2003), and for the Lhasa terrane and Tethyan Himalaya (IN-2: Cotte et al. 1999). All the V s - d e p t h profiles derived from surface-wave inversions for Tibet are significantly low compared to global crustal values (Fig. 3), as expected from the refraction and waveform-modelling Vp results. Additionally it is clear that there is a strong low-velocity channel in the mid-to-lower crust of the Lhasa terrane, placed at 2 0 - 4 0 km depth (IN-3: Rapine et al. 2003), 3 0 - 5 0 km (IN-2: Kind et al. 1996) or 3 0 - 7 0 k m (IN-2: Cotte et aL 1999). This channel becomes less pronounced to the north in the Qiangtang terrane because the overall lowercrustal velocity is lower in the Qiangtang than in the Lhasa terrane (Rapine et al. 2003). The

Fig. 3. Surface-wave velocity models along IN-I, -2 and -3. S-wave velocity-depth models of Cotte et al. (1999) for the Tethyan Himalaya and southern Lhasa terrane and of Rapine et aL (2003) for the northern Lhasa terrane and Qiangtang terrane are shown, each displaced laterally by 1 km s 1from the previous model to allow easy viewing, and to convey a sense of possible continuity of channels of unusually low velocity (grey-shaded regions). Possible dominant flow directions (see text and Fig. 8) are shown by south-directed arrow in southern grey region, and east (outof-the-page)-directed arrowhead/bullseye in northern grey region. Note possibility of multiple, vertically stacked, independently flowing weak zones, e.g. in northern Lhasa terrane. Dotted curve across and to right of each velocitydepth profile is the mean crustal P-wave velocity of Christensen & Mooney (1995) (only available from 5 to 50 km depth), with horizontal bars showing 1or from the mean at each depth, scaled by the mean crustal Vp/Vs ratio from Christensen (1996). Note the normal S-velocities in the shallow crust, but S-wave velocities in the low-velocity zones (shaded) that are one to two standard deviations below expected crustal velocities. Dashed lines at 75 km and 80 km in the Lhasa terrane and 70 km in the Qiangtang terrane are likely Moho depths, and highlight the unusually low Sn velocity of the Qiangtang terrane, <4 km s-1.

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channel apparently disappears south of the IndusYarlung suture (Kind et al. 1996) or becomes shallower and is restricted to the upper half of the crust (Cotte et al. 1999) (Fig. 3). Yuan et al. (1997) suggest that the strongest low-velocity zone is collocated with the seismic bright spots imaged in the Yangbajain graben (IN-2: Brown et al. 1996; Makovsky & Klemperer 1999), but also find a possibly analogous feature beneath Lhasa, confirming that the unusual low-velocity structure of southern Tibet is not restricted to the active extensional grabens. Figure 3 shows one possible correlation of different crustal low-velocity zones on the different surface-wave velocity profiles, and my inferred crustal flow directions (see below). The unusually low crustal Vp (e.g. PASSCAL: Owens & Zandt 1997) can be understood without the necessity of forming Tibet from abnormal crustal compositions if lower (and perhaps also middle) crust has been lost from Tibet by lateral flow (Haines et al. 2003), and if there is a small percentage of partial melting in the crust. Even though Vp in Tibet is unusually low, Vp/Vs is variously determined as being normal or high, potentially indicative of low-degree partial melts in the crust (PASSCAL: Owens & Zandt 1997; Rodgers & Schwartz 1998; Vergne et al. 2002; IN-3: Tian et al. 2005). Locally in the Yangbajain graben of the Lhasa terrane, large S delay times relative to P delay times also seem to require several per cent of melt averaged over the entire crustal thickness (SF-3: Him et al. 1995). Eclogitization could also reduce measured crustal Vp by placing the Moho above eclogitized lower crust (Haines et al. 2003) but may be precluded below the Qiangtang terrane by higher mantle temperatures there. Partial eclogitization is the likely cause of observed unusually high lower-crustal velocities (c. 7.4 km s -1) beneath the High Himalaya (HIMNT: Schulte-Pelkum et al. 2005) and subMoho velocities >8.4 km s i where the Indian craton is subducting beneath the Tethyan Himalaya (SF-I: Henry et al. 1997; Sapin & Him 1997). Gravity modelling suggests that eclogitization may continue to develop within the Lhasa terrane lower crust (Cattin et al. 2001). Eclogitization might also occur at the northern margin of Tibet where Asian lithosphere may be subducting beneath Tibet (Kind et al. 2002; Shi et al. 2004). Mafic lower crust seems to be absent north of the Kunlun Fault (SF-4 and -6: Vergne et al. 2002) and a 'crust-mantle mix' zone with Vp = 7.37.8 km s -1 above a mantle velocity of 8.2 km S - 1 beneath the Altyn Tagh (TQ: Zhao et al. 2006) might represent partial eclogitization. Miocene adakitic lavas - globally rare, and thought to represent melting of garnet-bearing sources, hence presumably eclogitized crust - are now known from both

the southern Lhasa terrane (Chung et al. 2003) and the Songpan-Ganzi terrane (Wang et al. 2005). A fundamental parameter that is well-determined by seismic velocity-depth models is crustal thickness, which varies from 70-80 km in southern Tibet to 6 0 - 7 0 k m in northern Tibet (e.g. WT: Kong et al. 1996; IN-2, -3, PASSCAL: Kind et al. 2002). Some of the earliest controlled-source experiments in Tibet were used to infer large and abrupt Moho offsets beneath the Himalaya, IndusYarlung suture and Banggong-Nujiang sutures (SF-1, -2, -3: Him et al. 1984a, b). These Moho steps were interpreted using the fan-profiling method in which there are large ambiguities between velocity and structure, and have largely been discounted by subsequent workers (Molnar 1988) and in the light of new data (IN-3: Zhao et al. 2001; HIMNT: Schulte-Pelkum et al. 2005). However, at the northern margin of Tibet more robust observations of abrupt changes in crustal thickness suggest abrupt changes in crustal rheology. There is good evidence from receiver functions for a large crustal thickness decrease of 10-20 km northwards across the Karakax-Altyn Tagh fault over a lateral distance <20 km (TB: Kao et al. 2001; SF7: Wittlinger et al. 2004). A similar abrupt offset and evidence for southward underthrusting of the Tarim block is also seen across the Altyn Tagh fault where it separates the Tarim and Qaidam basins, from both teleseismic tomography (SF-5: Wittlinger et al. 1998) and wide-angle refraction profiling (TQ: Zhao et al. 2006). Likewise, Moho offsets are also seen beneath the Qaidam border fault on the PASSCAL, SF-4 and SF-6 profiles (Zhu & Helmberger 1998, Vergne et al. 2002), suggesting that a strong Qaidam basin fills the role of the Tarim basin further west, failing to thicken in response to collisional stresses, and marking the northern limit of major Tibetan crustal thickening. This implied strength of the Tarim and Qaidam basins is also seen in their elastic thicknesses, Te = c. 50 km, that significantly exceed values on the Plateau where Te is typically < 2 0 km (Braitenberg et al. 2003; Jordan & Watts 2005). From receiver-function profiles along SF-4 and -6 Vergne et al. (2002) suggest an additional but potentially much smaller Moho offset below the Jinsha suture. Galv6 et al. (2002a) interpret wide-angle refraction profile SF-6 to show that the largest crustal thickening occurs beneath the Kunlun fault, but a lack of continuous reflection data makes this argument less compelling than the data demonstrating a more abrupt Moho step further north. However, these data (Galv6 et al. 2002a; Vergne et al. 2002) arguing for a gradual transition in crustal thickness across northeastern Tibet may correspond to a region of somewhat elevated Te SE of Qaidam Basin (Jordan & Watts 2005).

CRUSTAL FLOW IN TIBET Seismic attenuation: evidence for high temperatures and partial melts

Theoretical models and laboratory measurements show an increase of seismic attenuation (or decrease in its inverse, Q) with increased temperature, and particularly with the presence of partial melt. Unusually high attenuation was one of the first geophysical anomalies recognized in Tibet, both in the crust (Ruzaikin et al. 1977) and in the upper mantle of northern Tibet (Barazangi & Ni 1982; Ni & Barazangi 1983). More recent waveform modelling studies have refined the degree and extent of these anomalies. The region of strong Sn (upper mantle) attenuation found by Ni & Barazangi (1983) and by McNamara et al. (1995, from PASSCAL) also corresponds with unusually strong crustal attenuation, detected by its effect on both Pnl phases (Rodgers & Schwartz 1998, from PASSCAL) and on Lg (trapped postcritical crustal shear-wave) (Fan & Lay 2003, from permanent broadband stations). Both Rodgers & Schwartz (1998) and Fan & Lay (2003) suggest that the high attenuation is a manifestation of widespread partial melting of the crust of northern Tibet. Regional variability in Lg (crustal) Q from 80 to 100~ across Tibet seems no larger than uncertainties in its estimation (Fan & Lay 2002). Southern Tibet has a less attenuating mantle (Ni & Barazangi 1983; McNamara et al. 1995) and crust (Fan & Lay 2003) on a regional scale. The strongest attenuation (Q = c. 60) yet reported from Tibet comes from Lg measurements along the INDEPTH-2 profile in the Yangbajain graben, believed from the lack of frequency dependence to be due to intrinsic attenuation and not to scattering, so presumably due to hydrothermal and magmatic fluids in the upper crust (Xie et al. 2004). Q increases southwards to c. 100 between the Indus-Yarlung suture and the Kangmar Dome, and to c. 300 beneath the High Himalaya (cf. typical values of c. 200 in tectonically active western North America and >650 for stable central and eastern North America: Xie et al. 2004).

High conductivity zones: evidence for fluids and crustal weakening

Natural-source magnetotelluric (MT) studies in the Sino-French collaboration of the early 1980s demonstrated anomalously high conductivities at shallow depths north and south of the IndusYarlung suture zone, both within the Yangbajain graben and outside the Yadong-Gulu rift beneath Yamdrok Tso (Pham et al. 1986). The high conductivities were interpreted as magma (Pham et al. 1986) because of the correlation with a region of

49

anomalously high heat flow (Francheteau et al. 1984; Jaupart et al. 1985). A similar correlation exists between very high electrical conductivity and the inferred magmatic source of the highest heat-flow region of NW India (Harinarayana et al. 2004). From 1995 to 2001, INDEPTH recorded c. 1600 km of broadband and long-period MT data along six profiles that together cross the entire Tibet Plateau in a corridor from 89 to 94~ (Wei et al. 2001; Spratt et al. 2005), corroborating the older results and providing much new information on the existence of fluids at depth in the crust. Two additional MT transects cross the NW Himalaya at 78~ (himprobe: Gokarn et al. 2002) and the Nepal Himalaya at 85~ (nf: Lemonnier et al. 1999), the latter continuing into Tibet at the same longitude (in-800: Unsworth et al. 2005; wt: Kong et al. 1996). Three salient features emerge from the combined data-sets: (1) generally high crustal conductance (integrated conductivity over depth; a far better resolved parameter than the absolute conductivity at any depth) across the entire Tibetan plateau (in-100, -500, -600: Wei et al. 2001; Unsworth et al. 2004) (Fig. 4); (2) extraordinarily high conductivity below 15km in the Yangbajain graben, southem Lhasa terrane (in-100 and -200: Li et al. 2003), collocated with the seismic bright spots discussed below; and (3) a remarkable continuity of conductivity structure along the entire Himalayan arc from 77 to 92~ (himprobe, nf/in800, in-100, in-700: Unsworth et al. 2005). Wei et aI. (2001) show that along the main INDEPTH transect (in-100, -500, -600), from the Himalayan crest to the Kunlun fault, crustal conductance everywhere exceeds typical values for the continents (Fig. 4). North of the Kunlun fault, and south of the STD, conductance declines into the normal range. In the Tethyan Himalaya and southern Lhasa terrane (in-100) the highest conductivities are seen above 30 km depth, and the conductivity anomaly may be confined to the crust; in the northern Lhasa terrane, Qiangtang and SongpanGanzi terranes (in-500, -600) the highest conductivities are seen below 3 0 k m depth. In two regions of Tibet conductance is above 104 S, exceeding the normal upper bound by more than an order of magnitude (Wei et al. 2001): (1) across the Indus-Yarlung suture and in the southem Lhasa terrane (29-30~ probably associated with the seismic bright spots, and probably confined to the crust; see also Li et al. 2003); and (2) locally in the east-central Qiangtang terrane (in-600 but not in-500, at 33-34~ east along strike from abtmdant Late Miocene and younger xenolith-bearing volcanic rocks (discussed below; Turner et al. 1996; Ding et al. 2003), and certainly extending into the upper mantle (see also Unsworth et al. 2004). The only sensible explanation for the

50

S.L. KLEMPERER

Fig. 4. Lithospheric electrical conductivity of Tibetan Plateau, modified after Wei et al. (2001). Top: conductivity models for in-100, -500, -600 (see Fig. 1). Data gap exists at 30.5~ profiles are offset at 33~ (vertical dashed line) where in-600 is 150 kin east of in-500. Horizontal dashed line at 20 km depth emphasizes typically resistive upper crust except in Yangbajain graben and beneath Tethyan Himalaya. Note in-500 lies outside, and northern part of in-100 lies within, the Yangbajain graben thereby showing that high lower-crustal conductivity in the Lhasa terrane is not only related to the Neogene east-west extension. Horizontal dashed line at 70 km depth marks probable minimum Moho depth south of the Banggong-Nujiang suture to emphasize the likely resistive mantle in southern Tibet, and marks the probable maximum crustal thickness north of the Banggong-Nujiang suture and so emphasizes the likely conductive mantle of the Qiangtang terrane. Bottom: total conductance to 100 km depth. Grey-shaded region is range of typical crustal conductance, 10-1000 S and horizontal dashed line is mean Phanerozoic lower-crustal conductance of 400 S (Jones 1992).

high conductance beneath Tibet is interconnected fluid along the grain boundaries of the rock (Wei et al. 2001), either partial melt or saline aqueous fluids. Wet rocks and partially molten rocks are both far weaker than their dry counterparts (Kirby 1984; Rosenberg & Handy 2005), thus fluid-rich crust, irrespective of the fluid type, must represent weaker areas able to accommodate deformation by crustal flow. The high observed conductance implies unusually weak crust below 30 km throughout Tibet, with the additional implication of weak upper crust in southern Tibet and weak upper mantle in northern Tibet. Li et al. (2003) provide detailed models for the Yangbajain graben region that satisfy both the high conductivity (Chen et al. 1996) and the seismic bright spots (Brown et al. 1996; Makovsky & Klemperer 1999). The MT data alone cannot distinguish between end-member models of a > 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 <10 km distance requires the anisotropy to be at least partially in the crust. (b) Anisotropic crustal velocity model at AMDO from receiver-function inversions, from Sherrington et al. (2004): vertical lines show minimum, average and maximum velocities for anisotropic layers. Numbers beside each layer are trend/plunge of fast direction and percentage anisotropy for each of five anisotropic layers. (c) Radial anisotropy in the Tibetan crust from Shapiro et al. (2004), showing on the left the surface-wave dispersion inversion for a point in western Tibet (34~ 84~ and the fit to the observed dispersion curves using the 8% radial anisotropy in the middle crust shown on the right. Depth scale is identical in (b) and (c).

known only from within the graben, perhaps merely because seismic-reflection coverage outside the graben is as yet sparse (Haines e t al. 2003; Ross e t al. 2004): seismic surface-wave studies indicate that similar low-velocity zones exist both within and beyond the graben (discussed above: Yuan e t al. 1997). The lack of correlation of the inferred region of partial melt with the young north-south grabens (which may be relatively superficial features, confined to the hanging wall of the STD: Cogan e t al. 1998; Hurtado e t al. 2001) is further demonstrated by the essential similarity of four widely spaced MT transects along the Himalayan arc (himprobe, nf/in-800, in-100, in-700: Unsworth e t al. 2005). This similarity strongly suggests that the remarkable conductivity of southern Tibet first recognized by Pham e t al. (1986) may be extrapolated laterally over 1000 km, counter-claims (Harrison 2006) notwithstanding. In all four transects, high mid-crustal conductivities clearly extend south of the IndusYarlung suture into the Tethyan Himalaya, to the North Himalayan gneiss domes (including Kangmar dome in the east, Tso Morari complex in the west: Fig. 1). The top of the conductivity

anomaly appears to shallow to the south, and in the best-resolved eastern transects the anomalous region also thins to the south. This clear geometric pattem is explicitly interpreted by Unsworth e t al. (2005) as the signature of a southward-directed channel flow; they show that reasonable assumptions about the degree of partial melt required to produce the observed conductivity result in strength reduction by a factor of about four. Whereas the highest conductivities in southern Tibet are most reasonably explained by brines in conjunction with magma, in northern Tibet it has previously been suggested that brines do not play a significant role (Wei e t al. 2001; Unsworth e t al. 2004). The xenoliths that are our best record of the lower crust in the Qiangtang terrane had anhydrous mineral assemblages resistant to widespread crustal melting at the time of eruption, 3 million years ago (Hacker e t al. 2000). Thus either brines or melts in the lower crust are likely to be derived from below, and any saline fluids have probably infiltrated the crust since 3 Ma. Irrespective of the availability of brine or melt, fluids must form a physically interconnected network to be conductive, requiring a fluid-mineral dihedral angle < 6 0 ~ for

52

S.L. KLEMPERER

typically low lithospheric fluid contents. Despite previous statements that aqueous fluids will not connect to form conductive pathways at the depths of the northern Tibet conductors (Wei et al. 2001), the unusual thickness of Tibetan crust (>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 <0.5 s due to the long dominant periods (>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

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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 <6.2 km s -1 (Makovsky et al. 1996) precludes a large number of mafic intrusions, so the reflectivity seems most likely to represent older flow-related flattening structures. Some

indirect temperature

estimates

Direct measurements of crustal heat flow and inference of high lower-crustal temperature are corroborated by a variety of less direct approaches. Alsdorf & Nelson (1999) model the pronounced satellite magnetic low observed over Tibet as indicating that the Curie isotherm (c. 550~ resides in the upper crust (c. 15 km depth) across the entire plateau. Mechie et al. (2004) believe they can identify the c~ - / 3 quartz transition as a seismic reflector along the INDEPTH-3 transect, thereby constraining the temperature to be 700~ at 18 km depth in the southern Qiangtang terrane, and 800~ at 32 km depth in the northern Lhasa terrane. Hacker et al. (2000) describe xenoliths from the Qiangtang terrane that equilibrated at >_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 <750~ and upper mantle beneath the northern plateau in the range 950-1100~ (Ruppel & McNamara 1997). Although the abundant Late Miocene to Recent potassic volcanics of northern Tibet have been interpreted as representing high temperature

56

S.L. KLEMPERER

(c. 1100~ melting of the subcontinental mantle, this temperature was probably not attained above c. 125 km depth, well below the Moho (Turner et al. 1996). Miocene (but none yet younger than 8 Ma) potassic volcanics are also known from southern Tibet (Ding et al. 2003), representing lower degrees of partial melt (Williams et al. 2004), therefore presumably lower melting temperatures than in northern Tibet, and are coupled with calc-alkaline adakites of the same age range that may represent partial melts of an eclogitic lower crust (Chung et al. 2003). The youngest igneous rocks in southem Tibet (away from the syntaxes) are c. 7 Ma leucogranites (Li et al. 1998) and though these only indicate relatively shallow (600800 MPa), relatively low-temperature (750800~ melting (Patino Douce & Harris 1998), the time for conductive cooling of material at c. 20 km depth is > 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

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(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

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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.

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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.

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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

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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

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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

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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.

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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

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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

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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,

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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).

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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)

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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

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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

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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

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homogeneous channel flow, LHO-2 illustrates Mode 2, heterogeneous channel flow, and LHO-3 illustrates Mode 3, hot fold nappes. The models are identical except for the properties of the lower crust, as described below.

Model LHO-1: homogeneous channel flow Model LHO-1 is a typical Mode 1 laterally homogeneous model with a uniform 10 km thick lower crust with model rheology B*(DMD/5). This scaling achieves an effective strength that is intermediate between very strong diabase (B*(DMD)) and intermediate granulite (e.g. Pikwitonei granulite (Mackwell et al. 1998) with effective strength B*(DMD/IO)). The pro- and retro-sides of the model (Figs 3-6, parts 1 & 2, respectively) indicate how the two sides of the model evolve with and without surface erosion, respectively. Figure 3 shows the material distribution and the deformation of a passive, initially rectangular, Lagrangian marker mesh for the pro- and retro-sides of the model. The bold vertical mesh lines are numbered relative to the surface suture (the initial boundary between the pro- and retro-sides of the model), labelled '0' and located above the model S point. Figure 4 shows the corresponding distribution of crustal radioactive heating and the temperature field. Figure 5 shows the distribution of the second invariant of the stress, and Figure 6 shows the velocity, plotted as vectors, and the second invariant of the strain rate field. Convergence is symmetric with Vp = 1.5 cm a -i, VR = - 1 . 5 c m a -1 and Vs = O. During the initial 25 million years the main style of deformation, shown by the velocity vectors and Lagrangian marker grid, is characterized by diachronous near-pure-shear thickening of the upper- and most of mid-crust, the development ofa subhorizontal shear zone near the base of the mid-crust, and the viscous decoupling of the relatively weak lower mid-crust from the stronger B*(DMD/5) lower crust (Figs 3b & 6b, 20 My). The lower crust is weakly sheared and thickened where the basal boundary condition forces it to detach near the centre of the model. This effect is probably not realistic; Beaumont et al. (2004) argued that lower crust is most likely subducted during orogenesis because orogenic antiformal cores comprising thickened lower crust, like that seen here, are not observed in natural orogens. However, lower crust is not subducted in this model to be consistent with the next two models. The temperature field is closely linked to the evolving distribution of heat-producing material. During diachronous crustal thickening there is some radioactive internal self-heating, but significant thermal disequilibrium remains owing to

vertical advection of the temperature field during crustal thickening (Fig. 4b & c, 20 and 30 My). Thermal re-equilibration, by radioactive selfheating and thermal diffusion, occurs with a timescale of close to 20 million years during which time the temperature in the lower crust reaches 800~ (Fig. 4b). This self-heating timescale is much shorter than the 50-200 million years required for lithospheric-scale thermal relaxation. At approximately 25 million years, channel flow starts in the melt-weakened retro-mid-crust, where T > 750~ and is soon followed by an equivalent flow in the pro-crust (Figs 3c & 6c, 30 My). Although the flows develop against the thickened lower crust, this strong antiform does not cause the channel flow by acting as a backstop. Similar channel flows occur in models where the strong core is absent (Beaumont et al. 2004; Jamieson et al. 2006). The minor asymmetry in the flow results from erosion on the pro-side of the model. The oppositely directed channel flows subsequently tunnel outward such that their tips evolve with the temperature field, coinciding with the 750~ isotherms (Figs 3, 4 and 6), which are also close to the edges of the orogenic plateau that develops in the centre of the model (Figs 3 and 4, 30-60 My). Channel flow is restricted to the region beneath the plateau and does not penetrate into the foreland crust, which is too cold. The only significant difference between the two sides is the erosional uplift and exhumation of the pro-flank, which causes tectonic thickening of the mid-crust but is not sufficiently intense to exhume the channel, which continues tunnelling. Flow in the channel beneath the plateau reaches velocities of approximately 0.75 m a - l ; strain rates exceed 10 -13 s -1 in the boundary layers but the second invariant of the stress in the channel does not exceed 1 MPa. This model illustrates homogeneous channel flow (Figs 3 and 6).

Model LHO-2: heterogeneous channel flow In many orogens the crust of the colliding continents may be heterogeneous. This is almost certainly true in the Himalayan-Tibetan orogen, where the Indian and Asian crusts have different compositions and, moreover, the earlier accretionary history may have given the Asian crust considerable internal heterogeneity. Although we have not undertaken an exhaustive sensitivity analysis of the effects of crustal strength variations, we have a range of model results that include upper, mid- and lower crustal heterogeneities. Model LHO-2 (Figs 7-10) provides some insight into the effect of variations in lower crustal properties on

CRUSTAL FLOW MODES the thermal-tectonic style of the models. We focus on the relative styles of deformation of the mid- and lower crust and their differences compared with homogeneous lower crust, LHO-1. The only difference between models LHO-1 and LHO-2 is that the lower crust in the interior of LHO-2 comprises alternating 250 km wide zones with B*(DMD) and B*(DMD/IO) rheologies. The external parts of the model have lower crust with B*(DMD). The high- and low-viscosity regions in the lower crust therefore have a nominal viscosity contrast of 10, designed to correspond approximately to the difference between dry, refractory mafic lower crust and intermediate granulite lower crust (e.g. Pikwitonei granulite, see above). However, this factor of 10 contrast is modulated by the non-linear effect of power-law flow and temperature variations. The strong lower crustal blocks are therefore nominally a factor of 2 stronger than LHO-1 lower crust, and the weak blocks are a factor of 5 weaker. Model LHO-2 results show a complexly deformed crust that can be understood as the superposition of two main deformation phases. Phase 1 activates and deforms the zones of weaker lower crust in the transition zone between the foreland and the plateau (Fig. 7b). The style is very similar to the deformation of a finite-width salt layer as sediment progrades over it (Lehner 2000); the horizontal pressure gradient in the transition zone between plateau and the foreland acts in the same way as the pressure gradient caused by the prograding sediment (Gemmer et al. 2004). It squeezes and evacuates the weak lower crust, then thrusts it and the overlying crust pro-ward on the pro-side, and in the opposite direction on the retro-side, as allochthonous tongues or nappes over adjacent regions of strong lower crust (Fig. 7b-d, 3 0 50 My). Shears at the leading edges of the tongues propagate upward through the crust, and the allochthonous tongues and their overburden become uplifted and transported. Where lower crust is evacuated it is replaced by subsiding midcrust, and these regions preferentially shorten and thicken during further contraction (e.g. vertical markers - 2 to - 3 and - 4 to - 5, Fig. 7b). In Phase 2, a channel flow develops in the heterogeneous crust created in Phase 1. The tongues of overthrust weak lower crust become entrained in the channel flow (Fig. 7d & e, 5 0 - 6 0 My). The remaining zones of strong lower crust are transported into the centre of the plateau and detached at S, where they are incorporated into an antiformal stack (Fig. 7 b - e , 3 0 - 6 0 M y ) similar to that in LHO-1. The temperature distribution, redistribution of radioactive crust, velocity field, strain rate, and stress in LHO-2 are different from those of LHO- 1,

117

but less so than the deformation (Fig. 7) might suggest. Channel flow (Figs 7 and 10) develops beneath the plateau in both cases. The implication is that heterogeneous lower crust may make the geometry and composition of the channel flows similarly heterogeneous, e.g. in the heterogeneous flow mode the channels may transport detached lumps of much stronger, distinctly different composition, high metamorphic pressure, granulitic or eclogitic lower crust. Widespread channel flows can develop even under these circumstances, provided that the viscosity of most of the mid-crust becomes sufficiently low and the lumps are not too large to be transported.

Model LHO-3: hot fold nappes The evolution of a representative model, LHO-3, designed to test the response of an orogen to collision with successively stronger blocks of lower continental crust is shown in Figures 11-14. The model is symmetric except that one flank of the orogen is mildly denuded by slope-dependent erosion and the other is not. The upper and midcrust are uniform and the only lateral variation in properties comes from the 250 km long, 15 km thick lower crustal blocks in which the effective power-law viscosity, based on Dry Maryland Diabase B*(DMD), is successively reduced by factors of 4, 8, 12, 16 and 20 toward the centre of the model from both sides (Table 1). This scaling creates effective viscosities ranging from B*(DMD) in the external crust, through B*(DMD/IO) (intermediate Pikwitonei granulite; Mackwell et al. 1998) to half this value, B* (DMD/20), in the centre of the model. The entire lower crust has ~bey = 15~ but this is not important because deformation occurs in the ductile regime. The model is highly idealized, and is designed more as a physics/mechanics experiment to test how different strength lower crustal blocks will be absorbed by the model orogen system than as an attempt to model a natural system. In this experiment the blocks that are inserted become progressively stronger with time. The experiment determines when lower crustal blocks appear to be weak, and therefore deform and are incorporated into the orogen or, in contrast, when they are strong and act as indentors. The model represents a development of the vice-type models described by Ellis et al. (1998). The model exhibits a three-phase evolution. During Phase 1 convergence, the crust containing the weaker lower crustal blocks diachronously shortens and thickens by nearly uniform contraction in the upper and mid-crust (Fig. 1 lb, parts 1 & 2, 30 My). A ductile shear zone develops at the base of the crust, detaching the overlying weak lower

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crustal blocks from the basal boundary condition that represents kinematically underthrusting mantle lithosphere (Figs l lb & 14b, 30 My). Diachronous thickening of the radioactive crust is relatively fast and creates thermal disequilibrium owing to the vertical stretching (Fig. l lb, 30 My). This thermal disequilibrium is reduced during Phase 2, a period of radioactive self-heating and thermal relaxation that produces hot, ductile lower crust, highly ductile mid-crust and a relatively cool, strong, frictional-plastic upper crust (Figs 12 & 13, b & c, 30-40 My). Phase 2 is also diachronous and typically takes c. 20 My after crustal thickening ends (Fig. 12c, 40 My). Phases 1 and 2 occur sequentially as weak lower crustal blocks are inserted, thickened, absorbed and heated, as the model orogen becomes progressively wider and hotter. The onset of Phase 3 coincides with the arrival and underthrnsting of a lower crustal block that cannot be absorbed by Phase 1-style deformation because it is too strong and resists decoupling. This effect is initially progressive - blocks with rheology based on B*(DMD/20), B*(DMD/16) and B*(DMD/12) decouple easily and there is no significant change in deformation style. However, the B*(DMD/4) block offers some resistance to decoupling, forcing additional contraction on the interior of the system which responds by developing large-scale lower-crustal folds (Fig. l lc, 40 My). The transition to Phase 3 becomes fully developed with the arrival of the B*(DMD) lower crust. It does not decouple and, consequently, acts as an indentor/plunger that forces weak middle and lower crust into large-scale, gently inclined, ductile fold nappes rooted at the Moho (Fig. l ld, 50 My). Some of these are then expelled over the indentor and either inserted into the middle crust (Fig. 1 le, part 2, 65 My) and/or exhumed to the surface by erosion (Fig. 11, part 1, d & e, 50 and 65 My). Surface denudation during Phase 3 determines the relative amount of uplift and exhumation of the fold nappes versus their horizontal transport once inserted into the mid-crust (pro- versus retrosides; Fig. 11). If there is little or no erosion (retro-side), the nappes remain buried and are transported together with the overlying crust, which shows little deformation associated with nappe insertion (Fig. l le, part 2, 65 My). As explained below, the hot fold nappe style of crustal flow is favoured by weak lower crust in the interior of the orogen. The extent of weakening is related to the incubation time, the duration of Phase 2 for each part of the model crust (see discussion below).

Evolution of topography in LHO models The evolution of the topography in models LHO-1 to LHO-3 is shown (Fig. 15) for two possible

isostatic balances in which the density difference between the crust and mantle is either 500 or 600 kg m -3. The 500 kg m -3 results are comparable to natural orogens, and predict average plateau elevations of approximately 5500 m for models LHO-1 and LHO-2. Model LHO-3 has a higher mean plateau elevation because the strong lower crust is thicker than in the other two models. In all three cases the topography has the triangular initial shape expected for small, bivergent, critical wedge-type orogens. In LHO-I and LHO-2 this geometry grows self-similarly during the first 20 million years, also as expected for critical wedges, and the maximum elevation reaches 7 to 8 km. After c. 20 million years, radioactive self-heating is sufficient to weaken the mid-crust progressively in the centre of the orogen and the geometry evolves to a central plateau flanked by younger, stronger critical wedges. In LHO-1 and LHO-2 the elevation of the plateau is lower than the local high in the centre of the orogen, which correlates with the antiformal core of strong lower crust (Figs 3 & 7). As noted above, this core does not form if lower crust is subducted. In LHO-2, the plateau evolves less uniformly than in LHO-1 because the topography is sensitive to the absorption of the alternating strong and weak lower crustal blocks. LHO-3 differs from the other two models because the lower crustal blocks within the orogen are initially weak, giving low-taperangle bivergent critical wedges (e.g. Fig. 15, 15 My). Later the topography takes the form of a plateau 'bookended' by high-taper-angle strong lower-crustal wedges (e.g. Fig. 15, 37-65My). The strength of these bounding critical wedges maintains a higher, narrower plateau in LHO-3 compared to the other models.

Upper-mantle-scale models The models described in the previous section treat the coupled thermal-mechanical deformation of the crust in a self-consistent manner subject to the assumed basal kinematic velocity boundary conditions. These conditions can be interpreted either as symmetric convergence and subduction of the two mantle lithospheres (Figs 6, 10 & 14), or as advancing subduction of the pro-mantle lithosphere coupled to the retro-mantle lithosphere (Fig. 2), a style also referred to as ablative subduction (Tao & O'Connell 1992; Pope & Willett 1998). Is this prescribed subduction of the mantle lithosphere dynamically consistent? We describe here two upper-mantle-scale (UMS) models to show that subduction is dynamically consistent for two particular sets of mantle lithosphere properties. An investigation of the range of parameter values for

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LHO-1

7 6

E5

-scale for _ A 9 = 500 kg/m 3

scale for Ap = 600 kg/m 3

--

m

v

4 t~

~3 2

b

-500

6 S-point

500

--, 0 1000 Distance, x (km)

LHO-2

6

scale for _ Ap = 500

,~.

scale for kg/m 3

_

v

~t~ 4 -

2 1

0

LHO-3

v

~t:n 4

0 S-point

500

scale for

7 6

-500

1000 Distance, x (km) scale for

!

1:

2 1 0 -500

() S-point

500

1000 Distance, x (km)

Fig. 15. Evolution of the topography for models LHO-1 to LHO-3 shown with respect to the S point, x = 0 (My = millions of years). The height scale is shown for two representative isostatic balances which depend on @ = Ap/pm, where Ap = Pm Pc and Pm and Pc are the mantle and crustal densities, respectively. The scale of the height is most sensitive to Ap and the results are therefore shown for Ap = 500 and 600 kg m 3, corresponding to @ = 0.156 and 0.182, respectively. -

-

which the models exhibit subduction will be published elsewhere. Upper-mantle-scale orogenic thermal- mechanical models with viscous-plastic rheologies have been presented by Pysklywec (2001), Pysklywec et al. (2000, 2002) and Pysklywec & Beaumont (2004). Rayleigh-Taylor (R-T) instabilities in mantle lithospheres with linear and non-linear viscosities have been investigated by, for example, Conrad &

Molnar (1997), Houseman & Molnar (1997), Molnar et al. (1998) and Neil & Houseman (1999). The results of the viscous-plastic experiments demonstrate several modes of mantle lithosphere deformation including subduction, double subduction and slab breakoff, in addition to the viscous R-T dripping. However, most of this work focused on the early stages of continent-continent collision. Here, our concern is flow modes in large hot orogens, therefore

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we describe two examples that illustrate the types of crustal flow and mantle lithosphere behaviour that may occur during prolonged continent-continent collision.

Description of UMS model experiments Both models are two-dimensional and the domain is 2000 x 600 k m (Fig. 16). The A L E finite-element techniques and the reasons for the choice of model properties are explained in the numerical calculation section and the Appendix. The models include the lithosphere and upper mantle (Fig. 16, Table 2) and are laterally uniform except for a narrow weak zone in the crust and uppermost mantle designed to represent a simplified suture that localizes the initial deformation (Fig. 16). No precursor phase of oceanic subduction is considered in these experiments. The lithosphere boundary

conditions are specified at the sides of the model domain and are designed to correspond approximately to the collision of India with Asia. Prolithosphere, equivalent to India, converges from the left at a uniform velocity (in both depth and t i m e ) , Vp = 5 c m a - t , against a stationary retrolithosphere, VR = 0 cm a-1, at the fight boundary corresponding to Asia. In contrast to the CS models described above, velocity is not specified at the base of the crust, nor anywhere inside the model. The subduction advance/retreat velocity, Vs, is instead determined by the dynamical evolution of the model. The sides and base of the model have flee-slip boundary conditions and a small, uniform, symmetric, outward leakage flux of material is specified through the sublithospheric mantle parts of the side boundaries to balance the flux of pro-lithosphere into the model (Table 2). No surface erosion or deposition

Fig. 16. Configuration and principal properties of the upper-mantle-scale models, LHO-LS 1 and LHO-LS2 (see also Table 2). These models are the same except for the minor difference in the reference densities of their respective mantle lithospheres. Notation tb = 15 ~ --~ 2 ~ implies strain-softening of the internal angle of friction in this case over the range of strain of the second invariant of 0.5 to i.5. Effective viscosity 7: B*(WQx 5) (Wet Quartz rheology, scaled by 5); B*(DMD/IO) (Dry Maryland Diabase rheology scaled down by 10); B*(WOI x 10) (Wet Olivine rheology scaled up by 10), p = density given at reference temperatures; thermal coefficient of volume expansion = 3 x 10 -5 ~ ~. Note lower crustal density change corresponding to the 'basalt-eclogite' metamorphic phase transition. Model domain is 2000 • 600 km and comprises the lithosphere, thickness 100 km and sublithospheric mantle. Lithosphere converges asymmetrically from left at 5 cm a -~ . Boundary conditions on sublithospheric mantle are free slip with no material flux across the base. The sides have a small uniform outward material flux that balances the flux of lithosphere into the model. White region at left is the initial narrow weak zone. There are no surface processes. Bold frame shows area displayed in Figures 17 and 18; note that the position of this frame migrates with time in these figures.

CRUSTAL FLOW MODES occurs in these models, which are designed for comparison with the simple tunnelling mode of homogeneous channel flow (Beaumont et al. 2004, figs 12a and 13a). This approach was chosen in order to focus on the effect of the mantle lithosphere behaviour on the crustal channel flow. The mechanical and thermal properties of the UMS models are described in the numerical calculation section and in Table 2; key properties are summarized in Figure 16. In particular, they include frictional-plastic flow, power-law creep, crustal radioactive heating, melt-weakening, and the granulite to eclogite phase transition in the lower crust. The crust is similar to that in the CS models, except that the frictional-plastic rheology strainsoftens from ~beff= 15 ~ to 2 ~ over the range 0.5 to 1.5 of the second invariant of the strain. The mantle lithosphere strain-softens in the same manner. Lower-crustal density changes from 2950 to 3 1 0 0 k g m -3 during the granulite-eclogite phase transition. This density increase is chosen to be relatively small because only a fraction of the crust is considered to transform to high-density eclogite. The scaled power-law creep parameters for the model layers are given in Figure 16 and are discussed above and in the Appendix. Density varies among the model layers and with a volume coefficient of thermal expansion of 3 x 10 -5 ~ -1. The only difference between the two models is in the reference density of the mantle lithosphere, which is 3 3 0 0 k g m -3 in model LHO-LS1 and 3 3 1 0 k g m -3 in model LHO-LS2, resulting in a nominal average density difference between the mantle lithosphere and sublithospheric mantle of 40 and 30 kg m -3, respectively. The results show that the model behaviour is very sensitive to this 0.3% difference in mantle lithosphere density.

UMS model results: models LHO-LS1 and LHO-LS2 Model results are described using the 'pro- retro-' terminology because the dynamic behaviour is similar to the prescribed subduction in the CS models. During the initial stages of convergence in both models (e.g. Fig. 17a for LHO-LS1) the mantle lithosphere asymmetrically underthrnsts and subducts at a relatively low angle in a 'platelike' manner with little internal deformation. By 9 My, however, the behaviours of the subducted slabs diverge. The denser mantle lithosphere in LHO-LS2 begins to sink, in addition to subducting, and the lower part of the slab steepens and dips at high angle (Fig. 18a). In contrast, the slab in LHOLS1 resists subduction and the retro-mantle lithosphere deforms to accommodate the contraction.

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This subsequently develops into advancing (Fig. 17b) and then double subduction (Fig. 17c), during which the subduction point advances dynamically leading to a net subduction zone advance of approximately 700 km between 9 and 33 million years, corresponding to an average Vs = 3.2 cm a -~. At approximately 30 million years, the buoyancy of the double slab becomes sufficiently negative that viscous necking starts, leading to break-off of the double slab at 42 million years (Fig. 17d), by which time the subduction point has advanced by 900 km at an average velocity of 2.7 cm a 1. This style of advancing subduction, at approximately half the overall convergence rate, is effectively the same as that prescribed in model HT1 (Beaumont et al. 2004), indicating that the prescribed basal velocities are compatible with a dynamical model with properties like LHO-LS1. The detached lump of mantle lithosphere remains in the model domain and tends to circulate upward because it approaches neutral buoyancy as it heats and thermally expands. In nature an equivalent lump may sink into the lower mantle before it approaches thermal equilibrium. LHO-LS1 also develops a mid-crustal channel flow similar to those in equivalent CS models where the channel tunnels outward and is not exhumed by erosion (Beaumont et al. 2004, fig. 1 l a). The main difference from the CS models is that the lower crust does not subduct efficiently but instead tends to accumulate near the subduction point (Fig. 17b-d). Unlike CS model LHO-1 where the lower crust forms a large antiform, the eclogitic lower crust in LHO-LSI pools at the base of the isostatically depressed crust. This difference occurs because the eclogitic lower crust is denser (3100 versus 2700kg m -3) and weaker (B*(DMD/IO) versus B* (DMD/5)) than the lower crust in LHO-1. In contrast to LHO-LS 1, the mantle slab in LHOLS2 is slightly denser and becomes unstable, necks, and breaks off much earlier, between 9 and 12 My (Fig. 18b). The slab is sufficiently dense that it subducts without significant deformation of the retromantle lithosphere. There is, however, still a significant component of subduction-zone advance between 9 and 18 million years. Between 18 and 21 million years, the subducting slab begins to sink such that its motion is vertically downward along much of its length. Sinking is faster than the overall convergence rate and the system changes to subduction-zone retreat, creating a progressively widening region between the slab and the retro-mantle lithosphere that is synchronously filled by the rapid influx of low viscosity, hot (1000 - 1300~ sublithospheric mantle (Fig. 18c). This region widens to approximately 200 km by 27 million years (Fig. 18d). The model therefore displays a combination of subductionzone retreat and mantle delamination. The

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Fig. 17. Model LHO-LS 1 results showing evolution of upper-mantle-scale deformation. Panels show the model materials (see Fig. 16) (dark-grey regions are eclogite-facies lower crust), a sparse version of the Lagrangian tracking grid, the velocity field (arrows, scale at bottom), and selected isotherms (V = H). t = Elapsed model time, ~c = total convergence. No surface processes. Note the progressive 800 km movement of the panel windows toward the left as the model evolves, designed to keep the subducted slabs near the centre of each panel. Crustal channel flow is well developed by 30 million years (My).

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Fig. 17. (Continued)

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Fig. 18. Model LHO-LS2 results showing evolution of upper-mantle-scale deformation. Panels show the model materials (see Fig. 16) (dark-grey regions are eclogite-facies lower crust), a sparse version of the Lagrangian tracking grid, the velocity field (arrows, scale at bottom), and selected isotherms (V ---- H). t ---- Elapsed model time, Ax = total convergence. No surface processes. In this case the 400 km movement of the panel windows as the model evolves is to the right. Crustal channel flow is restricted to the retro-crust but is well developed by 30 million years (My).

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Fig. 18. (Continued)

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delamination of the mantle lithosphere from the crust is very efficient because it creates net subduction-zone retreat despite the continued convergence of the prolithosphere. The delamination velocity therefore exceeds 5 cm a-1. The transition to retreating subduction is markedly different from the basal boundary conditions used in the Himalayan-Tibetan crustal-scale models (Beaumont et al. 2001, 2004). It has more in common with the behaviour envisaged in the Willett & Beaumont (1994) retreating subduction model, except that the polarity is reversed. Despite the different style of subduction, LHOLS2 also develops a mid-crustal channel flow (Fig. 18c & d), in this case confined to the retroside of the system. This restriction occurs because delamination and subduction-zone retreat occur beneath the converging pro-crust so fast that it does not have time to melt-weaken before it is transferred across the migrating subduction point to the retro-side of the system. The overall width of the channel zone is similar to that of LHO-LS 1 although the location of subduction beneath the plateau and region of channel flow is different. LHO-LS2 achieves an end-member geometry in which the subducting mantle lithosphere continuously peels away from the crust beneath the leading edge of the plateau and subducts at one side of the orogen. In LHO-LS1 the subduction zone advances beneath the plateau. In the context of the HimalayanTibetan system, these two results correspond approximately to subduction of the Indian mantle lithosphere respectively beneath the IndusTsangpo and Bangong sutures. The two UMS models illustrate how sensitive the behaviour of the mantle lithosphere may be to small differences in the density contrast between the mantle lithosphere and the sublithospheric mantle. This sensitivity is enhanced by high temperatures in the lithosphere, which render it weak and prone to changes in the style of subduction. The behaviours seen in these models (and others not reported here) may all occur in nature depending on the ambient conditions during continent-continent collision. The results also indicate that crustal channel flows develop during dynamical subduction that exhibits both subduction-zone advance and retreat. This demonstrates that model crustal channel flows are not an artefact of the assumed basal boundary conditions in the CS models. However, the full range of parameter combinations needs to be investigated before drawing additional conclusions.

answered in a general way using the T-M diagrams (Fig. 1) adapted to show where flow modes are predicted in T-M space (Fig. 19). Mid-crustal flows will not occur in small cold orogens because typical quartzo-feldspathic crust has a viscosity that is too high at the ambient temperatures. However, orogens that are rich in limestone and evaporite (e.g. calcite, anhydrite and halite rheologies), which are much weaker than quartzdominated lithologies, may develop these flows in the small cold parts of T-M space. Mode boundaries in T-M space (Fig. 9) are therefore sensitive to the composition of the lithosphere. For example, the flow modes we have described are common in passive-margin salt tectonics (e.g. Lehner 2000; Gemmer et al. 2004), despite the small size and cool temperatures of these systems. For typical quartzo-feldspathic crust, homogeneous channel flows are restricted to the hot regions of T-M space. Their lower limit (Fig. 19) is the threshold at which increasing temperature, which lowers viscosity, and increasing orogen magnitude, which amplifies the gravitational-driving force, combine to drive channel flow. However, because orogen magnitude is limited by the maximum thickness of the crust (c. 70 km), the range of gravitational force is smaller than that of the variation of viscosity. It follows that the gravitational forcing cannot overcome high crustal viscosity and by implication the crest must be very weak, and presumably hot, for channel flows to occur. Volcanic arcs may represent a small hot orogen end-member that could develop localized channel flows even at relatively small magnitude (Fig. 19). The tectonically driven hot fold nappes mode can occur in a much larger part of T-M space, including

Discussion F l o w m o d e s in t e m p e r a t u r e - m a g n i t u d e space

Can we predict which flow modes will operate in different types of orogens? This question can be

Fig. 19. Interpretation of flow modes discussed in this work in regard to the positions where they may operate in T-M space

CRUSTAL FLOW MODES the homogeneous channel domain (Fig. 19). All that is required is a sufficiently hot and thick orogen interior that nappes will be expelled and injected at the mid-crustal level during indentation. The hot fold nappe behaviour illustrated by model LHO-3 is a variation of that produced by the mechanical vice model, which has been applied to the Newfoundland Appalachians (Ellis et al. 1998). Vice-type deformation, where weaker crust is squeezed between 'jaws' of stronger crust, can occur in most of T-M space. Hot fold nappes (Fig. 19), however, form only when the upper parts of the vice jaws are sufficiently weak that they cannot resist expulsion of the nappes when weak material is expelled from the interior over the vice or indentor. Jamieson et al. (2004a) have interpreted part of the Grenville orogen to record the diachronous evolution of hot fold nappes, possibly with superimposed heterogeneous channel flows, during Mesoproterozoic collision on the Laurentian margin. We suspect that similar tectonically driven styles will be recognized in many North American orogens, which developed by successive collisions against and accretion to the cratonic core which acts as an indentor. In particular, the Trans Hudson, central and southern Appalachian, and southern Canadian Cordilleran orogens, and the Archean Slave and Superior cratons, are prime candidates for these flows. The heterogeneous channel flow mode is transitional between the other flow modes described here and has both tectonic forcing, required to activate and evacuate weak lower crust, and gravitational forcing required for channel flow. The fully developed form of this flow mode therefore overlaps with the homogeneous channel flow region of T-M space because both require gravitationally driven flow (Fig. 19). The tectonic evacuation of weak lower crust alone can, however, occur in a larger region of T-M space, which grades into the hot fold nappes domain (Fig. 19). In nature, unlike our two-dimensional models, ductile flow will not be restricted to the direction of convergence. In many situations, flow (sub)parallel to the strike of the orogen may be preferred over expulsion of hot nappes, and channel flow may also be directed around strong enclaves of crust. The inferred outward movement of crust beneath the eastern flank of the Tibetan Plateau (e.g. Clark & Royden 2000) is one example of a three-dimensional flow, and Hatcher & Merschat (2006) describe evidence for Palaeozoic orogenparallel flows in the southern Appalachians.

Effect of thermal relaxation and incubation time on crustal flows Comparing the timescales of external and internal orogenic processes helps to predict the flow styles in the model experiments. We define an external

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timescale, the incubation time, to be the lag time between rapid tectonic thickening of the crust during contractional orogenesis and a subsequent external process, such as indentation, that acts on the system. This definition corresponds to that of England & Thompson (1984) in which erosion was the external process. Note that the incubation time varies with position within the model because tectonic thickening is diachronous. The orogen response to indentation will depend on whether the incubation time is long or short compared with timescales of internal processes, defined below, that are required for radioactive heating and thermal relaxation to achieve particular thermal-rheological thi'eshold states within the crust. In the case of the hot fold nappes mode, let ~'HN be the delay time necessary since rapid tectonic crustal thickening for self-heating and thermal relaxation to achieve the thermal-rheological state required for the hot nappe type of response seen in LHO-3. That is, hot nappes will be the flow style in models like LHO-3 when the incubation time is approximately equal to "rHN. For shorter or longer incubation times the response may be quite different. Although we have not measured ~rm accurately for the models described here, it is related to the thermal relaxation timescale and is estimated to equal the 2 0 - 3 0 million years required for radioactive self-heating to heat the thickened mid-crust to T > 700~ (figs 4b, c, 8b & 12b; see also Medvedev & Beaumont 2006), when ductile nappes will easily form during indentation. Although related, the thermal relaxation timescale and ~'HN differ in that the former is the characteristic timescale that measures decay of thermal disequilibrium, whereas the latter is the time required to achieve a particular thermal-rheological state. Tectonic thickening of cold crust with a low level of radioactivity will be associated with a characteristic thermal relaxation timescale, but the crust may never become weak enough for the hot nappe response, giving the system a finite thermal relaxation timescale but an infinite rI-XN. Such orogens are subcritical with respect to any of the three flow modes described here (Fig. 19). Conversely, crust that is already hot and weak before thickening can have a ri-~ that is zero or only a fraction of the thermal relaxation timescale. Model LHO-3, which creates hot fold nappes, satisfies the condition that the incubation time, which varies from 20 to 50 million years with lateral position at the time of the onset of indentation at 55 million years (Fig. 11), was equal to or greater than ran (20-30 million years). Earlier indentation, at 30 million years for example, may just have caused more crustal thickening. Model LHO-1 illustrates the analogous situation in regard to the onset of channel flows. We define

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"rcv as the delay since the time of rapid crustal thickening, required for the onset of gravitationally driven channel flow. Given that channel flow requires weaker crust than is necessary for the development of hot nappes during indentation, ZCF is normally larger than ZI-INfor the same material properties. This relationship explains LHO-3 behaviour, in which the initial response to indentation is to produce fold nappes (Fig. 1 ld), because the incubation time is greater than "ri-iN but less than "rcr. Later, at 65 million years, a more pronounced channel flow is superimposed on the system (Fig. l i e ) indicating that the incubation time is now greater than "rCF. Model LHO-2 also illustrates the effects of incubation. During Phase 1 deformation, the weaker lower crustal regions are detached and evacuated during their incorporation into the model orogen (Fig. 7b-d). These regions are already sufficiently hot and weak that they do not require incubation in order to form hot nappes. In contrast, the overlying mid-crust initially deforms by shortening and thickening (Fig. 7b). The gravitationally driven channel flow (Fig. 7 c - e ) develops only during the diachronous second phase, and has a ~cv of approximately 30 million years. In summary, comparison of the thermal incubation time with ZHN and ZCF provides a guide to the flow modes that will develop in the models. Diachroneity of crustal tectonic thickening implies that different flow modes can coexist in different regions of the model. For natural orogens we may know the timing of tectonic thickening of the crust but lack information on the thermalrheological evolution. In the absence of good estimates of "rHN and ZcF, indirect evidence of crustal viscosity and temperature can be derived from large-scale topography and magmatism. Plateau development in collisional orogens demonstrates that the crust, and possibly the entire lithosphere, cannot sustain thickness variations against gravitational forces and has flowed to equilibrate the pressure in the crust below the plateau to a lithostatic state. This state indicates the crust is weak and that it is more likely to flow over a lower crustal indentor than to shorten and thicken against it. The development of a plateau is therefore a measure of a thermo-rheologically weak crust, and the delay from crustal thickening to initial plateau development is an approximate estimate of "rHN and a lower bound on ~'CF. Magmatism, particularly involving widespread crustal melts, is an independent indicator that crustal temperature is high and that the crust is therefore probably weak. The delay from orogenic crustal thickening to the onset of magmatism can also be used as an upper bound estimate of ~'HN, and for crustal melts also provides an estimate of zCF.

Infrastructure a n d superstructure All three LHO models illustrate the development of differing styles of deformation at different levels of the crust corresponding to what is termed 'infrastructure and superstructure' in the classical geological literature. The terminology refers to contrasting styles of deformation and metamorphism in the upper, superstructure, and lower, infrastructure, levels of the crust. The superstructure preserves early, low-grade deformation, typically with upright contractional structures, whereas the infrastructure is a ductile, high-grade, migmatitic level with younger, gently dipping structures that overprint early structures (Culshaw et al. 2004; see also Williams et al. 2006). In the LHO models the superstructure develops during crustal shortening and thickening and the infrastructure is superimposed by the flows that develop later in the mid- and lower crust. The main differences among the models are in the cause and timing of infrastructure development. In model LHO-3 (Fig. 11), the infrastructure comprises lower and mid-crustal nappes partly overlying the underthrust indentor and decoupled from the upper crust by a reverse-sense shear zone at the top of the highly strained ductile mid-crust (Fig. 1 ld & e). The upper crustal superstructure remains relatively undeformed after Phase 1 shortening except where exhumed by syntectonic erosion (pro- versus retro-sides of Fig. 1ld & e). In contrast, Phase 1 structures in the mid- and lower-crustal infrastructure are strongly overprinted by Phase 3 flow (Fig. 1 l d & e). From an observational perspective, the three-phase evolution of model LHO-3 leads to what would be recognized geologically as an old, but not reworked, contractional upper crust underlain by and decoupled from mid- and lower crust that records the initial contraction, thermal relaxation, and superimposed deformation activated by Phase 3 collision with the indentor. The mid- and lower crust become weak, with "rHN approximately 20-30 million years, but hot-nappe deformation is not strongly activated until the indentor collides much later. In this example the infrastructure is created by the tectonic indentation process, not by internal gravitational flow, after a long incubation time. In contrast, the infrastructure in models LHO-I and LHO-2 becomes sufficiently weak during thermal relaxation that it deforms and flows under gravitational forces alone (Figs 3 and 7). Under these circumstances the development of superstructure/ infrastructure relationships in LHO-1 is governed by the delay timecale ~'CF for channel flows and does not require external forcing by an indentor or some other tectonic process. In LHO-2 the formation of the lower crustal infrastructure starts

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during the initial crustal shortening and thickening and is subsequently overprinted by the channel flow on the ~'cF timescale.

Conclusions Numerical models have been used to investigate crustal flows in large hot orogens in plane strain at the crustal and upper-mantle scale. The flow styles are divided into three types, homogeneous, heterogeneous, and hot fold nappe modes, and the conditions under which each will operate are assessed in the framework of the orogenic T-M diagram, which plots orogen characteristics in terms of the two principal controls, temperature and magnitude. We draw five main conclusions. (1)

(2)

(3)

Gravitationally driven mid-crustal channel flows, exemplified by the homogeneous and heterogeneous modes (Fig. 20), are most likely to occur in Giant and Supergiant members of the large hot orogen family (Figs 1 & 19), such as the Himalayan-Tibetan system. Gravitationally driven flows require orogenic crust to be weak (low viscosity) and therefore hot under most circumstances. These flows are facilitated by gravitational forces that result from the large difference in potential energy between the tectonically thickened interior and normal thickness exterior crust in large hot orogens. These flows may have been more prevalent in Archean orogens, if they were indeed hotter than equivalent-sized contemporary orogens. Tectonically forced flow modes, exemplified by the hot fold nappes mode and the tectonic component of the heterogeneous flow mode (Fig. 20), may occur in Giant and Supergiant orogens (Figs 1 & 19). More importantly, they can occur in large hot orogens that are not hot and/or large enough to undergo gravitationally driven flows. In particular, the hot fold nappes mode is predicted for accretionary and collisional orogens where the orogen experiences late-stage collision/indentation by strong crust, for example older refractory crust such as a cratonic nucleus, or cold oceanic crust. Both crustal-scale (CS) models with kinematic mantle subduction basal boundary conditions and upper-mantle-scale (UMS) models can develop homogeneous crustal channel flows. To a first approximation the general characteristics of these flows are insensitive to the effects of advancing or retreating subduction of the underlying mantle lithosphere. However, a moving subduction zone will change its position relative to the crustal

Fig. 20. Summary diagram of the three crustal flow modes investigated in this work, together with their characteristics and the requirements for each of them to operate.

(4)

channel and its final position may vary from a location beneath the edge of the plateau (Fig. 17) to one beneath the centre of the plateau (Fig. 18). The temperature-magnitude (T-M) diagram, which we have introduced here, provides a framework for the classification of orogens, ranging from small cold to large hot, and for

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(5)

their d e v e l o p m e n t with respect to a main sequence (Fig. 1). W e suggest that natural orogens can be analysed, at least conceptually, by their evolution in T - M space and that this approach offers a w a y to predict w h e n different types of flow m o d e s m a y occur. These flow modes have been inferred to have developed in some ancient orogens, including the Mesoproterozoic western Grenville orogen (Jamieson et al. 2004a), and the Palaeozoic Inner Piedmont o f the southern Appalachians (Hatcher & Merschat 2006). In North America, which has grown outward by successive collisions against and accretion to the cratonic core, which could have acted as an indentor, we anticipate that the tectonically driven styles will be recognized in m a n y large hot orogens, including the Trans-Hudson and southern Canadian Cordillera orogens, and parts o f the Archean Slave and Superior cratons.

This research was funded by NSERC Discovery Grants to Beaumont and Jamieson. Beaumont also acknowledges support from the Canadian Institute for Advanced Research, the Canadian Research Chairs programme, and from an IBM Shared University Research grant. The models used software developed by P. Fullsack, Dalhousie University. The work benefited from discussions with colleagues including N. Culshaw, D.Grujic and especially the late D. Nelson. We thank the reviewers G. Houseman and J. Platt for thorough reviews and R. Law for editorial advice. We are grateful to the conference organizers for providing a stimulating environment in which to respond to M. Harrison, and to D. Nelson for originally encouraging us to tackle this problem.

Appendix: Design of crustal- and upper-mantle-scale models Philosophy of numerical approach to problem solution and model parameterization What are the criteria for the development of geodynamical models and how complex should they be? In the design of numerical models there is a trade-off between those that are overly simplified/specified and therefore fail to demonstrate important types of behaviour because relevant physical processes are omitted/suppressed, and those that are overly complex, displaying characteristics that are difficult to interpret owing to the large number of possible interactions. Our motivation is to illuminate the most basic physics behind orogenic evolution. We therefore choose a numerical methodology that is robust and includes the ability to solve the underlying coupled mechanical and thermal problems that operate at orogen scales. We specifically avoid 'simulation', in which the models may be overconstrained with the intent of reproducing or 'mimicking' a

particular natural setting in detail. We prefer to view our models as numerical experiments designed to investigate the types of processes that occur within models of collisional orogenesis with boundary conditions that are deliberately simplified by comparison with nature. For example, the velocity boundary conditions for the crustal-scale (CS) models HT1 and H T l l l described in Jamieson et al. (2006) and their upper-mantle-scale (UMS) equivalents are purposely restricted to be uniform approximations of the natural Himalayan-Tibetan system. Our thesis is that this simplified approach will reveal the underlying firstorder processes.

Advantages and limitations of the CS and UMS model designs We choose the simplest model design that is compatible with the first-order processes and features of natural orogenic systems - in this case, large, hot, collisional orogens. The CS and UMS model designs described in this paper have both limitations and advantages. Limitations include: (1) the 2D plane-strain restriction (no flow of material out of the plane of the model); (2) the crustal scale of the CS model (no mantle dynamics); and (3) the choice of basal kinematic boundary conditions. The advantages include: (1) the fully dynamical solution of the flow calculation within the CS crustal section subject to the boundary and surface process conditions, and the dynamical solution at the upper-mantle scale in UMS models; (2) the ability to include pressure-dependent plastic (Drucker-Prager) rheologies, corresponding to Coulomb failure and Byerlee's law, and a first-order approximation of the effects of pore-fluid pressures (CS and UMS) and a parametric model for strain-softening in UMS; (3) the inclusion of thermally activated power-law viscous creep; (4) the coupled thermal-mechanical nature of the calculation; and (5) the Arbitrary Lagrangian Eulerian (ALE) formulation of the finite-element problem, which both allows for sufficiently accurate calculations at medium scales within the problem domain, and includes the calculation of the evolving shape of the model domain such that orogen geometry, topography, plateau growth, surface processes and the gravitational feedback effects of changing geometry, and large deformation, are easily and naturally incorporated in the calculation. The basic design of the ALE numerical model has been described elsewhere (Fullsack 1995; Jamieson et al. 2002; Beaumont et al. 2004) and was summarized in the Numerical Calculation section above. The same CS numerical model is used in the calculations described by Jamieson et al. (2006).

Model complexity and selection and tuning of model properties Even with the simplifications described above, the models may appear to include a large number of parameters whose

CRUSTAL FLOW MODES values are poorly known. These can, however, be grouped into only four property sets: (l) the mechanical properties required to specify a three-layer crust (CS) and lithosphere and mantle (UMS); (2) the associated thermal properties; (3) the velocity boundary conditions; and (4) the properties of the surface processes model. All of these play important roles in natural systems and cannot be neglected in the models. Although we show only a selection of the results, they are based on extensive sensitivity analyses in which a reference model is established and then tested for its sensitivity to variations in one or, at most, two of the properties at a time, e.g. time variations in the intensity of the surface processesf(t), or the spatial variation of kinematic boundary conditions. There are three important steps in the model design: (1) the selection of a reference model; (2) the choice of parameter variations to be used in the sensitivity analysis; and (3) the assessment of the results for robust outcomes. The approach is reductionist in that a direct cause-andeffect relationship between parameter variation and model behaviour is sought. Although some of these relationships can be interpreted to be robust, the behaviour is commonly a system response involving the dynamics of one or more feedback loops that cannot be demonstrated to be uniquely related to a single input parameter. Our experience with sensitivity analyses yields some confidence in attributing cause-and-effect relationships. It also indicates when the model outcomes become very sensitive to small variations in several input variables. In such cases, it is important to establish the range of expected variability in the model context. Equivalent natural systems can also be expected to show a range of behaviours owing to their inherent natural variability. However, it will likely be impossible to attribute a specific cause-and-effect relationship for most specific natural examples, because we commonly do not know the system properties and their associated variations accurately. Below we describe tests for some boundary conditions, the rationale for some specific parameter choices, and explain how the HT model series was developed from a simpler reference model.

Testing the basal boundary conditions in CS models In CS models the basal velocity boundary conditions are specified kinematically to correspond to assumed behaviours of the mantle lithosphere, for example subduction, advancing subduction, or pure shear thickening. The UMS model experiments provide an opportunity to test these assumptions by removing the specified velocities at the base of the crust and, instead, model the dynamics of the interaction between the lithosphere and underlying mantle. The observed model behaviours range from advancing double subduction, through subduction, to subduction-zone retreat, and include shortening and

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thickening of the mantle lithosphere and various forms of convective instabilities of the mantle lithosphere (e.g. dripping, slab break-off; Pysklywec et al. 2000). Mantle subduction is the preferred mode when the early stages of deformation correspond to underthrusting of one mantle lithosphere beneath the other. In the models, subduction is facilitated by a weak shear zone between the two converging lithospheres; in nature, this might be inherited from a phase of oceanic subduction prior to continent-continent collision. Distributed lithospheric contraction and thickening occurs in the absence of significant zones of weakness that could act to break the symmetry of pure shear thickening. It cannot be demonstrated that mantle subduction necessarily accompanies continent-continent collision. However, as described in this paper, results from UMS models which predict dynamic mantle subduction are compatible with those from Himalayan-style CS models with kinematic subduction. Many UMS models, with a range of properties, exhibit subduction with combinations of subduction zone advance and retreat that are controlled by the density contrast between the mantle lithosphere and sublithospheric mantle. When the density contrast is large there is also a tendency for repeated slab break-off events. Therefore, the possibility of punctuated subduction of mantle lithosphere must be considered, possibly associated with reversals in subduction polarity (Pysklywec 2001).

Scaling of laboratory power-law creep flow laws We choose to base the flow laws in the models on a reference set of well constrained laboratory experimental results: wet quartz (WQ) (Gleason & Tullis 1995; meltabsent Black Hills quartzite), dry diabase (DMD) (Mackwell et al. 1998; dry Maryland diabase), and wet olivine (WO) (Chopra & Patterson 1984; wet Aheim dunite; Karato & Wu 1993). Laboratory-derived flow laws are subject to significant uncertainties associated with the measurements on individual samples, the variability of measured results among samples of similar rock types, the range of deformation mechanisms, the effects of water fugacity, and the known and unknown errors in extrapolating the laboratory results to natural conditions. We have therefore chosen to limit the complexity and to base our model rheologies on a few reliable data-sets in order to minimize the number of sources of error while allowing some variation in the model viscous flow properties. Flow laws for rocks that are stronger/weaker than the base set are constructed by linearly scaling up/down the values of B* (Equation 3). This approach is used to approximate other material rheologies. The scaled viscosities can either be interpreted in terms of the effects of composition or the consequences of water-saturated versus water-poor (wet versus dry) conditions. This is

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valid if the exponent of the water fugacity term is close to unity, and therefore the effect of water scales linearly with A in the flow law (e.g. Hirth et al. 2001). Alternatively they can be interpreted as synthetic model rheologies. Given that relative ductile flow of different materials in the models is mainly a consequence of their viscosity contrast, the simple scaling guarantees that the viscosity contrast is always given by the scaling factor under the same ambient conditions. This approach simplifies the interpretation of the model results and is the principal reason for choosing it: instead of having results in which all of the parameters in the power-law creep flow law vary (Equation 2), only the effective viscosity varies as B* is scaled. We believe that this scaling is an appropriate way to test the sensitivity of the models to the effect of wet versus dry conditions or to a moderate change in composition. For example, B*(DRY) is in the range 1050 x B*(WET), and B*(WQ x 5)approximates conditions when flow is influenced by a mineral such as feldspar that has a higher effective viscosity than wet quartz for the ambient conditions. We choose dry Maryland diabase (B*(DMD)) to represent the strongest lower crustal rheology, knowing that a comparison demonstrates that B* (DMD/IO) corresponds closely to the power-law flow properties of intermediate composition granulite (Pikwitonei granulite: Wilks & Carter 1990; Mackwell et al. 1998). Given uncertainties in the composition and other properties of the lower crust, we argue that a reasonable approximation of power-law creep of the lower crust can be based on proxy materials ranging from B*(DMD) (the strongest end-member) to B*(DMD/20) (weak lower crust). We in no way imply that the lower crust is diabase. Similarly, in the UMS models, ductile flow of the mantle is based on the power-law rheology of olivinecontrolled rocks; we use B*(WO) (wet Aheim dunite: Chopra & Patterson 1984) as the reference rheoiogy, knowing that this flow law corresponds closely to that of wet olivine (Karato & Wu 1993). To a first approximation, dry olivine has an effective viscosity that is as much as 50x that of wet olivine for mantle lithosphere conditions. We therefore use B*(WO) for sublithospheric mantle, assumed to be water-saturated, and B*(WO x 10) for continental mantle lithosphere that is considered to be relatively water-poor. The effect of the activation volume is not included in the calculation of the power-law creep flow laws. In the lithosphere, pressure is sufficiently low that the activation volume effect on viscosity is not significant. In the upper mantle the effect could be large, but prediction of the effective viscosity for wet olivine is complicated by the pressure and temperature dependence of water fugacity and whether the system behaviour is open or closed (Karato & Jung 2003). For the purposes of the demonstration models we omit both of these effects, but limit the sublithospheric minimum viscosity to 1019 Pa s, which is somewhat larger than the predicted watersaturated values (Karato & Jung 2003).

Design of Himalayan-Tibetan

(HT) models

The HT series models were developed from a large hot orogen CS reference model similar to Model 1 of Beaumont et al. (2004) but with no melt-weakening or erosion. The reference model has Vp = 2.0 cm a -1 and Vs = VR=0. The undeformed crust has ~be~-= 15 ~ throughout, and comprises a 25 km thick upper/midcrustal layer with B*(WQ) and a 10 km thick lower crust with B*(DMD). The lower crust is not subducted, there is no melt weakening, and surface processes are not included. Thermal properties are those used by Jamieson et al. (2002, 2004b). The model contains two layers with contrasting heat production, A1 = 2.0 puW m -3 (020km) and A2 = 0.75 txW m -3 (20-35km). These values were chosen to represent continental margin crust (Jamieson et al. 2002); upper-crustal heat production, in particular, lies within the range reported from GHS lithologies (e.g. Huerta et al. 1998, and references therein). As described in the Numerical Calculation section above, values of other thermal parameters (K, K, Cp, p) are identical in both layers and lie in the mid-range of those normally attributed to continental crust (e.g. England & Thompson 1984). Results from the reference model show significant departures from the first-order properties of the Himalayan-Tibetan orogen. A number of physically justifiable modifications were therefore made which led to the HT series of models, from which representative model HT1 was subjected to detailed analysis (Beaumont et al. 2004; Jamieson et al. 2004b, 2006). The five essential modifications incorporated into the HT series models in order to produce model results compatible with observations are listed below. Model thermal properties were not adjusted. (1) Change velocity boundary conditions. For consistency with estimates of average India-Asia convergence velocity, the HT series models have Vp = 5 cm a 1. The models are viewed in the fixed Asia reference frame, VR = 0, with advancing subduction, Vs = 2.5 cm a -~. Royden et al. (1997) and Beaumont et al. (2004) demonstrated that advancing subduction is required to reproduce the general planform geometry of the Himalayan-Tibetan system and the surface position of the Indus-Tsangpo suture. For reasons noted above, velocity boundary conditions remain constant during each of the CS model experiments. (2) Subduct lower pro-crust. Accumulation of lower pro-crust in the model orogen produces a large lower crustal antiform, inconsistent with data from the Himalayan-Tibetan orogen (e.g. Model 1 of Beaumont et aI. 2004). The lower crustal layer on the pro-side of the model system (corresponding to India) is therefore subducted along with the pro-mantle lithosphere. This is consistent with mechanical coupling between strong lowermost crust and upper mantle in mature continental crust, and with lithosphere-scale interpretations of seismic data from the orogen (Owens & Zandt 1997).

CRUSTAL FLOW MODES Because the lower pro-crust is detached and subducted at the S-point, it behaves like the mantle directly beneath it and is not deformed during model evolution. As the overlying crust thickens and heats up, it becomes mechanically decoupled and detached from the lower crust, which is overridden as the orogen propagates towards its foreland. (3) Include melt-weakening. As shown by Beaumont et al. (2004, Model 3 versus Model 1), models without melt-weakening produce, at best, inefficient channel flows restricted to the region beneath the central model plateau. Including a parameterized viscosity reduction over the temperature interval associated with melting (Beaumont et al. 2001, 2004; Jamieson et al. 2002) produces efficient channel flows extending to the plateau edge. This is consistent with seismic evidence that some melt is present under the present-day Tibetan Plateau (Nelson et al. 1996; Klemperer 2006) and with observations that GHS gneisses (exhumed equivalents of postulated channel) are typically migmatitic. In HT models, melt-weakening is restricted to the middle and upper crustal (quartzo-feldspathic) layers and does not affect the lowermost (granulitic) crust. (4) Include surface denudation. In the absence of erosion, the channel flow zone 'tunnels' into the surrounding crust at a rate controlled by its thickness and the temperature field (Royden 1996; Beaumont et al. 2004; Medvedev & Beaumont 2006). In order to exhume the channel it is necessary to erode the plateau flank. In the HT models, surface denudation is controlled by the interaction of surface slope, a spatial function (g(x)), and a time function (f(t)). Local surface slope is calculated within the model. To a first approximation g(x) is a measure of the spatial variation of aridity (0 = dry, 1 = wet) across the model, and f(t) combines the effects of long-term climate variations, the bedrock incision rate constant, and a parameter that scales the model surface slopes, which are determined on a 10 km spatial resolution, to include higher riverbed slopes at smaller scales. A more detailed denudation model is not justified because the model is cross-sectional, and therefore cannot represent planform drainage patterns, and the scaling effect in f(t) for local slopes at less than 10 km spatial resolution is not known accurately. All HT models are run for an initial set-up phase ( 0 24 million years; 5 4 - 3 0 Ma) without surface denudation. This is not a significant factor in the later stages of model evolution (the focus of our work to date), and is designed to achieve a model state with an embryonic plateau and mid-crustal channel flow as a precursor to testing model sensitivity to denudation. The results axe similar with moderate denudation during the set-up phase but the times to develop the plateau and channel flow are somewhat longer. In model HT1, erosion rate is high from 24 to 39 million years (30-15 Ma), which initiates efficient channel extrusion, and then declines gradually from 39 to 54 million years (15-0 Ma) towards present-day values. Similar model results are obtained using somewhat different denudation functions (e.g. Model 3 of Beaumont

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et al. 2001). However, successful models require a period of rapid erosion (f(t) large) after 24 million years (30 Ma) in order to initiate channel exhumation, and a decline from the maximum rate (f(t) decreasing) in the last 15-20 million years of model evolution in order to produce model ages for peak metamorphism and cooling that lie within the observed range. As noted by Jamieson et al. (2004), 'GHS' cooling ages predicted by HT1 are too young, suggesting that recent erosion rates should be even lower. With the additional provenance and detrital mineral data that have recently become available (e.g. DeCelles et al. 2004; Amidon et al. 2005; Najman et al. 2005; Najman 2006), different denudation functions might be chosen for a future series of models. The original HT1 design is retained by Jamieson et al. (2006) in order to complete the analysis of that particular model. (5) Include three crustal layers. The modifications noted above lead to the development of a Himalayan-scale model orogen with extrusion of a mid-crustal channel on timescales of 50-55 million years. However, a crustal structure comprising three laterally continuous layers with contrasting mechanical properties produces significant improvements in the model. In particular, a weak upper crustal layer that is capable of detaching from underlying middle crust allows the formation of an asymmetric overthrust structure at the orogenic front and domes in the region between the plateau flank and the suture (Beaumont et aL 2004; Jamieson et aL 2006). The theology of the uppermost layer (0-10 km) is given by B*(WQ) with ~bey= 5 ~ representing sedimentary rocks of the upper crust with high pore fluid pressures. The middle crustal layer (10-25 km) uses B*(WQ x 5) with ~bey= 15 ~ representing quartzo-feldspathic granitic and/or metasedimentary rocks. As described above, the upper and middle crustal layers are subject to melt-weakening where T _> 700~ The rheology of the lower crust (25-35 km) is given by B*(DMD), with ~be~.= 15 ~ representing lower crustal granulite. Similar results are obtained with B*(DMD/5). The lower crustal layer is not subject to melt-weakening. Some HT series models use variations on this simple three-layer structure, which are described where specific model results are presented.

Essential model requirements for channel flow We use the terms channel flow and extrusion to describe the general process of orogen-scale, confined, pressure-driven flow (analogous to pipe flow; Turcotte & Schubert 1982, p. 237) and the ejection of the channel material toward the surface near the end of the flow zone. In order to generate channel flow in the model, the only requirements are reduced viscosity, ~ey -< 1019 Pa s, and a pressure differential sufficient to drive a flow with that viscosity. In the HT models, both the pressure differential and the reduced viscosity result from crustal thickening. The pressure differential comes from the potential energy difference resulting from the contrast in crustal thickness and elevation

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between the plateau and the foreland, and the viscosity reduction is associated with high temperatures generated by heat production in thickened crust. Beneath the plateau, material flux through the channel is related to its thickness and viscosity (scales with h3/~; e.g. Royden 1996) and the rate of channel propagation is limited by the rate at which adjacent material becomes hot and weak enough to be incorporated into the advancing channel flow (Beaumont et al. 2004; Medvedev & Beaumont 2006). The active or previously active (fossil) channel is exhumed by focused surface denudation. Extrusion between coeval thrust- and normal-sense shear zones occurs where material in an active channel (T >_ 700~ is pumped or forced towards the surface; by analogy with pipe flows, the surface represents the open end of the pipe. Since the temperature at the model surface is 0~ channel material cools during extrusion at a rate determined largely by the rate of denudation. In the models, and probably also in nature, the geometry of the channel is significantly modified during extrusion. In the model, deformation superimposed on the channel material at this stage generally involves flattening and thinning. By implication, structures in natural exhumed channels should record features formed during active channel flow as well as features superimposed during extrusion, exhumation and cooling. It might therefore be difficult to determine unambiguously whether or not channel flow has occurred from structural analysis of specific exhumed sections. Model-data

comparisons

What are the most effective tests of the models? The feasibility of any numerical model for orogenesis must be tested against data from real orogenic systems. Conversely, the feasibility of conceptual models based on geological or geophysical data, and of kinematic models based on predefined geometries, should be tested against the physics of the system as a whole. Are the assumptions physically realistic? In either case, the tests should be designed to reflect the first-order properties of the model system on the appropriate scale. If the models fail the first-order tests, second-order features are irrelevant. If the models pass the first-order tests, it must be determined whether second-order model predictions are robust, and therefore testable, and whether the second-order features themselves are well enough characterized to constitute valid tests of a specific model. Given that orogenesis is a response to the behaviour of the lithosphere during convergence, the present models are designed on the scale of the crust and upper mantle. This imposes numerical limitations on model resolution and there is a corresponding limit to the scale at which model predictions can be reliably compared with data from specific orogenic transects. A further limitation on model-data comparisons is the two-dimensional, planestrain, nature of the numerical models presented here. The Himalayan-Tibetan system displays remarkable

along-strike continuity (e.g. Hodges 2000), which allows reasonable first-order model-data comparisons for the central part of the orogen. However, where three-dimensional effects are important, e.g. in the vicinity of the Himalayan syntaxes, specific model-data comparisons become tenuous. In comparing our model results with data from natural orogens, we first assess consistency with crustal- or lithospheric-scale features before making comparisons with specific seismic, structural, metamorphic, slxatigraphic or geochronological data-sets. In compiling geological or geophysical data, we look for regional-scale consistency in order to distinguish general (first-order) properties of the system from those controlled by local features. Similarly, matching the details of a particular type of data (e.g. a specific P-T-t path) is less important than consistency with combinations of data (e.g. P-T-t path style combined with peak grade profiles and geochronology). The first-order test of the channel flow model is the existence of mid-crustal channels with large-scale flows characterized by velocities on the order of 1 cm a -1. This has not yet been detected directly. In the Himalayan-Tibetan system, indirect evidence for channel flow includes a variety of geophysical data from the Tibetan Plateau, as summarized by Klemperer (2006), the magnetotelluric evidence (e.g. Unsworth et al. 2005) and a range of geological data summarized by Jamieson et al. (2004b). While indirect evidence may not constitute a diagnostic test, the ability of the homogeneous channel flow model to account for a wide array of disparate features of the orogen suggests that the simple model captures many essential elements of the behaviour of the system. We conclude that channel flow models in general provide a reasonable first-order explanation for the thermaltectonic and lithological evolution of the Himalaya and southern Tibet. In exposed mid-crustal levels of ancient orogens, a number of geological observations could constitute tests for the former existence of channel flows. These include: (1) the presence of coeval normal- and thrust-sense shear zones bounding a regionally extensive zone of migmatite or some other material inferred to have had low viscosity (relative to underlying and overlying rocks) at the time the shear zones were active; (2) a transition from an 'inverted' metamorphic sequence across the basal thrust-sense shear zone into a 'normal' metamorphic sequence across the upper normal-sense shear zone; (3) evidence that ductile flow in the low viscosity zone post-dated compressional deformation in overlying crust by c. 20-25 million years (time needed to initiate melt-weakening in thickened crust); (4) discontinuity between upper and lower crustal structures across the ductile flow zone; (5) evidence for substantial lateral transport of low-viscosity material along structures that were shallow-dipping at the time that they formed. Beaumont et al. (2001, 2004, this study) and Jamieson et al. (2004b, 2006) have demonstrated that both CS and UMS models are sensitive to small variations in

CRUSTAL FLOW MODES parameters such as crustal strength, denudation history, and upper mantle density. Within the range of natural variability of these parameters, the model system can respond in different ways to produce a variety of features observed in different places and/or times in the evolution of the orogen. The resulting variability does not extend to its first-order features, i.e. the generation and exhumation of mid-crustal channel flows, but can produce significant differences in the surface expression of the underlying processes. Under these circumstances a model could potentially be 'tuned' to achieve a desired effect, for example to explain the details of a specific transect. As discussed above, model tuning to fit second-order features provides little or no insight into processes, and the resulting match does not constitute a valid test of the model. However, far from being a weakness of the HT model series, its sensitivity to variations in parameters that are demonstrably variable in nature should be regarded as one of its strengths. This is in itself an important test of the model. Models that fail to predict natural variability are inadequate. It follows that the expectation that one specific model should explain all the features of an orogen is wrong, and conversely there is no unique model against which all observations should be compared.

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softening. Journal of Geophysical Research, 108 (B 10), 2496. DOI: 10.1029/2002JB002026. JAMIESON, R. A., BEAUMONT, C., FULLSACK, P. & LEE, B. 1998. Barrovian regional metamorphism: Where's the heat? In: TRELOAR, P. & O'BRIEN, P. (eds) What Controls Metamorphism and Metamorphic Reactions? Geological Society, London, Special Publications, 138, 23-51. 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., CULSHAW, N., BEAUMONT, C., NGUYEN, M. H. & SLAGSTAD, T. 2004a. 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 Abstracts Volume. Geological Society, London. JAMIESON, R. A., BEAUMONT, C., MEDVEDEV, S. NGUYEN, M. H. 2004b. 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. & GOD1N, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 165-182. KARATO, S. & JUNG, H. 2003. Effects of pressure on high-temperature dislocation creep in olivine. Philosophical Magazine, 83, 404-414. KARATO, S. & Wu, P. 1993. Rheology of the upper mantle: a synthesis. Science, 620, 771-778. 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 Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 39-70. LEHNER, F. K. 2000. Approximate theory of substratum creep and associated overburden deformation in salt basins and deltas. In: LEHNER, F. K. 8r URAI, J. L. (eds) Aspects of Tectonic Faulting. Springer-Verlag, Berlin, 21-47. MACKWELL, S. J., ZIMMERMAN,M. E. & KOHLSTEDT, D. L. 1998. High-temperature deformation of dry diabase with application to tectonics on Venus. Journal of Geophysical Research, 103, 975-984. MEDVEDEV, S. & BEAUMONT,C. 2006. Growth of continental plateaus by channel injection: models designed to address constraints and thermo-mechanical consistency In: LAW, R. D., SEARLE, M. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision

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deformation and lower crustal flow in eastern Tibet. Science, 276, 788-790. RUSSELL, H. N. 1914. Relations between the spectra and other characteristics of the stars. Popular Astronomy, 22, 275- 294, 331 - 351. [Excerpt with commentary appears in LANG, K. R. GIN~ERICH, O. (eds) 1979. A Source Book in Astronomy & Astrophysics, 1900-1975, Harvard University Press, Cambridge, Mass, 212-220; also in SHAPLEY, H. (ed.) 1960. Source Book in Astronomy, 1900-1950, Harvard University Press, Cambridge, Mass., 253-262. SHEN, F., ROYDEN, L. H. & BURCHFIEL, B. C. 2001. Large-scale crustal deformation of the Tibetan Plateau. Journal of Geophysical Research, 106, 6793-6816. TAt, W. C. & O'CONNELL, R. J. 1992. Ablative subduction: A two-sided alternative to the conventional subduction model. Journal of Geophysical Research, 97, 8877-8904. TURCOTTE, D. L. & SCHUBERT, G. 1982. Geodynamics: Applications of Continuum Physics to Geological Problems. John Wiley & Sons, Chichester. UNSWORTH, M. J., JONES, A. G., WEI, W., MARQUIS, G., GOKARN, S. G., SPRATT, J. E. & the INDEPTH-MT team 2005. Crustal theology of the Himalaya and Southern Tibet inferred from magnetotelluric data. Nature, 438, 78-81. VANDERHAEGHE, O., MEDVEDEV, S., FULLSACK, P., BEAUMONT, C. & JAMIESON, R. A. 2003. Dynamic evolution of orogenic wedges and continental plateaus: insights from thermal-mechanical modeling of convergent orogens. Geophysical Journal International, 153, 27-51. WASCHBUSCH, P., BATT, G. & BEAUMONT, C. 1998. Subduction zone retreat and recent tectonics of the south island of New Zealand. Tectonics, 17, 267-284. 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. WILKS, K. R. & CARTER, N. L. 1990. Rheology of some continental crustal rocks. Tectonophysics, 182, 57-77. WILLETT, S. D. 1999. Orogeny and orography: The effects of erosion on the structure of mountain belts. Journal of Geophysical Research, 104, 28,957-28,981. WILLETT, S. D. & BEAUMONT, C. 1994. Subduction of Asian lithospheric mantle beneath Tibet inferred from models of continental collision. Nature, 369, 642-645. WILLETT, S., BEAUMONT, C. & FULLSACK,P. 1993. A mechanical model for the tectonics of doublyvergent compressional orogens. Geology, 21, 371-374. WILLIAMS, P. F., JIANG, 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. P. & GODIN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 221-235.

Growth of continental plateaus by channel injection: models designed to address constraints and thermomechanical consistency S. M E D V E D E V t'2'3 & C. B E A U M O N T

1

aDepartment of Oceanography, Dalhousie University, Halifax, NS, B3H4J1 Canada 2Fachrichtung Geologie, Freie Universitiit Berlin, Malteserstrasse 74-100, Berlin, 12249 Germany 3Present address." Physics of Geological Processes, University of Oslo, PO Box 1048 Blindern, 0316 Oslo, Norway (e-mail: [email protected]) Abstract: Weak, possibly partially molten, middle crust may exist and deform by channel flow beneath continental plateaus, thereby significantly influencing their dynamics. The role of channel flows in the transition zone between the plateau and the foreland is, however, unclear. We develop successively more complete approximate models for the channel injection (CI) mode in which differential pressure pumps channel material from beneath the plateau into the transitional crust, which thickens it and widens the plateau. The motivation is to improve our understanding of the controls on the growth of continental plateaus and the interactions in the transition zone, and to gain more insight into the results of more complex numerical models. In model CI- 1, a channel with constant viscosity and thickness exists in the transitional crust and the pumped material accretes/freezes above and below the channel. Although results compare favourably with the geometry of some natural examples, this model is incomplete because the connection between the transition zone and the plateau is not considered. Model CI-2 includes a decrease in channel viscosity when the channel depth exceeds a critical value, D*, a proxy for onset of melt weakening or low viscosities at high temperatures. The model completes the connection to the plateau, but relies on the arbitrary choice of D*. Model CI-3 is more physically based, and considers the channel viscosity and thickness to depend on temperature, calculated by an associated thermal model that includes radioactive self-heating, and advection of heat by channel material. This model demonstrates self-consistent plateau widening if the channel viscosity decreases at the critical temperature, T*. Acceptable comparisons with the topography of Tibet are achieved with transition zone viscosities that decrease from 10~9-1022pas to subplateau values of 1018-10 w Pa s, with T* of 700-750~ Additional analyses and tests are used to determine the range of parameter values for which CI models are both internally consistent and compatible with observations. Additional modes of deformation in the transition zone, viscous thickening (VT) and plastic translation (PT), may also be important.

High temperatures and partial melting may lead to low-viscosity channel flows in the mid- to lower crust beneath continental plateaus (Fig. la). Following earlier work demonstrating the feasibility of such channel flows (e.g. Bird 1991; Kaufman & Royden 1994), several recent studies indicate that these flows lead to decoupling of deformation with focused flow in the channel and overlying crust that remains relatively stable (e.g. Royden 1996; Beaumont et al. 2001, 2004; Shen et al. 2001; Vanderhaeghe et al. 2003). These models also predict that plateaus grow outward by tectonic thickening and accretion of adjacent crust in the transition zone between the plateau and the surrounding foreland (Fig. lc) and that channel flows may subsequently convert the transition zone to become part of the plateau. In contrast, Clark & Royden (2000) proposed that crust in the transition zone may thicken by direct injection of channel material (Fig. lb) and showed

this style to be compatible with the crustal thickening of the eastern margin of the Tibetan plateau. In this proposed channel injection (CI) mode no crustal shortening in the transition zone is required, as suggested by the limited shortening observed from the eastern margin of the Tibetan plateau. A primary motivation for further development of channel injection models, as described here, is to improve our understanding of the first-order controls on the growth of continental plateaus and the interactions that occur within the transition zone. The second motivation is to gain more insight into predictions of numerical models which produce realisticlooking results but are difficult to interpret in terms of the basic physical controls (Beaumont et al. 2001, 2004; Jamieson et al. 2004). We investigate some of the basic factors controlling channel flows, and ask whether they contribute to the outward growth of plateaus as in the channel injection mode.

From: LAW, R. D., SEARLE,M. P. & GODIN,L. (eds) ChannelFlow, Ductile Extrusionand Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 147-164. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

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Fig. 1. Diagram showing possible models for the growth of a continental plateau and its relationship to the adjacent transition zone. (a) There is strong evidence for weak mid-/lower crust below plateaus and channel flow may also occur at depth. Less is known about the relationship between the plateau and the transition zone. Three processes may occur in this zone adjacent to a neutral growing plateau: (b) channel injection (CI); (c) crustal thickening by accretion, in this case considered to be viscous thickening (VT); (d) outward flow of unstable brittle upper and mid-crust, termed plastic translation (PT). Darker grey areas represent zones of high strain rate.

The concepts that form the basis for this work (Fig. 1) rely on the existence of weak crust beneath plateaus. The existence of weak subplateau crust is also consistent with the 'flat' plateau topography in that weak mid- to lower crust implies that the viscosity here is too low to create sufficient deviatoric stress to maintain long-wavelength components of the topography, and therefore plateaus flatten like pancake batter. In contrast, steep topographic gradients in the plateau-flanking transition zones (Fig. 1a) imply stronger, higher effective viscosity crust, and/ or a change in boundary conditions that allows these zones to laterally support (like bookends) the plateau against its tendency to flow outward. In the limit, very weak subplateau crust will undergo outward channel flow driven by the pressure gradient between the plateau and the foreland (Fig. l a), but can this limited differential pressure inject channel material into the transition zone? Possible mechanisms for crustal thickening in the transition zone can be modelled as channel injection (Fig. lb; Clark & Royden 2000), viscous tectonic thickening (Fig. lc; Royden 1996; Shen et al. 2001), and plastic translation (Fig. ld; Lehner 2000; Gemmer et al. 2004), in which the upper brittle, frictional-plastic crust decouples and rides outward above the ductile substratum. Channel injection can thicken the transitional crust without surface shortening, but viscous thickening and

plastic translation require tectonic contraction, and in the last instance extension as well. Two end-member settings also exist for plateau growth: the 'actively convergent margins' of the plateau for which contraction is imposed at the lithospheric scale by large-scale plate motions (e.g. the north and south margins of the Tibetan plateau), and the plate-tectonically 'neutral margins', which are less affected by plate-scale process but may be tectonically contractional owing to the gravitationally driven outward flows of the plateau material (e.g. the eastern margin of Tibetan Plateau). Clearly, channel injection can, in principle, be the sole mechanism that thickens the crust in the transition zone of neutral margins, but there must be a component of tectonic contraction of the crust in actively convergent margins. This distinction separates plateau transition zones into two classes: the actively convergent type modelled in Beaumont et al. (2001, 2004) and Vanderhaeghe et al. (2003); and the neutral type modelled by Clark & Royden (2000). The threedimensional models (Royden 1996; Shen et al. 2001) can address both types of boundary. The two fundamentally different types of boundary raises a 'chicken-and-egg' issue in regard to the role of channel injection in crustal thickening and the growth of plateaus. Can injection of channel material occur directly into normal continental

GROWTH OF PLATEAUS BY CHANNEL INJECTION crust, as modelled by Clark & Royden (2000), or must there be a precursory phase of tectonic thickening followed by thermal incubation to develop the hot, weak conditions under which channel injection will operate efficiently? In this paper, successively more complete approximations to the evolution of a continental plateau and the adjacent transition zone by the channel injection mechanism within crust above a stable isostatically adjusting mantle lithosphere are examined. The first two models, CI-1 and CI-2, largely ignore the thermal aspects of the problem. In CI-3 an associated thermal model is introduced which allows both the thickness and the temperature/viscosity of the channel to be predicted. The last part of the paper addresses the question of mechanical and thermal consistency by restricting the acceptable range of model parameter values to those thermal conditions for which the mechanics of channel injection will both operate efficiently and give results compatible with observations. This analysis also includes a comparison of the CI model with simplified models of viscous thickening (Fig. lc) and plastic translation (Fig. ld) of the crust.

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The possible range of parameter values is large, therefore we restrict the models to those that produce similar geometries to those of Beaumont e t al. (2004) for the transition zone (plateau high approximately 5 km, transition zone 100-150 km) at a similar time during their evolution (c. 20 million years). The ranges of consistent rheological and boundary conditions are then compared among models and with natural examples.

Channel injection theory The simple shear thin sheet approximation (Medvedev & Podladchikov 1999; Lehner 2000) and mass balance (Equation 1) describe the flow in a linearly viscous, locally isostatically balanced, crustal channel at + ~ = 0

Oh2

(1)

aq

at ~-Vx = ~ The geometry and boundary conditions (Fig. 2) are as follows: $2+ is the upper boundary of the

Fig. 2. General channel injection (CI) model. Three crustal zones are considered: reference undeformed foreland crust; the transition zone (where crust thickens by accretion of injected material from/within the channel bounded by S+); the plateau (characterized by nearly flat topography, $3 = S3p = const, and constant thickness crust, Hc = tt~p). Crust has uniform density p. The boundary condition is the flux of material from the plateau into the channel, qbMaterial flow in the transition-zone channel comprises a Poiseuille flow driven by the transitional-zone topographic gradient. The underlying mantle lithosphere, density Pro, is stable and moves vertically in response to local isostatic adjustment of the system.

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channel material, ~ = (Pro - Pc)/Pm is the isostatic amplification factor and Pm and Pc are respectively the densities of mantle and crust. The material flux q(x) through the channel, thickness h(x), assuming lithostatic pressure and uniform density of the crust, is s+

pgh 3 0S3 12/x Ox

q = as7 Vx dz -

(2)

where S + and S~- are the boundaries of the channel, /x is the viscosity of the channel material, and S 3 is the upper surface of the crust. The variation in $3 induces lateral pressure gradients which drive the Poiseuille flow. The horizontal velocity inside the channel is Vx(z) =

m(z

- si-)(s~- - z) os3

2~

Ox

(3)

The general model of the channel injection (Fig. 2) assumes three zones in the crust and an underlying stable mantle lithosphere that undergoes local isostatic adjustment. In the crust, material between S + and S 3 and between S2 and SM is pre-existing crust above and below the channel, and the material in this zone moves only vertically and acts as a passive load. Material between S + and S{ is the channel material actively involved in the current flow. The material between S + and S +, and between $1 and $2 is material that accreted to the crust from the channel, and is passive once accreted. Equations 1 and 2 can be solved for a range of assumptions concerning the channel thickness h(x), and the channel viscosity /z(x), as described in the next sections.

C h a n n e l injection m o d e l CI-I: transition zone with uniform p r o p e r t y c h a n n e l This model, in which h and /z are constants, was used by Clark & Royden (2000) to calculate channel flow in the plateau-bounding transition zone and the results were applied to the eastern flanks of the Tibetan Plateau. In addition to the uniform thickness and viscosity they assumed that material accretes (freezes) to the channel walls as soon as it crosses S~, thereby thickening the region between S1+ and S + and S~- and S~-. This material then acts as part of the passive overburden by applying its weight to create a pressure gradient along the channel. In this model the geometry of the channel does not change with time (h = const), even though the velocity field and topography do evolve. Equations 1 and 2 can be combined for

this case, and using $3 - S + = const (i.e. constant thickness of the initial crust above the channel) gives:

0S3

OS+ * p g h 3 0 ( 1 0S3"~ Ot -- 12 Ox \-~--~x ]

(4)

Using the assumptions of Clark & Royden (2000), that the channel has constant viscosity and the flux into the model is constant (/x = const, qb = const), we solved Equation 4 analytically using the Laplace method ((~zi~ik 1980, chapter 7) for the evolution of the upper surface:

S3(x, t) = qb

exp -~-~

- x erfc (5)

where S 3 ( x , t = 0 ) = 0 is the initial position of upper surface, a = (~pgh3)/3tz, erfc is the complementary error function, and t is time. Based on this approach, Clark & Royden (2000) estimated the distribution of viscosity in putative crustal channels under the north, east and SE margins of the Tibetan Plateau using the current topography of those margins ($3). However, as they acknowledge, the results are critically dependent on parameters that cannot be accurately estimated (i.e. the values of h and qb) and whether h, qb and /x are actually constant. If uncertainties in these parameters are taken into account (for example, h ranges from 5 to 20 km), the resulting estimates for channel viscosity, /x, can vary by orders of magnitude even when qb is known because the control parameter o~ scales as h3/i x. Therefore, an uncertainty in h of a factor of slightly more than 2 causes an order of magnitude uncertainty in/x. Figure 3 illustrates a typical result for model CI-1. Assuming constant flux of material into the system, qb = 10- 6 m 3 s- 1per m along-strike (which is equivalent to an average rate of 3.3 mm a -1 in a 101on thick channel), results in elevating the plateau side of the transition zone to 5.5 km in 20 million years. Although the results in Figure 3a were obtained for the special case of 1 0 k m thick channel with /x = 3 x 102~ Pa s, exactly the same results can be achieved for 20km thick channel with /X2o = 2.4 x 1021 Pa s, where the subscript indicates the assumed thickness of the channel in kilometres which yields this estimate of the viscosity. While the results of Clark & Royden (2000) compare well with the topography of the chosen margins of the Tibetan Plateau, there is a need to address additional questions that cannot be

GROWTH OF PLATEAUS BY CHANNEL INJECTION

Fig. 3. Results for channel injection model CI-1. The special case considered here includes:/z = 3 x 1020Pa s, h = 10 km. Flux of material into the system, qb = 10-6 m3 s -1 per m along strike (average rate is 3.3 mm a -1 in 10 km channel), results in the plateau margin being elevated to 5.5 km in 20 million years. Results are the same for constant h3/t z (e.g./z15 = 1021Pa s (for h = 15 km) or /-~2o= 2.4 x 1021Pa s (for h = 20 km). Note: (1) results do not depend on the vertical position of the channel within the crust; (2) model does not address connection with plateau. answered by channel injection model CI-1. For example, how is the flux qb determined dynamically and would the available differential pressure drive this particular flux? How can this model of the transition zone be matched to the geometry of the plateau itself, and how are the overall dynamics of these two regions related?

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can be interpreted to be the consequence of partial melting. Thus, viscosity in the channel is ~ = / z (x, t, D*), and the evolution of the system is described by the same Equation 4 as in the model CI-1. This model can address both the development of the transition zone (d < D*) and the adjacent plateau (d > D*). The boundary flux qb and viscosity /z(d) were chosen in a similar way to model CI-1 (Fig. 3) and to Clark & Royden (2000), so that the edge of the plateau grows to a height of 5.5 km in 20 million years, and also so that geometry of the transition zone is similar to that in the previous section (Fig. 3). For example, the result shown in Figure 4 displays a plateau-like geometry of the orogen if the viscosity of the channel decreases by a factor of 100 under the plateau (/z(d > D*) = / ~ ( d < D*)/100, Fig. 4b,c). The absolute value of viscosity can be estimated, as in the model CI-1, within the uncertainty of the control parameter h3/tz (Fig. 4c). The dynamic link between the plateau and its adjoining transition zone results in more conservative estimates in general of the viscosity under the plateau ( / z ~ 1018-1019 pa s) than estimates presented in Clark & Royden (2000), where estimates based on the model CI-1 gave/x ~ 1016 Pa s for equivalent h values. Although model CI-2 is very simple, it expresses a bulk behaviour similar to full two-dimensional (2D) numerical models of plateau evolution (Vanderhaeghe et al. 2003). Once the plateau is established, it grows laterally and the geometry of the transitional zones remains self-similar (Fig. 4a) in the same manner noted by Vanderhaeghe et al. (2003). A simple kinematic analysis allows the rate at which the plateau edge advances, V,, and the corresponding rate of plateau widening to be estimated. For the boundary conditions implemented here (i.e. the edge of plateau grows from 0 up to 5.5 km in 20 million years) the edge of the plateau advances by the width of the transition zone in the same timescale ( V a ~ 5 m m a - ~ ) . Mass balance relates Va to the boundary flux by qb = VazMr-Ic, and thus

Channel injection model CI-2: constant thickness, variable viscosity channel

The approach of Clark & Royden (2000) can be combined with that of Royden (1996) which used a depth-dependent model of viscosity. Based on the assumption that temperature increases with depth in the crust and that viscosity correspondingly decreases, we propose a model in which the channel develops at the base of the upper crust (Fig. 4). In this model there is a step decrease in the channel viscosity when the centre of the channel, depth d, reaches a critical depth D* (Fig. 4). This decrease

Va = qb/AHc ~ const

(6)

where AHc = Hop - Hco ~ const is the difference in the crustal thickness between the foreland and the plateau. Although the lateral variations of the plateau crustal thickness are small, the plateau crust must thicken internally (Fig. 4a) to drive the subplateau channel, because the rate of flow in the channel is proportional to the local slope of topography. This requirement, however, is an artefact of the thin-sheet approximation used here to describe the system. Beaumont et al. (2004) describes channel

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S. MEDVEDEV & C. BEAUMONT

Fig. 4. Results for channel injection model CI-2. Special case considered here: thickness of the channel h = 10 km; /z --- 102~ Pa s if the depth of the centre of the channel, d(x, t) < D* = 50 km, and/z = 10 ~s Pa s otherwise. Flux of material into the system, qb = 4 • 10 6 m 3 s- 1 per m along strike (average rate is 1.5 cm a- l in l0 km channel) results in plateau margin elevated to 5.5 km in 20 million years. Note: (1) decrease in viscosity under transitional crust leads to plateau-type topography; (2) self-similar widening of the plateau is characterized by the plateau advancing velocit~r Va and by kinematic waveform of evolution of upper surface $3 = S3(x § Vat); (3) results are the same for constant h //x; (4) channel flux varies spatially: under the plateau q ~ qb = const, and in the transition zone flux decreases.

flow under a nearly flat plateau using full 2D numerical calculations for which the driving force for the channel flow is not the local, but global changes in topography. W e therefore simplify m o d e l CI-2 and assume constant thickness plateau crust, H c p = c o n s t , and, correspondingly, A H c = c o n s t in subsequent analysis. The rate o f plateau edge advance, Va, is not a material velocity, nor is it associated with the motion o f any physical part of the system. The self-similar evolution of the plateau is similar to the propagation of a kinematic w a v e such that evolution o f the topography in the transition zone can be expressed as S 3 ( x , t) = S 3 ( x - A t W a , t - At) for any positive x, t and At ( < t ) . W e use this property and assume that other model parameters (for example, temperature in the thermal model; see Appendix) also evolve in the f o r m o f a laterally propagating kinematic wave. Similar geometries and d y n a m i c s of growth (5.5 k m over 20 million years) can be obtained for models CI-1 and CI-2 (Fig. 5) only if the viscosity of the channel under the transition zone in m o d e l

CI-2 is approximately three times smaller than in model CI-1. M o d e l CI-2 also requires a four times higher boundary flux into the system. The

6

C1-1 ~1o = 3"1020 Pa.s (t :20yyr, qb=10 "s m2/s)

5 ~ N'3

,%

CI-2: Blo = 102~ Pa.s (t=50 Myr, qb=4.10 -6 m2/s)

2 1

0 0

,

,

,

i

20

,

,

-

.

.

.

.

40

,'~

60

-In.

"~

.

80

.

.

.

100

X, km

Fig. 5. Comparison of the final geometry of channel injection model CI-1 (dashed line, see details on Fig. 3) and transitional zone predicted by the channel injection model CI-2 designed to achieve the same geometry (solid line, see details on Fig. 4). Similar geometry is, however, obtained for different viscosities and flux into the system because the model CI-2 accounts for the connection to plateau. See text for details.

GROWTH OF PLATEAUS BY CHANNEL INJECTION discrepancy is caused by an oversimplification in model CI-1, which does not account for the additional flux required to build the advancing plateau in addition to the flux into the translating transition zone. Building the plateau margin comprises two processes: first, the development of the transition zone as in model CI-1; and second, the conversion of the margin to plateau, which is the process that requires the extra material flux. In this regard model CI-2 is more complete than CI-1. The kinematic connection between the transition zone and the plateau edge in model CI-2 also results in a time-dependent flux into the transition zone at a fixed location, even though the boundary flux qb is a constant. Under the plateau, where the thickening rate is negligible, the flux in the channel q ~ qb = const. The channel material flux then decreases across the transition zone to zero at the toe. Thus, the model CI-1 assumption of a constant flux into the transition zone cannot be combined with the plateau advancing with a constant rate Va-

C h a n n e l injection m o d e l CI-3: t e m p e r a t u r e dependent channel properties While the resulting geometry of model CI-2 looks realistic (Fig. 4), we question whether the physical foundation of the model is correct. Is depth a good proxy for temperature in determining channel properties? Would a better model based on thermal evolution of the crust result in different channel properties and flow regimes? Vanderhaeghe et al. (2003) compared depth- and temperature-dependent representations of crustal viscosity in models of plateau evolution. Although the topographic shape of the resulting plateau can be similar for both rheologies, the internal structure of the plateau crust is substantially different. In the model in which viscosity decreases with depth (see also Royden 1996; Shen et al. 2001), a lowviscosity weak layer of constant thickness develops below the plateau and is reflected by the flat topography of the plateau. In contrast, the temperaturedependent viscosity model of the crust results in large (up to 20 km) variations in thickness of the low-viscosity weak layer and, consequently, inversely correlated variations in thickness of the overlying upper, strong and cold layer (Vanderhaeghe et al. 2003). The geometry of the weak crustal layer determines the structure of the channel flow in the crust; therefore, we can expect different interactions between the overburden and the channel for the depth- and temperature-dependent models. Here we combine model CI-2 with an approximate thermal model (Appendix) to analyse evolution of the plateau and the transition zone when the viscosity and geometry of the channel is temperature-dependent. We assume that instead of a

153

critical depth, D*, the channel viscosity decreases when the channel temperature reaches a critical value, Z*, corresponding to the onset of significant melt weakening. The channel thickness, initially h0, increases to include the mid- and upper crust where T > 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-<~ h2 h d h3

S3p Mechanical parameters t* g P Pm

ix Thermal parameters T(x, z, t) Ttop

7* qm

Cp

K K A1 A2

Definition

Value

Thickness of the model crust Initial crustal thickness Crustal thickness at the edge of plateau Thickness of deformable part of the crust (Fig. 2) Thickness of the crustal channel in the CI model (Fig. 2) Depth of the middle of the channel Thickness of the lower, undeformable crust (Fig. 2) Uplift of topography at the plateau edge ($30 = 0, Fig. 2)

35km 65 km 10--+ 40kin h0 : 10km do = 2 0 k m

10km 5.5 km

Thickening time Acceleration due to gravity Density of the crust Density of the lithospheric mantle Isostatic amplification factor (1 - P/Pm) Viscosity of the channel material

2 x 107years 9.8 m s -2 2700 kg m -3 3300 kg m -3 0.18 Variable (Pa s)

Temperature (Eqn A 1) Temperature on the top surface (boundary condition for Eqn A1) Critical temperature (Fig. 6) Mantle heat flux (boundary condition for Eqn A l) Heat capacity (Eqn A1) Thermal conductivity (Eqn A1) Thermal diffusivity (= K/pCp, Eqn A l) Volumic rate of heat production in the upper crust (initially 20 km thick, Fig. A2) Volumic rate of heat production in the lower crust (initially between 20 and 35 km deep in the crust, Fig. A2)

(~ O~

Criteria f o r initial and transient consistency of thermal and mechanical models Three criteria can be used to analyse consistency between the mechanical and thermal models.

1.

Initial conditions: if channel flow is the mech-

2.

anism that tectonically thickens the crust, temperatures in the crust must be high enough, and associated viscosity must be low enough, to allow viscous channel flow to occur under the initial ambient conditions. Temperature change: the temperature change in each crustal column, while it is in the transition zone, should be such that the final temperature meets the requirements for channel viscosity to decrease by melt weakening, so that the transition to plateau topography will occur.

Timescale of temperature change: the timescale for increase in temperature and onset of melt weakening of the transition zone must agree with that of the mechanical thickening

750~ 20 mW m -2 750 J kg -1 K - t 2Wm-lK-1 10 -6 m 2 S-1 1.8 p~W m -3 0.75 txW m -3

model, so that the mechanical and thermal plateau transitions are compatible. The first criterion addresses the initial thermal state of undeformed crust. Our reference model (Table 1, Fig. 6) with ho = 10 km is consistent because the temperature at the top of the channel is 400~ (Fig. A2), and the threshold for viscous creep in quartz-controlled crust is approximately 300~ Moreover, any equivalent foreland crust is consistent with this temperature requirement if the radioactive heat production is 0.75 b W m -3 or higher. However, a requirement for a thicker channel, e.g. ho = 15 kin, is marginally too cold (290~ at the top of the channel) unless the thermal parameters are changed. Increasing radioactive heat production would, for example, require A = 5.4 p~W m -3 for the corresponding channel to be 20 km thick. The second criterion determines whether thickening the crust results in efficient heating. Crustal thickening results in transient heating, but this occurs with a characteristic timescale often longer

GROWTH OF PLATEAUS BY CHANNEL INJECTION

155

Fig. 6. Results for channel injection model CI-3. Special case considered here:/z = 1020 Pa s if temperature in the channel is below critical T < T* = 750~ and/x = 4 x 1018 Pa s if T > 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)
Fig. 13. Geometry and forces for simplified analysis of the stability of brittle upper crust of the transition zone (block with bold outline). The brittle block is underlain by a mid-crustal channel and material accreted from the channel. Only forces that contribute to the horizontal balance are indicated.

(13)

(15)

Equation 15 can be solved for the critical limiting value of 4~, q~* (below which the block becomes unstable) using the geometry (Fig. 6) and properties (Table 1) of our reference model. Here we use h~, = 15 km, based on the depth to the 400~ isotherm calculated using the associated onedimensional thermal model (Appendix) and shown in Figure A2. Crust below this level will almost certainly be viscous. If hwp = 40 km, ~b* = 21.5 ~ The block will be unstable if & < oh*. It is also worth noting that erosion which occurs sufficiently rapidly to reduce the thickness of the frictionalplastic layer will tend to destabilize the system by

GROWTH OF PLATEAUS BY CHANNEL INJECTION reducing Fp and/or F~. This effect is seen in the numerical models (Beaumont et al. 2004, fig. 7) and leads to the transition from channel flow beneath stable plateau and transitional crust to unstable outward flow of the upper crust together with the channel. For example, for the current reference model a reduction of the frictional-plastic layer thickness to 10 km would render the block unstable for 4' = 21.5~ and requires an increase to 4'* = 37 ~ if the system is to remain stable. If the system becomes unstable, the overburden fails, the block glides outward with velocity Vc, and the force balance (Equation 10) becomes F~ + Fpv - F~ - Fnx -4- F , - FC - F v = 0

(16)

where the forces Fpb and Fo are replaced by their limiting values because the frictional-plastic overburden has failed. F c = ix L c V c / h is an additional force owing to the basal drag of the Couette flow in the viscous channel on the overburden (Gemmer et al. 2004), and L c = Xp - xo. F v is the additional viscous force that results from the extension of the viscous crust above the channel at the edge of the plateau. This force can be estimated by assuming that the triangular region necks viscously with a horizontal length scale L,, which gives

Fp =

Ii +(x~) -~~x')2tZw-~Vc dz =

2txw(hwe - hb)-~,, Vc

(17)

assuming that the triangular block has a uniform viscosity/Zw in the region that necks. The following results are obtained assuming hb=15km, Lc=100km, Ln=30km, and /Zw = 102~ Pa s, and using other parameter values from Table 1. The maximum velocity Vc = 8 mm a -1 occurs when 4)= 0 ~ and ranges from 6 mm a - l (4, = 5~ through 5 mm a -1 (4, = 9~ to 2.75 mm a -1 (4, = 15~ These results demonstrate that passive plateau margin transition zones can be close to extensional failure and the outward flow velocities can be significant if the viscosity of the crust above the channel at the plateau margin is approximately 1022 Pa s. The outward flow could still be significant on geological timescales even if this viscosity were an order of magnitude higher.

Conclusions Simple modelling has been used to improve our understanding of the first-order controls on the growth of continental plateaus and the interactions that occur within the transition zone at the plateau margin. The simplified approach of this study also helps us to gain more insight into the results of

161

numerical models which produce realistic-looking geometries, but are difficult to interpret in terms of the basic physical controls (Beaumont et al. 2001, 2004; Jamieson et al. 2004). Three channel injection (CI) models. We developed successively more complete approximations to the evolution of a continental plateau and adjacent transition zone by the CI mode. CI-1, the simplest approximation (based on Clark & Royden 2000), assumes a constant viscosity and thickness channel in which excess material accretes in zones above and below the channel. While the predictions of this model compare well with some natural examples, this model does not consider relations between the plateau and transition zone. CI-2 includes a decrease in channel viscosity, predicted to occur on melt weakening or at high temperatures, when the crustal thickness exceeds a critical value, D*. This model completes the connection between plateau and transition zone, but relies on an arbitrarily chosen critical thickness of the crust. CI-3 is more realistic and considers the channel viscosity and thickness to be temperaturedependent. CI-3 includes an associated thermal model which accounts for radioactive self-heating of the transition crust and for heat advection as channel material flows from beneath the plateau. CI-3 demonstrates self-consistent plateau widening if the channel viscosity decreases when its temperature exceeds a critical value T*. The simplicity of the models allows analytical solutions for some cases, and provides direct estimates of the parameters that control channel flow. Acceptable results for the topography of Tibet, for example, are achieved by model CI-3 when channel viscosity is in the range 1018-1019 Pa s beneath Tibet and 1019-1022 Pa s beneath the adjacent areas, and the critical temperature for onset of partial melting is in the range 700-750~ These values agree with those determined from fully coupled thermal-mechanical models. Consistency o f thermal a n d mechanical models. The first thermal consistency test requires the initial thermal-viscous structure of the crust to be compatible with crustal thickening by channel injection. The second thermal consistency test asks whether the temperature reaches the melting point, thereby initiating melt weakening, at the appropriate place and time. Comparison of the thermal evolution of model CI-3 with a model in which transitional crust tectonically thickens uniformly (PS) shows that CI is a much more effective mechanism for heating the transitional crust. The third, transient, thermal consistency test assesses whether the development of the melting conditions in the transitional crust coincides with the end of thickening. Premature melt weakening will, for example, result in flattening the transitional-zone

162

S. MEDVEDEV & C. BEAUMONT

topography before it reaches the level of the plateau, and that is incompatible with lateral growth o f the plateau. The requirement for consiste n c y significantly restricts the range of compatible p a r a m e t e r values. Rheological consistency. Rheological consiste n c y compares the m o d e l predictions o f channel viscosity for the transitional crust with viscosity estimates f r o m laboratory experiments. Modelling o f wide ( 1 . 5 - 2 x 103 k m ) transitional zones, for example, with the CI-3 m o d e l requires channel viscosity to be about 2 - 1 0 x 1017 Pa s, w h i c h is incompatible with laboratory data for quartz-based rheologies without 'melt weakening'.

Viscous thickening (VT) and plastic translation (PT) of the transitional crust. The compatibility o f the CI models, which assume only vertical passive motion o f the upper crust, was compared with possible V T and PT deformation of the upper crust. Predicted CI-3 topography o f the transition zone can be r e p r o d u c e d by viscous thickening o f the crust, but corresponding V T model estimates o f the crustal viscosity are up to two orders of magnitude higher than those for CI. This result indicates that topography alone is not a sufficient basis from which to estimate crustal viscosity. The m o d e o f deformation must also be known. Stability estimates for the failure and outward flow o f the plateau margin, PT m o d e deformation, indicate that neutral plateau transition zones m a y be close to failure by plastic translation of the crust above a channel.

determined from the first and second spatial derivatives from 2D calculations), term 6 of the 2D thermal equation

OT A Ot--pCp

OT OZT OT 02T VZOZ+K~za--Vx~x+KOx~

(A1)

may be neglected. This assumption is also in agreement with the thin-sheet character of our mechanical model, where horizontal variations are assumed to be small. In Equation A1, T is the temperature, t is time, A is the rate of radioactive self-heating, Cp is the specific heat at constant pressure, Vz and Vx are vertical and horizontal components of velocity, K = K/(pCp) is the thermal diffusivity, and the numerical values of the parameters are given in Table 1. Pure shear (PS) model. A specific example of pure shear crustal thickening is solved. This model is then used as a reference for comparison with the CI models. We assume that upper and middle crust is thickened at a uniform rate by a factor of 2.2 in 20 million years resulting in a total crustal thickness of 65 km, at which time the crust is assumed to become stable plateau crust (Fig. A1). In addition to the simplification noted above,

The research was funded by Natural Sciences and Engineering Research Council grant to C.B.S.M. has also been supported by the Deutsch Forschungsgemeinschaft (Sonderforschungsbereich Grant 267). We thank the reviewers, R. Buck and T. Gerya, and the lead Editor, R. Law, for their constructive comments. M. Clark is thanked for providing data on Tibet margins topography.

Appendix: the associated thermal model This appendix develops an approximate onedimensional (1D) vertical model for temperature in the crust that is thickened by channel injection. It is used in conjunction with channel injection models to investigate how channel flows may develop when viscosity depends on crustal temperature. To emphasize the role of the mode of thickening, we compare the CI mode with the model in which the upper and middle crust thickened by pure shear (PS) mode. Vanderhaeghe et al. (2003) demonstrated that the 2D cross-sectional thermal evolution of plateau crust can in some circumstances be explained approximately by analysis of a reduced vertical 1D thermal problem. We use the same approach here. Based on the result that vertical temperature variations are much larger than horizontal ones (as

Fig. A1. Thermal evolution of pure shear model. (a) Evolution of temperature at base of the middle crust (5 km above lower crust, dotted line in b). (b) General view of evolution of a vertical Lagrangian crustal column. Dark grey indicates upper and middle crust that thickens uniformly, light grey is lower, undeformable crust. The dash-dot line separates crust into highradioactivity (A = A l) and low-radioactivity (A = A1) parts. The underlying mantle lithosphere (the depth of the model is 90-100 km) has zero radioactivity. The isotherms (T* = 750~ solid line; T = 700~ dashed line) indicate compatibility of the thermal and mechanical models.

GROWTH OF PLATEAUS BY CHANNEL INJECTION in this model we also neglect horizontal advection (term 5, Equation A1) on the basis that this term is much smaller than vertical advection (term 3, Equation A1) because both velocity components are small, but the vertical temperature gradient is larger. The simplified equation is solved numerically and Figure A1 shows the result for the reference model. Note that the upper, highly radioactive part of the crust thickens significantly from an initial 20 km to 44 km. Channel injection model CI-3. To apply the thermal model to channel injection in the CI-3 model we use the self-similar nature of the lateral growth of the plateau (Figs 4 and 6). We assume that the temperature field also evolves in a self-similar manner, that is, T(x, z, t) = T(x + VaAt, Z, t + At). This implies that during steady-state plateau growth, the vertical thermal structure migrates laterally with the same velocity as the plateau edge, V, (Equation 6). Furthermore, in CI-3 the horizontal advection of heat cannot be neglected because Vx in the channel may be large. The self-similarity of the temperature profile gives OT Ox

.

.

l O T ,~ .

Va at

.

l(p~p

lZ 3T K--O2T'~ z - - -4-

Va

OZ2]

' OZ

(A2)

where the first equality results from aT 3t

--

163

where ho is the initial thickness of the channel. In addition, the depth-dependence of the horizontal velocity Equation 3 can be expressed via flux in the channel, qp, using Equation 2:

ydz)

6qp

= ~-(z

(A7)

- s ; ) ( s + - z)

Combining Equations A6 and A7 we obtain channel heat advection coefficient inside the channel of thickness h: Oc

V~ -- Va --

6(h2 - ho) h----------~ ( z -

S7)(S

+ -

z)

(A8)

We solve Equation A1 (for the pure-shear model) and Equation A4 (for the CI-3 model) numerically using a Lagrangian approach and a finite-difference approximation. The 1D mesh consists of 200-300 grid points with variable density of the mesh (more dense in the upper and mid-crest). The bottom boundary condition (heat flux, qm) is set initially at the depth of 95 kin. To illustrate the results we consider two model properties: the temperature evolution for a point located 5 km above the base of the upper crust (Fig. A2a), and the time when the temperature at this location reaches the critical value (Fig. A2a,b). The model behaviour will be consistent with the channel injection behaviour and melt

T(x, z, t -4- At) - T(x, z, t) lim At-~0 At

= lim at-~o

T(x, z, t + At) - T(x + VaAt, z, t + At)

At 8T

(A3)

The approximate equality in Equation A2 is the zerothorder approximation of the horizontal temperature gradient in the system. Combining Equations A1 and A2 gives the higher-order approximation for the thermal evolution of crust thickened by channel injection:

3T

--

Ot

-----

(l§

a

E OT

-

z ~ Z -4-

K--

OZ2

(A4)

Here, 0c is the channel heat advection coefficient. It is zero outside of the channel and 0c = Vx/Va inside the channel. It can be seen that the horizontal channel flux accelerates the thermal evolution in the vicinity of the channel by a factor (1 -4- 0o). The distribution of Oc can be found from the mass balance. First, Equation 1 and self-similarity of geometry of the transition zone give Oqp _ 3x

3h2 Ot

--

Oh2 Ox

V a -

(A5)

The requirement that qp = 0 in the undeformed foreland allows integration of Equation A5: qp = Va(ha - ho)

(A6)

Fig. A2. Thermal evolution of channel injection (CI) model. (a) Evolution of the mid-channel temperature. (b) General view of thermal and thickening evolution of a vertical Lagrangian crustal column. Levels of grey (from lighter to darker) indicate: upper and low undeformable crust; material accreted from the channel; high-viscosity channel in the transition crust; lowviscosity channel beneath the plateau. See details of design in Figure A1. Note that highly radioactive portion of the crust thickens much less than in pure shear model.

164

S. MEDVEDEV & C. BEAUMONT

weakening to form a plateau if, at 20 million years, the lower part of the upper crust reaches a high enough temperature, c. 750~ (At = 0, meaning the temperature is reached at the specified 20 million year time). Significant lead (At < 0) or lags (At > 0) in the time at which this condition is achieved implies the thermal evolution is not consistent with the mechanical model which assumes that the transition to a plateau occurs at 20 million years. The same consistency conditions can be applied to the plateau model in which the transition zone thickens by pure shear mode (Fig. A1). We consider two types of thermal consistency. The CI-3 model requires an exact match between the mechanical and thermal models (7* = 750~ At = 0). However, in the consistency analysis our uncertainty in the critical temperature is acknowledged, and we consider a range, 700-750~

References BABEYKO, A. YU., SOBOLEV, S. V., TRUMBULL, R. B., ONCKEN, O. & LAVIER, L. L. 2002. Numerical models of crustal scale convection and partial melting beneath the Altiplano-Puna plateau. Earth Planetary Science Letters, 199, 373-388. BEAUMONT, C., JAM1ESON, 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. & [VIEDVEDEV, 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 / 2003JB 002809. BIRD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit.

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of

Geophysical

Research,

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10275-10286. BUCK, W. R. & SOKOUTIS, D. 1994. Analogue model of gravitational collapse and surface extension during continental convergence. Nature, 369, 737-740. CLARK, M. K. & ROYDEN, L. H. 2000. Topographic ooze: Building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703-706. GEMMER L., INGS, S., MEDVEDEV, S. & BEAUMONT,C. 2004. Salt tectonics driven by differential sediment

loading: Stability analysis and finite element experiments. Basin Research, 16, 199-218. DOI: 10.1111/j.1365-2117.2004.00229.x. GLEASON, G. C. & TULLIS, J. 1995. A flow law for dislocation creep of quartz aggregates determined with the molten salt cell. Tectonophysics, 247, 1-23. 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. DOI: 10.1029/2003JB002811. JOHANNES, W. & HOLTZ, F. 1996. Petrogenesis and

Experimental

Petrology

of

Granitic

Rocks.

Springer-Verlag, Berlin. KAUFMAN, P. S. & ROYDEN, L. H. 1994. Lower crustal flow in an extensional setting: Constraints from the Halloran Hills region, eastern Mojave Desert, California. Journal of Geophysical Research, 99, 15723-15739. LEHNER, F. K. 2000. Approximate theory of substratum creep and associated overburden deformation in salt basins and deltas. In: LEHNER, F. K. & URAI, J. L. (eds) Aspects of Tectonic Faulting. Springer-Verlag, Berlin, 21-47. MEDVEDEV, S. 2002. Mechanics of viscous wedges: Modeling by analytical and numerical approaches. Journal of Geophysical Research, 107, B6, 10.1029/2001JB 000145. MEDVEDEV, S. E. & PODLADCHIKOV, Yu. Yu. 1999. New extended thin sheet approximation for geodynamic applications - II. 2D examples. Geophysical Journal International, 136, 586-608. 0ZI~IK, M. N. 1980. Heat Conduction. John Wiley, New York. 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. SHEN, F., ROYDEN, L. H. & BURCHFIEL, B. C. 2001. Large-scale crustal deformation of the Tibetan Plateau. Journal of Geophysical Research, 106, 6793-6816. VANDERHAEGHE, 0., MEDVEDEV, S., BEAUMONT, C., FULLSACK, P. & JAMIESON, R. A. 2003. Evolution of orogenic wedges and continental plateaux: insights from crustal thermal-mechanical models overlying subducting mantle lithosphere. Geophysical Journal International, 153, 27-51.

Provenance of the Greater Himalayan Sequence and associated rocks: predictions of channel flow models R. A. J A M I E S O N 1, C. B E A U M O N T 2, M. H. N G U Y E N 1'2 & D. G R U J I C 1

aDepartment of Earth Sciences, Dalhousie University, Halifax, N.S., Canada, B3H 3J5 (e-mail: [email protected]) 2Department of Oceanography, Dalhousie University, Halifax, N.S., Canada, B3H 4J1 Abstract: Numerical models for channel flow in the Himalayan-Tibetan system are compatible with many tectonic and metamorphic features of the orogen. Here we compare the provenance of crustal material in two channel flow models (HT1 and HT111) with observations from the Himalaya and southern Tibet. Thirty million years after the onset of channel flow, the entire model crust south of the India-Asia suture still consists only of 'Indian' material. The model Greater Himalayan Sequence ('GHS') is derived from Indian middle crust originating < 1000 km south of the initial position of the suture, whereas the Lesser Himalayan Sequence ('LHS') is derived mainly from crust originating > 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 (_<5%) of in situ partial melt (e.g. van der Molen & Paterson 1979; Rosenberg & Handy 2005). However, any other geological process producing a reduction in effective viscosity by a factor of about 10 over the same temperature range would have a similar effect in the model. Exhumation in both models is controlled by surface denudation, which varies with time, distance and local surface slope (Beaumont et al. 2004; Jamieson et al. 2004). The time- and spacedependent denudation functions, identical in models HT1 and H T l l l , are pre-defined (f(t), g(x); see Fig. 1, Table 1). Until 30 Ma there is no erosion; this is followed by a high erosion rate until 15 Ma and a gradually declining erosion rate until the end of the model (details in fig. 10 of Jamieson et al. (2004) and appendix of Beaumont et al. (2006)). Slope-dependent surface denudation is focused on the pro-side flank of the orogenic plateau. The local surface slope is determined by the interaction of model tectonics with the imposed denudation functions, leading to some differences in denudation patterns between the two models (Figs 3 & 4).

Crustal-scale model results Tectonic and thermal results from model HT1 have been presented elsewhere (Beaumont et al. 2004; Jamieson et al. 2004) and only those features relevant to provenance are summarized here. Channel flow initiates at c. 30Ma and channel extrusion begins at c. 15Ma (Fig. 3). Initially the extruded channel is symmetrical, but a transition to an asymmetrical overthrust structure occurs between 15 and 6 Ma (Fig. 3) in response to detachment and outward flow of overlying upper crust (e.g. Medvedev & Beaumont 2006; fig. 16d of Beaumont et al. 2004). 'Tethyan' crust above the channel remains thick and strong enough to retain its coherence, and model HT1 does not develop domes between the suture and the erosion front (Fig. 3).

PROVENANCE AND CHANNEL FLOW MODELS

167

Fig. 1. Model parameters and initial conditions; a full list of parameters is presented in Table 1. (a) Mechanical model showing initial distribution of crustal layers. (b) Initial thermal structure showing temperature (isotherms, 100~ intervals) and velocity fields (short lines); the model is started in conductive steady-state, with Vp = Vs = 0. (c) Surface denudation model. Local surface slope calculated by model, f ( t ) controls intensity of erosion for specified time intervals; g(x) specifies erosion only on pro-side plateau flank. Further explanation in text and Beaumont et al. (2006, appendix). (d) Mechanical and thermal properties of crustal layers. Upper and middle crustal layers subject to melt weakening. Model HT111 differs from HT1 only in having a thin weak layer (~beff= 2 ~ 4.0-7.5 km initial depth) embedded in the upper crust (column at far right). Further details in text, Jamieson et al. (2004), Beaumont et al. (2006, appendix).

R.A. JAMIESON ETAL.

168

T a b l e 1. Parameters used in models HT1 and H T l l l (see also Fig. 1) Parameter

Mechanical parameters Vp VR Vs

0 D Pc.rust Pmantte

HTI: HTlll: HT 1 and HT 111: ,~v

eft

= B*('f2)(1-n)/2n

.! exp[Q/nRTK]

12 R TK B*(WQ) ( 0 - 1 0 km) B*(WQ • 5) ( 1 0 - 2 5 km) B*(DMD) ( 2 5 - 3 5 km)

'melt weakening' (B* (WQ) and B*(WQ x 5) only)

Thermal parameters pC,,(~T/at + v 9 VT) = KVZT+A Ce K

Meaning Pro-side (convergence) velocity Retro-side velocity S-point velocity (subduction advance) Subduction dip angle Flexural rigidity (isostasy model) Crustal density Mantle density Effective internal angle of friction upper crust upper crust Middle and lower crust General equation for effective viscosity Second invariant of strain rate tensor Gas constant Absolute temperature Wet Black Hills quartzite flow law (after Gleason & Tullis, 1995) Modified wet Black Hills quartzite flow law Dry Maryland diabase flow law (after Mackwell et al. 1998) Linear reduction in effective viscosity over T range 7 0 0 - 7 5 0 ~

q,,. A1 ( 0 - 2 0 kin) A 2 ( 2 0 - 3 5 km) TMoho Surface denudation slope • f(t) • g(x) slope f(t)

Denudation model Local surface slope (pro-flank of plateau) Time function

g(x)

Spatial (climate) function

T~ qm

5 cm a -1 O c m a -1 2.5 cm a - I 20 ~ 1022 Nm 2700 k g m -3 3300 k g m -3

5~ (0-10 ~ ) 5 ~ (0-4.5 & 7 - 1 0 ~n) 2 ~ ( 4 . 5 - 7 km) 15 ~ ( 1 0 - 3 5 km)

1S -1

8.314 J mo1-1 K - l K n = 4.0 B* = 2.92 • 106 Pa S1/4 Q = 223 kJ mol-1 B* = B*(WQ) • 5 n, Q as above n = 4.7 Q = 485 kJ mo1-1 B* = 1.91 • 105 Pa s 1/47 r/7oo = flow law value r/75o = 1019 Pa s

Heat balance equation Heat capacity Thermal conductivity Thermal diffusivity (K = K/pCp, where pCp ~- 2 • 106) Temperature at lithosphere/asthenosphere boundary Basal mantle heat flux Initial surface heat flux Upper crustal heat production Lower crustal heat production Initial temperature at Moho

K

Value(s)

750 m 2 K 1 s - 2 2.00 W m - l K -1 1.0 • 10 -6 m 2 s 1 1350~ 20 m W m -2 71.25 m W m -2 2.0 p~W m -3 0.75 p~W m 3 704~

0.107 m a - l for t = 3 0 - 1 5 Ma declining for t> 15Ma 0 --+ 1 (dry --+ wet)

Model HT111 differs from HT1 only in having a weak (&eft = 2~ layer embedded in the upper crust (details below).

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Fig. 2. Material distribution at the start of the model (54 Ma). Shaded blocks in the mid-crustal layer are used for tracking purposes only; there is no lateral difference in material properties. The initial width of each block is 200 kin. Blocks are labelled for convenient cross-referencing with text: Ia-Ij originate on the 'Indian' side with Ia closest to the suture; Au, Am, A1 correspond to 'Asian' upper, middle and lower crust respectively; no lateral distinctions are made on the 'Asian' side. Labelled symbols show the initial positions of selected tracked points in model HT1 (squares) and HT111 (circles). 'GHS' P-T-t paths from tracked points G1, G2 (model HT1), and D1, D2, D3 (model H T l l l ) are shown in Figure 9; 'LHS' points (L3, DL) are shown for illustration only; P-T-t path for L3 was presented by Jamieson et al. (2004). Model HT111 differs from model HT1 only in having a thin, weak layer (~beg= 2 ~ embedded in the upper crust between 4.5 and 7.0 km depth. Until c. 24Ma the model evolves in a very similar manner to HT 1 (Fig. 4), but the weak upper crustal layer affects its subsequent tectonic evolution by facilitating detachment and outward flow of the crust overlying the channel. Moreover, because the incoming foreland upper crust is also weaker than its equivalent in HT1, it offers less resistance to outward-flowing upper and middle orogenic crust. Consequently, the orogen in model HT111 propagates much further to the south than it does in HT1 (cf. Figs 3 & 4), and for the denudation model used here the region between the suture and the orogenic front is much wider (c. 370 km versus c. 200 km). Enhanced southward flow of orogenic crust in H T l l 1 also enhances advection of isotherms with the channel, forming an inverted thermal 'lobe' that extends into the middle and upper crust (e.g. Fig. 4 at 6 Ma). Incoming middle crust is cooler and stronger than the overlying channel, and takes some time to reach the 700~ melt-weakening threshold. The effect in the model is to create a strong ramp consisting of cool 'Indian' middle and lower crust. The hot channel flow zone is forced up and over this ramp, further destabilizing the overlying upper crust. Outward flow of weak upper crust above the channel forms a dome underlain by channel material. The first dome forms at c. 10Ma, and is subsequently

translated southward along with the channel ('egg-in-snake' effect) and extruded at the erosion front between 5 and 0 Ma. This process is similar to the creation and expulsion of hot fold nappes in response to collision with a strong lower crustal indentor (Beaumont e t al. 2006). Structures initially formed in a dome over a mid-crustal ramp may evolve into strongly flattened nappes during extrusion (Grujic et al. 2004). Continued convergence leads to the creation of a second dome, which at 0 Ma lies c. 200 km north of the erosion front at a depth of c. 10 km (Fig. 4). A more detailed analysis of the mechanics of dome formation in large hot orogens will be presented elsewhere. In both models, outward flow of middle and upper plateau crust above the channel accompanies southward translation of the model suture (embedded in upper plateau crust) from its initial position directly over the S point to a final position c. 400 km south of the S point (Figs 3 & 4). As a result, the entire model crust south of the surface exposure of the suture consists of 'Indian' material. In both cases, the base of the model orogen as far north as the S point consists of a 10 km thick layer of undeformed 'Indian' lower crust (Figs 5 & 6). Because this layer is subducted it is not involved in pro-side crustal deformation (cf. model 1 of Beaumont et al. 2004), does not become incorporated into the channel, and does not extend north of the S point.

170

R.A. JAMIESON E T A L .

Fig. 3. Tectonic and thermal evolution, model HT1, for selected time steps; equivalent results for other times are presented by Beaumont et al. (2004) and Jamieson et al. (2004). All times are given in Ma (millions of years before end of model) to facilitate comparison with observations. &x = total amount of convergence (km). The upper panel in each pair shows a deformed passive marker grid and distribution of mid-crustal materials (shaded blocks) whose initial distribution is shown in Figure 2. Short vertical arrow marks surface position of suture in each panel; subsurface suture is shown by heavy black dots (best seen in lower panel of each pair). The graph above the model surface shows the spatial distribution of surface denudation for each time step; the vertical axis shows erosion rate (scale bar = 10 mm a-1). Maximum erosion rate (ema~) for each time step is also noted. Times of first appearance of GHS and LHS detritus (text in boxes) after DeCelles et al. (2004). Positions of 'MCT' and 'STD' at 6 Ma and 0 Ma determined by inspection of material distributions in the model; see Figures 7, 8, and Jamieson et al. (2004) for further detail. The lower panel in each pair shows the corresponding temperature field (100"C contours; 700~ isotherm = melt-weakening threshold), velocity field (short black lines) and distribution of heat-producing material (shaded region corresponds to material A1; Table 1, Fig. 1) (results from similar models and animations can be viewed at http://geodynamics.oceanography.dal.ca).

PROVENANCE AND CHANNEL FLOW MODELS The behaviour of upper and middle plateau crust in HT111 is affected by the embedded weak layer, leading to a significant difference in crustal structure between the two models north of the suture. In model HT1, the upper crustal layer immediately north of the suture (Au, Fig. 5) thickens significantly during the early stages of

171

convergence, leading to necking of the underlying mid-crustal layer (Am) and its eventual detachment from strong 'Asian' lower crust (A1) after c. 24Ma. By 0 Ma, 'Indian' crust not only underlies the entire region between the erosion front and the suture, but extends beneath 'Asian' upper and middle plateau crust for another 200 km to the north. In contrast,

Fig. 4. Tectonic and thermal evolution, model HT111. Details as in Figure 3. At 24 Ma the model is very similar to HT1, but note formation and expulsion of domes following 15 Ma.

172

R . A . J A M I E S O N ETAL.

9

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oo

!:::~ O ,..~ "~ O,3 t"-,I

d

~~ O,..~

az2, ~ o

""'

PROVENANCE AND CHANNEL FLOW MODELS

173

~az~

/, N ,~~

~ ."= -~ L,

.~~

9

o

it3

~ ,kT..~ o,~ -.-.,

174

R.A. JAMIESON ETAL.

in model HT 111, the weak upper plateau crust (Au, Fig. 6) does not thicken significantly. 'Asian' upper and middle crust (Am) remains coupled to 'Asian' lower crust (A1), and all are transported southward along with the suture above a thin (<20 kin), highly attenuated zone of 'Indian' crust. At the end of the model HT111, the upper two-thirds of plateau crust north of the suture consists of 'Asian' material, whereas in HT1 about half the upper plateau crust in the same region consists of 'Asian' material. We do not regard the model predictions for deformation in the vicinity of the suture to be robust in detail - for example, the properties of the suture zone and Indian versus Asian crust in the Himalayan-Tibetan system probably differ significantly from the laterally homogeneous material properties assumed here. However, the results illustrate that the behaviour of both 'Indian' and 'Asian' model crust is sensitive to subtle variations in upper crustal strength. The models also predict that 'Asian' upper and middle crust north of the suture, to a depth of c. 25-30 km, remained at T _< 700~ (Figs 3 & 4), and was therefore not part of an active channel, consistent with observations from exhumed mid-crustal levels of the Lhasa block (e.g. Kapp et al. 2005) and recently published magnetotelluric data (Unsworth et al. 2005).

Provenance of GHS and LHS In the Himalayan-Tibetan orogen south of the Indus-Yarlung Tsangpo suture, the entire crust, including the GHS and LHS, is interpreted to consist of the deformed and metamorphosed continental margin of northern India (e.g. DeCelles et al. 2000, 2004; Myrow et al. 2003; Richards et al. 2005). The isotopic signature of the GHS (eNa - 10 to - 14, TDM < 2.0 Ga, detrital zircons 500-1700Ma) indicates that it is a Neoproterozoic to early Palaeozoic Gondwanan succession, whereas the isotopically distinct LHS (eNa - 1 5 to - 2 0 , TDM > 2.0 Ga, detrital zircons 1600-2600 Ma) is inferred to represent a Palaeoto Mesoproterozoic sequence (e.g. Parrish & Hodges 1996; Hodges 2000; DeCelles et al. 2000, 2004; Richards et al. 2005). These isotopic contrasts have recently been used to delineate the MCT in central Nepal (Martin et al. 2005), although isotopic similarities between some LHS and GHS intervals suggest a common source area for part of their depositional history (e.g. Ahmad et al. 2000; Richards et al. 2005). Within the GHS, lithological units are laterally continuous along-strike for considerable distances, and a consistent regional lithostratigraphy can be recognized through much of the orogen (Hodges 2000, and references therein).

In models HT1 and HT111, the entire region between the model suture and the orogenic foreland is occupied by 'Indian' material (Figs 5 & 6). On a crustal scale, therefore, material distribution in the models is compatible with observations. The model 'MCT' (Figs 7 & 8) is the protolith boundary between outttowing (GHS) and inflowing (LHS) material as determined from the model at the end of its evolution (0 Ma). In model HT1, the offset across the 'MCT' is >600 km (Jamieson et al. 2004), whereas in H T l l l , the missing section is >400 km when dome material is taken into account (Fig. 2). In both cases, 'GHS' at the model surface is derived from distal 'Indian' upper and middle crust originating <1000km south of the model suture, whereas 'LHS' material is derived mainly from proximal crust originating >1400 km south of the suture (Figs 2, 7 & 8). The intervening material has largely been exhumed and removed by erosion during earlier stages of model evolution (e.g. Figs 5 & 6; 15 Ma versus 6 Ma). At the end of model HT1, the 'GHS' at the surface consists of distal 'Indian' middle crust, with material Ib overlain by material Ic, bounded by the 'MCT' at the base and the 'STD' at the top (Figs 2, 5 & 7). The exhumed material is derived from originally contiguous crustal elements (Ib, Ic) that flowed outward together in the upper part of the channel. Although these materials were subjected to extreme attenuation (Figs 5 & 7), their extrusion from mid-crustal depths to the model surface did not involve mixing with other crustal components. In nature, equivalent materials might appear to display a coherent lithotectonic 'stratigraphy', although intense deformation during channel flow and extrusion would likely obliterate true stratigraphic relationships. In contrast to the 0 Ma material distribution at the model surface, diverse 'Indian' middle crustal elements at depth have been interleaved by folding and transposition during ductile flow beneath the plateau (Figs 5 & 7). If these deeper levels of the channel were to be exhumed, it is extremely unlikely that laterally coherent lithological units could be recognized. In model HT1, the 'LHS' at 0 Ma consists of proximal 'Indian' middle crustal material Ih overlain by material Ii, with Ii separated from 'GHS' materials Ib and Ic by the 'MCT' (Figs 2, 5 & 7). Both 'LHS' materials have been extremely attenuated during their incorporation into the orogen and subsequent exhumation in the footwall of the channel, but have not been mixed with other materials in the exhumed region. As noted by Jamieson et al. (2004) the distribution of materials in the exhumed 'LHS' suggests regional stratigraphic inversion in the footwall of the model 'MCT'. This is not observed in nature (e.g. Robinson et al. 2003), indicating the limitations of

PROVENANCE AND CHANNEL FLOW MODELS

175

Fig. 7. Distribution of Lagrangian material points in the vicinity of the orogenic front, model HTI. Crustal-scale results shown in Figure 5. Both 6 Ma and 0 Ma (48 and 54 million years respectively since start of model) shown in order to bracket the likely range of times since the India-Asia collision (e.g. Hodges 2000; Myrow et al. 2003; DeCelles et al. 2004). Each symbol corresponds to a Lagrangian point tracked by the model. 'LHS' points (Ih, Ii) are shown with diamonds and 'GHS' points (11o,Ic) with circles. Model 'MCT' is the boundary between 'GHS' and 'LHS' materials, and the model 'STD' is the upper bound of 'GHS' material. Initial distribution of crustal blocks and tracked points for P-T-t paths (numbered squares) shown in Figure 2. Only those mid-crustal materials at the surface near the end of the model are shown. These have been separated for clarity as follows: panel set (a), materials Ib, Ic, Ih, Ii; panel set (b) materials Ib, Ii; panel set (c) materials Ic, Ih. Note that the exhumed 'GHS' corresponds mainly to the upper part of the channel flow zone. Material at the model surface originates from contiguous crustal domains (Ib, Ic in 'GHS', Ih, Ii in 'LHS') and is not mixed significantly with 'exotic' material during channel flow and extrusion. Asian upper and middle crust (Aura, grey region at far right of 0 Ma panels) is not transported significantly south of the suture (<20 km) at 0 Ma.

176

R.A. JAMIESON E T A L .

the HT1 model style to represent foreland foldand-thrust-belt deformation realistically. The creation and extrusion of domes in model H T l l 1 leads to a more complex, two-stage evolution for the 'GHS'. At 6 Ma, before the emerging dome has reached the surface, material distribution in stage 1 'GHS' resembles that at the end of model HT1 (Figs 7 & 8), with originally contiguous distal 'Indian' materials (Ib, Ic, Id; Fig. 2) extruded as attenuated sheets between the 'MCT' and 'STD' (Figs 6 & 8). These materials are not mixed with other crustal elements during extrusion and might therefore resemble a deformed but coherent lithostratigraphic sequence in nature. At 6 Ma, the 'LHS' in HT111 consists of proximal 'Indian' materials Ii and Ih with a geometry similar to that displayed by HT1. At this stage, the still-buried dome consists of 'Tethyan' upper crust underlain by distal 'Indian' materials (Ie, Id, Ic, Ib) derived from the upper part of the channel, which are in turn underplated by proximal 'Indian' materials (Ii, Ih) derived from the leading edge of the underthrust ramp. In contrast to model HT1, the final stage of dome extrusion (5-0 Ma) in HT111 strongly affects the geometry and distribution of material above the 'MCT'. The previously extruded stage 1 'GHS' becomes progressively shortened as it is overridden from the north, forming an upright syncline (Fig. 8). The extruded dome is dominated by distal 'Indian' materials (Ie, Id, Ic, Ib), resembling stage 1 'GHS' materials in both lithology and metamorphic grade. Model HT111 therefore predicts that at 0 Ma the exhumed GHS should comprise structurally lower stage 1 and structurally higher stage 2 components. The total width of the exposed high-grade region is considerably wider in H T l l l (c. 50 kin) than in HT1 (c. 15 km). Although it could be very difficult to distinguish between the two 'GHS' packages in nature, peak grade and age profiles from model HT 111 (Jamieson, unpublished data) suggest that discontinuities in structure, metamorphism, and age should be detectable with sufficiently detailed observations. For example, it might be possible to recognize distinct thrust- and normal-sense structures bounding the different high-grade 'GHS' packages, and stage 1 'GHS' should have significantly older peak metamorphic and cooling ages than stage 2 'GHS'. A more detailed discussion of the tectonicmetamorphic consequences of dome extrusion will be presented elsewhere. The extruded dome and its buried equivalent also contain minor amounts of underplated and infolded proximal 'Indian' materials (If, Ih, Ii) that lie below the model surface at 0 Ma (Fig. 8). A transect through a deeply dissected dome might therefore reveal a transition from 'GHS' protoliths into rocks with 'LHS' crustal signatures. This transition,

if it exists in nature, is unlikely to be exposed in any present-day NHG domes (see below), but its strongly deformed equivalent may be preserved in domes extruded at the orogenic front. Both models HT1 and HT111 predict that the 'LHS' should be derived from proximal 'Indian' crust that should be distinguishable from distal 'GHS' protoliths, consistent with most observations from the Himalaya (e.g. Parrish & Hodges 1996; DeCelles et al. 2000; Richards et al. 2005). Furthermore, some general predictions can be made about the provenance of sediments eroded from the orogen after 30 Ma (onset of erosion in both models). Initially, the material being eroded consists entirely of 'Tethyan' upper crust, but by 24 Ma 'Indian' mid-crust is being exhumed at the orogenic front. The first material to be exposed originates c. 1000 km south of the suture (Figs 3 & 4) and would probably have isotopic and detrital zircon signatures more akin to 'GHS' than to 'LHS' lithologies, which are not exhumed until after c. 20Ma. By 15 Ma, continued convergence and channel extrusion have juxtaposed distal 'GHS' protoliths (above the 'MCT') with more proximal 'LHS' material (below the 'MCT') at the erosion front, and this situation persists until the end of the model. Detrital zircon and isotopic data show significant changes in the provenance of Cretaceous to Miocene sedimentary rocks in the Lesser Himalaya (e.g. Najman & Garzanti 2000; DeCelles et al. 2004; Najman et al. 2005). Tertiary strata were deposited in the Himalayan foreland and became incorporated into the orogen as it propagated to the south; these rocks therefore preserve a record of the early unroofing history of the orogen. Eocene and older strata were derived from Tethyan source rocks, but Early Miocene strata record the influx of GHS detritus. Material with LHS signatures first appears in the Early to Middle Miocene, and both GHS and LHS detritus is present after that time. Within the spatial and temporal resolution of the models, results from HT1 and H T l l l appear to be generally consistent with the Lesser Himalayan provenance data.

Provenance of gneiss domes and leucogranites The north Himalayan gneiss (NHG) domes, which lie in the Tibetan Plateau south of the IndusYarlung Tsangpo suture, are cored by gneisses resembling those exposed in the GHS (e.g. Wu et al. 1998; Zhang et al. 2004). These rocks exhibit condensed metamorphic sequences, are locally cut by leucogranites, and are separated from overlying lower grade rocks of the Tethyan Series by ductile shear zones and brittle normal faults (e.g.

PROVENANCE AND CHANNEL FLOW MODELS

177

Fig. 8. Distribution of Lagrangian points for materials at or near the surface in the vicinity of the orogenic front, model HT111. Initial distribution of materials and tracked P-T-t points (labelled circles) shown in Figure 2. Crustalscale results shown in Figure 6. Panel set (a) shows materials Ib, Ic, Id, Ie, Ih, Ii; panel set (b) shows Ib, Id, Ii; panel set (c) shows Ic, Ie, Ih. Stage 1 'GHS' material distribution at 6 Ma resembles that in HT1 at 0 Ma (Fig. 7), but subsequent dome extrusion deforms stage 1 'GHS', broadens the high-grade region, and juxtaposes the more diverse dome material (stage 2 'GHS') against previously extruded stage 1 'GHS'. Model 'MCT' and 'STD' determined as for Figure 7.

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Lee e t al. 2000, 2004, 2006). Petrological and isotopic studies of Miocene leucogranites in both the Himalaya and the gneiss domes suggest that they were derived by melting of GHS source rocks (e.g. Inger & Harris 1993; Patifio Douce & Harris 1998; Harrison e t al. 1999; Zhang e t al. 2004). Material tracking in models HT1 and H T l l l shows that the model crust south of the suture is made up of 'Indian' material. Although channel flow has transported distal middle crust to the erosion front, the suture and its surrounding 'Tethyan' crust have also been transported southward. Channel flow at depth has not yet carried 'Asian' middle and lower crust more than c. 20 km south of the surface exposure of the suture. Both models therefore predict that leucogranites and gneiss domes located south of the suture should record 'Indian' provenance. Furthermore, continuous transport of material into the orogen from the 'Indian' side of the system provides a steady supply of fertile source material for diachronous leucogranite generation. As noted above, model H T l l l predicts that material within the

NHG domes should resemble the 'GHS', although relatively more distal (shallower) and proximal (deeper) 'Indian' crust is juxtaposed during dome formation (Figs 6 & 8). Model predictions for the region south of the suture are therefore compatible with each other and with observations from leucogranites and gneiss domes south of the suture. The two models make different predictions, however, for plutons and gneisses formed north of the suture. In model HT 1, 'Indian' material occupies the lower crust for 200 km north of the suture, suggesting that granites and gneisses originating at depths greater than c. 30 km should have 'Indian' crustal signatures in this region, although mixed signatures could be produced by assimilation of 'Asian' upper crust. In model H T l l l , however, 'Asian' middle and lower crust is transported southward along with overlying 'Tethyan' material, so that crustally derived plutons and gneisses north of the suture should have dominantly "Asian' characteristics. In this respect, data indicating an Asian (Nyainqentanglha) source for dacite dykes in the Lhasa terrane north of the suture (Harris e t al.

Fig. 9. Pressure-temperature-time (P-T-t) paths for selected 'GHS' points in model HTI (a) and model HTI 11 (b) compared with data relevant to generation of Himalayan leucogranites. Locations of tracked points (G I, G2, D1, D2, D3) at various times shown in Figures 2, 5, 6, 7, 8. Other P-T-t paths from model HTI presented by Jamieson et al. (2004). Alternative AI2SiO5 triple-point positions from Holdaway (1971; H71) and Pattison (1992; P92). Melting ranges and muscovite and biotite dehydration reactions from Peto (1976), Stevens & Clemens (1993), Patifio Douce & Harris (1998) and Gardien et al. (2000). Mineral abbreviations: MS, muscovite; BT, biotite; QZ, quartz; Kfd, K-feldspar; Ab, albite; And, andalusite; Sil, sillimanite; Ky, kyanite. P-T data from Davidson et al. (1997; Bhutan), Prince et al. (2001, Garwhal), Vison~ & Lombardo (2002; Makalu), and Harris et al. (2004; Sikkim). P-T-t paths from both models cross the muscovite dehydration melting field between 30 and 15 Ma, and therefore predict leucogranite generation at the observed times. Isothermal decompression in model HTI 11 (D1, D2) results from dome creation and expulsion; P-T-t paths pass through the andalusite + melt stability field after 12 Ma.

PROVENANCE AND CHANNEL FLOW MODELS 2004a; King et al. 2005) are more consistent with H T l l l than with HT1. A mid-Miocene dyke swarm with similar Asian geochemical affinities (King et aI. 2005) lies within 20 km south of the suture (N. Harris, pets. comm. 2005), near the southern limit of Asian middle crust predicted by models HT1 and HT111 (Figs 5, 6 & 7). In addition to isotopic constraints on source rock characteristics, petrological and geochronological studies of Himalayan leucogranites provide pressure-temperature-time (P-T-t) data that can be compared directly with model P-T-t paths. Many Himalayan leucogranites are interpreted to have formed at 20-25 Ma by dehydration melting of muscovite-bearing GHS protoliths at midcrustal depths (kyanite _+_sillimanite stable; e.g. Inger & Harris 1993; Patifio Douce & Harris 1998; Harrison et al. 1999; Harris et al. 2004b; Godin et al. 2006). Early (c. 40Ma) fluid-present melting has been inferred in some places (e.g. Prince et al. 2001), and both melting and emplacement are locally inferred to have taken place at relatively low pressure (<5 kb; e.g. Vison~ & Lombardo 2002). Plutons significantly older (>30 Ma) and younger (< 15 Ma) than the typical Miocene range have also been reported from various parts of the orogen (e.g. Edwards & Harrison 1997; Wu et al. 1998; Zhang et al. 2004; Godin et al. 2006). In models HT1 and HT111, P-T-t paths from the lower 'GHS' pass through the medium-pressure muscovite dehydration melting field between 30 and 15 Ma (Fig. 9), demonstrating that both models are capable of generating the necessary P-T conditions at the right time. P-T-t paths from the upper 'GHS' (G2, D3) lie mainly within the fluid-present melting field and some (e.g. G1, model HT1) pass through this field before 40 Ma (Fig. 9a). Both models can therefore account for fluid-present melting between c. 45Ma and 6Ma. None of the 'GHS' P-T-t paths from model HT1 pass through the low-pressure andalusite + m e l t field; this model therefore cannot explain andalusite-bearing leucogranites such as the Makalu pluton (Vison?t & Lombardo 2002). In contrast, dome formation and extrusion in model H T l l l leads to isothermal decompression; the resulting 'GHS' P-T-t paths stay in the melting field beyond 15 Ma, crossing into the andalusite + melt field after 12 Ma (Fig. 9b). Models of this type provide an explanation for both low-pressure leucogranites and the generation of post-15 Ma plutons. P-T-t paths from the second (buried) dome in HT111 (Jamieson, unpublished data) indicate that the conditions necessary for medium-pressure leucogranite generation are reached in dome cores by c. 24 Ma. This is consistent with relatively young intrusion ages reported from some NHG domes (e.g. Zhang et al. 2004; Lee et al. 2006).

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Discussion and conclusions It has been argued that the channel-flow model can be refuted by provenance data, including the absence of Lhasa block (Asian) material in the GHS, the lack of a Gangdese arc signature in GHS detrital or xenocrystic zircons, and the continuity of GHS stratigraphy along much of its length (e.g. Harrison 2004, 2006; Harrison & Aikman 2004). These claims are not consistent with the model results presented here. Although channel flow in the models operates for about 30 million years, Asian upper and middle crust are also transported southwards, so that by the end of the models Asian-derived channel material has not advanced significantly south of the surface position of the suture. The models therefore do not predict an Asian signature in the GHS, or in melts or detritus derived from it. On the contrary, model results are entirely compatible with Indian crustal signatures from a variety of metamorphic, igneous and sedimentary rocks in the southern part of the orogen. In the lower model orogenic crust, strongly deformed materials from diverse 'Indian' sources have been interleaved by transposition and folding (Figs 5 & 6). In both models, however, material in the exhumed middle crust, corresponding to the 'GHS' at the model surface, is derived mainly from originally contiguous protoliths that flowed outward together in the upper part of the channel. This material could therefore appear to retain a coherent stratigraphy (Figs 7 & 8). Within the temporal and spatial resolution of the models, we conclude that material tracking results from models HT1 and H T l l l are consistent with provenance data from the Himalaya and southern Tibet. An important implication of the model results presented here and elsewhere (Beaumont et al. 2004; Jamieson et al. 2004) is that channel flow and many other orogenic processes are inherently diachronous, involving substantial lateral transport of both heat and material. Material emerging today at the orogenic front has passed through several different thermal-tectonic regimes during its transit through the orogen; those regimes may persist today in the interior of the orogen. For example, dome formation and extrusion in model H T l l l is continuous - as one dome is extruded at the orogenic front, another forms above the mid-crustal ramp. Convergence in the Himalayan-Tibetan system is constantly bringing new material into the orogen. Incoming material may eventually become involved in channel flow and re-emerge at the orogenic front, but most of the material that enters the orogen remains deeply buried. The models suggest that the distinctive GHS and LHS provenance signatures observed today do not represent the composition of the

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orogenic crust as a whole, but only that portion that has recently reached the surface. Models HT1 and HT111 were chosen to illustrate contrasting tectonic styles in the region between the I n d u s - Y a r l u n g Tsangpo suture and the Himalayan front. Differences in tectonic style, including the width of the orogen, generation and extrusion of domes, and the behaviour of 'Asian' crust, are produced by a small difference in the mechanical properties of model 'Tethyan' crust. Previous work (Beaumont et al. 2004; Jamieson et al. 2004) showed that model HT1 is consistent with metamorphic and tectonic data from the central part of the orogen, and the present study shows that it is also compatible with a range of provenance data. Model H T l l 1, although not yet subjected to the same degree of scrutiny, is equally compatible with the provenance data summarized here. In addition, H T l l l can explain some observations (e.g. NHG domes) not accounted for by HT1. We interpret models HT1 and H T l l l to represent members in a spectrum of channel flow model styles that could be applied to different parts of the Himalayan-Tibetan system and/or to different stages in its evolution. Furthermore, the homogeneous channel flow mode represented by models HT1 and HT111 is only one of several possible flow modes in large hot orogens (Beaumont et al. 2006). It is conceivable that other flow modes (heterogeneous channel flow, hot fold nappes) could have operated in some parts of the Himalayan-Tibetan system at various times. We conclude that channel flow models in general provide a reasonable first-order explanation for the thermal-tectonic and lithological evolution of the Himalaya and southern Tibet. However, more than one model style may be compatible with the observations and no single model is capable of explaining all the features of the orogen. As noted by Beaumont et al. (2006, appendix), the models lack the spatial-temporal resolution required for transect-specific comparisons, and are sensitive to subtle variations in input parameters that are already greatly simplified by comparison with nature. Observations that can be integrated on a regional, crustal or lithosphefic scale, including provenance data from such diverse parts of the orogen as the Ganges basin and the High Himalayan leucogranites, provide important tests of first-order model predictions. This work was funded by NSERC Discovery to R. A. J., C. B. and D. G. C. B. acknowledges support from the Canada Research Chairs programme and an IBM Shared University Research grant, and D. G. from the Canadian Institute for Advanced Research. The ALE numerical model was developed by P. Fullsack, Dalhousie University, and material tracking was implemented with the assistance of S. Medvedev, Freie Universit~it, Berlin. The work has

benefited from discussions with, and constructive criticism by, a number of Himalayan-Tibetan researchers, including P. DeCelles, N. Harris, K. Hodges, L. Hollister, Y. Najman, and especially the late D. Nelson. The paper was improved by constructive and thorough reviews by R. Parrish, P. DeCelles, and the balanced advice of editor R. Law. We are grateful to the conference organizers for providing a stimulating environment in which to present and discuss this work, and to D. Nelson for originally encouraging us to tackle this problem.

References

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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 Thust in Nepal. Geology, 31, 359-362. 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. STEVENS, G. & CLEMENS, J. D. 1993. Fluid-absent melting and the roles of fluids in the lithosphere: A slanted summary? Chemical Geology, 108, 1-17. UNSWORTH, M. J., JONES, A. G., WEI, W., MARQUIS, G., GOKARN, S. G., SPRATT, J. E. & the INDEPTH-MT team. 2005. Crustal rheology of the Himalaya and Southem Tibet inferred from magnetotelluric data. Nature, 438, 78-81. VAN DER MOLEN, I. & PATERSON, M. S. 1979. Experimental deformation of partially-melted granite. Contributions to Mineralogy and Petrology, 70, 299-318. VISON.~, D. & LOMBARDO, B. 2002. Two-mica and tourmaline leucogranites from the EverestMakalu region (Nepal-Tibet). Himalayan leucogranite genesis by isobaric heating? Lithos, 62, 125-150. WU, C., NELSON, K. D., WORTMAN, G. ET AL. 1998. Yadong cross structure and the South Tibetan Detachment in the east central Himalaya (89~176 Tectonics, 17, 28-45. ZHANG, H., HARRIS, N., PARRISH, R. ET 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.

Kinematic dilatancy effects on orogenic extrusion B. G R A S E M A N N , M. A. E D W A R D S & G. W I E S M A Y R Structural Processes Group, D e p a r t m e n t o f Geological Sciences, University o f Vienna, Althanstrasse 14, Vienna A-1090, Austria (e-mail: Bernhard. Grasemann @ univie.ac.at) Abstract: We undertake kinematic modelling to explore the role of volume increase in a slab

extruding from an orogenic wedge with constant or decreasing slab width. Using a dilatancy term, we modify the velocity gradient tensor dependent on the stretching-rate factor, kinematic dilatancy and vorticity number. We use this to explore the previously largely ignored role of volume change in kinematic evolution of extrusive flow, considering area change for non-isochoric flow types with no deformation in the intermediate direction. By keeping individual parameters constant for geologically simple scenarios (e.g. finite strain, steady-state flow) we examine the interdependence of the reciprocal parameters (kinematic vorticity and dilatancy number) and note model situations where degrees of freedom are limited. These interdependent parameters thereby provide a set of rules for integrating and modelling real field data. In particular we observe that for extrusion flow with a constant slab (or 'channel') width, degrees of freedom in kinematic vorticity and volume change at given finite strains are very restricted. We compare scenarios of low and high strain and low and high volume change on anatexis (related to partial melting of fertile sedimentary rocks and release of water upon crystallization) for different parts of the Himalaya.

With the identification of synconvergence extension in active orogenesis (e.g. Burg 1983; Burchfiel & Royden 1985; Ratschbacher et al. 1991; Wallis et al. 1993; Edwards et al. 1996; Grasemann & Vannay 1999; Wiesmayr & Grasemann 2002), it has become apparent that in many large-scale, wedge-geometry accretionary or collisional orogens (for example, the Himalaya), metamorphic rocks are exhumed by transport in a subhorizontal direction towards the surface within a discrete, frequently gently dipping slab, bounded above and below by major crustal discontinuities that act as normal and thrust faults respectively (e.g. Hodges et al. 1992). Continuous underplating at depth (e.g. along the Himalayan Main Central Thrust, MCT) is reciprocally(?) accompanied by relative normal-sense displacement in the upper part of the wedge (e.g. along the Himalayan Southern Tibet Detachment System, STDS) resulting in exhumation of an orogenic wedge (e.g. the Higher Himalayan Crystalline, HHC). This architecture geometrically and kinematically describes an extruding layer or slab of the middle crust (e.g. Royden & Burchfiel 1987; Henry et al. 1997; Wu et al. 1998; Beaumont et al. 2002). Note that we use the term 'slab' to describe this extruding mid-crustal layer as it best suits the first-order geometry. This is distinct from, and is an internal portion of, the greater orogenic wedge (a region delineated by converging plate-type boundaries, which we do not consider here). Originally discussed as a mechanism for the exhumation of eclogites, such models describing

extrusion of metamorphic rocks above a thrust and below a co-genetic normal fault have included both 'channel flow' and otherwise general extrusion types (e.g. England & Holland 1979; Mancktelow 1995). We remark that the term 'channel flow' is notably ill-defined in the physics literature and may range from laminar, Couette to Poiseuille flow through a pipe, to frequently even include rather complex open-, turbulent- or vorticalchannel flow types (e.g. Munson et al. 2002). Each discrete flow type has its own strict and exact boundary conditions, none of which are readily suitable (e.g. inappropriate Reynolds number values) for describing the exhumation of a deforming rock slab between shear zones with different vorticity (cf. Grujic 2006). Even if a clear mechanical definition for channel flow of rocks (i.e. the middle crust) could be introduced into the mainstream geological nomenclature, it is very rare that sufficient geological (field, laboratory) constraints are ever obtainable to uniquely identify such a specific flow type. We therefore strongly suggest that the term 'channel flow' is used only in investigations associated with very clearly defined flow scenarios in models or in nature (e.g. England & Holland 1979; Mancktelow 1995; Royden 1996; Clark & Royden 2000; Beaumont et al. 2002, 2004; Jamieson et aL 2004). Where otherwise generally 'ductilely' (i.e. flow-) deformed mid/lower crustal rocks are observed in exhumed young/fossil orogens and flowaccommodated exhumation is inferred, we suggest

From: LAW, R. D., SEARLE,M. P. & GODtN, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 183-199. 0305-8719/06/$15.00 9 The Geological Society of London 2006.

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that the non-specific flow mechanism term 'extrusive flow' be used instead. Intimately associated with the exhumation of metamorphic rocks are anatectic processes of considerable magnitude frequently involving (for many orogens) the progressive dry melting of fertile pelitic (source) rocks (e.g. Vielzeuf & Holloway 1988) deeper in the orogenic wedge via decomposition of muscovite (and occasionally biotite), in association with extensive melt production, extraction, migration and coalescence (e.g. Deniel et al. 1987; Le Fort et al. 1987; Inger & Harris 1993; Harris & Massey 1994; Scaillet et al. 1996). In the Himalaya, for example, this process is expressed by: (1) abundant stromatic migmatite in sillimanite gneisses throughout the upper portion of the crystalline slab, as well as (2) numerous decametre-thick sheets and kilometre-thick tabular plutons (the High Himalayan Leucogranites) that underlie many of the high peaks, forming a belt that extends along the > 2 5 0 0 k m orogenic arc (e.g. Gansser 1964, 1983; Le Fort 1975; Scaillet et al. 1995; Edwards et al. 1996; Schneider et al. 1999; Wiesmayr et al. 2002). Both the High Himalayan Leucogranite belt and a second prominent belt located 5 0 - 1 0 0 k m to the north housing much smaller (Watts et al. 2005) leucogranite bodies (the Northern Himalayan Granites) have remarkable chemical and geochronological consistency along the arc (e.g. Edwards & Harrison 1997; Harrison et al. 1998, and references therein; Whittington & Treloar 2002). Melt production in the Himalaya was formerly regarded as a short discrete event (Debon et al. 1985, 1986; Harrison et al. 1998), or several events (Guillot & Le Fort 1995), that created the High Himalayan Leucogranites alone, and/or separately, the Northern Himalayan Granites. More recently, however, with constraints upon the geometry of Himalayan orogen provided by controlled-source CMP reflectors and wideangle velocity structure (Zhao et al. 1993; Makovsky et al. 1996; Nelson et al. 1996; Hauck et al. 1998), it has become accepted that convergence, exhumation and extrusion proceed in an overall steady-state manner and are directly linked with melting and plutonism (e.g. Wu et al. 1995; Henry et al. 1997; Beaumont et al. 2002). Surprisingly, however, despite the massive dimensions of reordering of material within the orogen as a result of widespread melt transfer and fluid expulsion upon crystallization, none of the popular models for Himalayan collision (or other collisional orogenic settings) incorporate any change in volume. This includes models involving either coeval underthrusting with extrusion (i.e. no regional addition of material, only slab boundaries narrowing) and/or true extrusion (i.e. adding material to the overall slab whereby some material

must extrude from the front of the orogen). In this contribution, we explore volume-change scenarios in a series of simple plane-strain kinematic models for extrusive flow incorporating area increase. Because plane-strain implies no change in the intermediate direction, all volume change is accommodated by area change in two dimensions. Nonisochoric flow types (which should be associated even with small volume (or area) changes) are likely to have a dramatic effect on the finite stretch and rotation of material lines (particularly the rate of change of material line orientations and lengths) as has previously been suggested (Ramsay & Graham 1970; Schwerdtner 1982; Passchier 1988; Fossen & Tikoff 1993). We focus on controlling parameters like finite-area change, finite deformation or kinematic vorticity number. These are parameters that can, under favourable circumstances, be quantified or derived from natural examples (O'Hara 1988; Vissers 1989; Wallis 1992; Srivastava et al. 1995; Hippertt 1998; Grasemann et al. 1999; Ring 1999; Xypolias & Koukouvelas 2001; Law et al. 2004). More rigorous treatments of this topic, based on continuum mechanical considerations, are given for example in Schwerdtner (1982), Passchier (1988) and Simpson & De Paor (1993) and references cited in these publications.

Modelling approach Thermal and/or mechanical models are often highly complex, incorporating a large number of parameters with great freedom of boundary conditions and input values, thereby providing great flexibility in model outcome. Although such models are a fascinating tool for examining nature and provoking new ideas, the wide range of parameters employed in such models impedes the ability to assess the importance of individual variables (i.e. which significant geological phenomena are operating and to what degree). We employ here instead a simple kinematic model as a framework and calculate how the values of chosen parameters vary. This allows us to explore the limits of freedom in boundary conditions and the interdependence amongst the various parameters. Such a kinematic modelling approach has a significant advantage in immediate prediction of the fate of material lines undergoing deformation. Moreover, the approach is practical in that most of what we measure (e.g. rock fabrics, palaeoisotherm, volume change on melting) may be specifically related to the behaviour of material lines. In many cases, such models can be tested in the field, for example by measuring the orientation distribution of feeder dykes under High Himalayan plutons. We consider a two-dimensional, homogeneously deforming unit cell to represent the cross-sectional

DILATANCY EFFECTS ON EXTRUSION area of the extruding and deforming orogenic slab (Grasemann et al. 1999; Vannay & Grasemann 2001). The upper and lower boundaries are normal and thrust faults, respectively, whose displacement decreases to zero to the left (Fig. 1), which, in the case of the Himalaya, is to the north, and nominally at the suture (the northern limit of displacement is uncertain). More complex extrusion geometries can be set up by combining such cells with different deformation characteristics, while simultaneously considering requirements of strain compatibility (Simpson & De Paor 1993). There are clear caveats in our approach: we ignore heterogeneous deformation, strain partitioning and mechanical effects, as well as periods of non-steady-state straining. This latter caveat is probably our greatest Achilles' heel: if we invoke a 10% volume change, or switch on or off any of the other boundary conditions it is highly unlikely that steady-state deformation can persist (see discussion below). These assumptions are necessary, however, in order to investigate the purely first-order effects of dilatant flow. Moreover we emphasize that in many mountain belts (the HHC of the Himalaya, for example) the first-order assumption of overall homogeneity in the extruded slab is borne out by the appearance of the rocks; although they have flowed, the slab has maintained

O)

Lo =

/e:

1(122,

185

a coherent tectonostratigraphy. Massively partitioned, turbulent or other complex-directional flow, that is usually indicated (e.g. by superposed folding) in many multiply deformed ancient shield areas that expose deep continental crust, is strikingly absent in the extruded Himalayan middle crust.

Continuum mechanics model I s o c h o r i c p l a n e strain The velocity gradient tensor processes the position vector of any particle at an instant time and generates the associated velocity vector for that particle at that given instant. Homogeneous flow types can be conveniently modelled by the velocity gradient tensor; the instantaneous velocity of particles at X with the Eulerian rate of displacement is:

Sij = LijXj

(1)

where L is the velocity gradient tensor. The column vectors of L can be used to construct a Mohr circle (De Paor & Means 1984) by plotting the angular velocities o) of material lines against stretching rate g (Fig. 1). The angle subtended at the centre of the Mohr circle by two points on its perimeter is double the angle between two material lines in real space (Mohr 1882). The coordinates of the centre of the Molar circle represent half of the instantaneous area change (A) and half of the vorticity (W) (Passchier 1991). In isochoric flow types A = 0 and therefore the Mohr circle is centred on the o~-axis. The intersections of the Mohr circle with the e-axis represent the eigenvectors of L (a~,a2), which are lines of no instantaneous angular velocity (Bobyarchick 1986). The cosine of the angle between the eigenvectors in real space gives the kinematic vorticity number Wk, which is here defined after Passchier (1987): W cos oz = Wk = - S

Fig. 1. Mohr circle of general velocity gradient tensor L (note convention of stretching rate ~ and angular velocity ~ogeometric space). Shown are relationships between instantaneous stretching axes (ISA), eigenvectors al,2 (lines of no instantaneous angular velocity) and angles oz and/3 whose cosines give kinematic vorticity (Wk) and kinematic dilatancy (Ak), respectively. Note that the Mohr circle chosen is for nonisochoric flow as circle is not symmetrical on o) axis. See text for further clarification.

(2)

where S is the stretching rate factor (the difference between the maximum and minimum instantaneous stretching rate). Wk quantifies the flow type between the end-members of simple shear (Wk = 1) and pure shear (Wk = 0). Note that because of the cosine function in Equation 2, we linearize Wk by transforming it into 'percentage simple shear', for which we propose the notation Ye, by:

(

yp = 100 1

cos

7r/2

]

(3)

We want to compare finite structures that develop for a range of different extrusion flow types that (in

B. GRASEMANN ET AL.

186

almost all cases) will be between the pure- and simple-shear end-members. It is therefore critical to find components of L that describe finite deformations that are comparable amongst the different flow types. However, there is no general agreement over what boundary conditions will lead to equivalent finite deformation when considering different flow geometries (e.g. Fossen & Tikoff 1997; Jiang 1998). Here, we use velocities for different flow geometries that are characterized by the same stretching-rate factor S over the same time increment (Grasemann et al. 2003). For a given Wk and S, the components of L can be calculated by (corrected after Passchier 1987, fig. 2):

/

stretch in the third dimension); for dilatant plane strain, the definition of L therefore has to be extended by the kinematic dilatancy number A/, (Passchier 1988, 1991). By analogy with Wk, Ak can be expressed by/3, which is half the angle subtended at the centre of the Mohr circle by the points of interception with the w-axis (points for which stretching rates is zero) and is here defined as (Fig. 1): cos fl =

A S

Ak = -

(7)

where A/2 = &4, which is the rate of change in area (equivalent to a in Passchier 1988). Including Ak, L can be defined as:

(4a)

o

Lq= 0

which can be simplified to:

SA (8a)

Lij =

-Ae

which simplifies to:

where Ae is half the difference between the eigenvalues. L can be easily converted into the finite Lagrangian deformation tensor D, given that 0 _< W~ < 1 (Provost et al. 2004):

Dij = exp(Lgjt)

(5)

which is given over a time increment (t) of l (compare Ramberg 1975, equation 38) by: 1

V/1 - W

D6=

2

Lij = ( AA +

AA-

)

which can be converted into the finite deformation tensor D over a time increment of 1, given that 0 < Wk < 1, using the method described by Provost et al. (2004): O.~

exp(S(Ak-- l--~Z-~k2)) ]

exp(S(Ak+

l~-W~)) (exp(S l~-W~)-l)w, I

,/,_w:

2

0

(%) (6a)

which simplifies to:

Dij =

exp(Ae) +exp(--Ae)~ ~) tana exp(-Ae) /

I

e x p ( S ( A , - 1-~TZ~)) ]

e x p ( - ~ l ~ - W k 2)

(exP(0A

(8b)

which simplifies to: {exp(&4 +Ae) exp(AA +Ae-- 1) ] tana / \ 0 exp(AA--Ae) /

Dij=| (6b)

Dilatant plane strain We wish to model extrusion flow which can incorporate volume change, and so must include a change in area assuming plane strain (i.e. no

(9b)

We emphasize that Ak has to be converted into percentage area change ~ p for a given finite deformation tensor D, which gives over a time increment of 1:

AAp=D11D22=lOO(exp(AkS)-l)

(10)

DILATANCY EFFECTS ON EXTRUSION

187

Fig. 2. Examples of finite deformations that include finite-area increase AAp = 50% (a-d) show: on left, geographic space depicting fates of an initial unit square after (a, b) simple shear or (c, d) isochoric general shear (shown in dark grey) and after (a, c) area change or (b, d) non-isochoric distortion (shown in light grey); on fight, the Mohr circle for the velocity gradient tensor L given for a stretching rate factor S -- 1 corresponding to the finite deformation integrated over a time period of 1. (a) Simple shear (Wk = 1) plus area change results in a proportional stretching and widening of the sheared area, which appears in the Molar space as a horizontal shift of the Molar circle to the fight. (b) Non-isochoric distortion with a fixed width of the sheared area appears in the Mohr space as a rotation of the Mohr circle around the eigenvector resulting in two intersections of the axis of zero instantaneous angular velocities (Wk = 0.915). (c) General shear (Wk = 0.9) plus area change results in a proportional stretching and widening of the sheared area, which appears in the Mohr space as a horizontal shift of the Mohr circle to the right. (d) Non-isochoric distortion with a reduction of the width of the sheared area appears in the Molar space as a rotation of the Mohr circle around the second eigenvector resulting in a markedly lower kinematic vorticity number (Wk = 0.55).

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Therefore the AAp is dependent only on Ak and S, which leads to the conclusion that if Wk and Ak are constant during deformation (i.e. steady-state flow), the volume will increase exponentially.

Finite h o m o g e n e o u s d e f o r m a t i o n wi th area i n c r e a s e An area increase during homogeneous deformation can be achieved by: (1) a positive elongation in both the x- and y-direction; (2) a positive elongation in either the x- or y-direction; or (3) a combination of positive elongation and negative elongation in the x- and y-direction, as long as the product of the diagonal terms in our tensor D is greater than 1. In the Mohr diagram for L, introduction of an area increase leads to a shift of the circle to the right (fig. 1; Passchier 1988, 1991). However, if the flow is a combination of distortion and area change (i.e. non-isochoric distortion) this shift to the right on the Molar diagram will not be restricted to the horizontal and consequently Wk will change. In order to model extrusive flow (whereby material must displace and lengthen in unequal amounts) we consider here non-isochoric distortion, where the thickness during deformation either remains constant (Fig. 2b) or decreases (Fig. 2d), both of which flow types modify the W/, of the corresponding isochoric flow (i.e. isochoric flow with the same thickness after deformation). Interestingly, non-isochoric distortion that retains a constant thickness can be revealed by rotation operation of the Mohr circle of L for the corresponding isochoric flow. In this operation, the centre of the Mohr circle of L for the corresponding isochoric flow is rotated about the point that corresponds to the second eigenvector of the flow resulting in an increase and decrease of Ak and Wk respectively (as achieved in Fig. 2b & d). If W1, of the corresponding isochoric flow and the non-isochoric distortion are given by Wk~ and Wk2, the angle of rotation is: q~ = cos -1 Wk2 -- cos -1 Wkl

(11)

As pointed out by Passchier (1991), flow types with area change are ambiguously classified in terms of simple and pure shear. For example, if simple shear is defined by a flow type having just one direction which is neither rotating nor stretching, then simple shear with area change cannot be classified as simple shear because its non-rotating direction is stretching (Fig. 2a). Similarly, non-isochoric distortion retaining constant thickness has two non-rotating directions, which in classical terminology are also associated with general shear. If such flows have a fixed width

(i.e. the extruding slab has constant thickness during deformation), one non-rotating direction is stretching and parallel to the slab boundaries. The other non-rotating direction is neither stretching nor shortening (Fig. 2b and 4, discussed below), a characteristic normally considered as typical of simple shear (Passchier 1991). It is not the intention of this paper to introduce new terminology to the continuum mechanics geology community, and we follow Passchier (1991) in keeping the terms pure and simple shear, noting these ambiguities in such non-isochoric flow conditions. Figure 3 shows contour plots for components of the finite deformation tensor for a range of deformations where the variable parameters are finitearea change (AAp, plotted on the x-axis) and the contribution of the simple-shear component (yp, plotted on the y-axis). As noted, we use percentage area change AAp and percentage simple shear ye, because Equations 2 and 7 are both non-linear. Our new deformation tensor (Equation 9a) is sensitive to the stretching rate factor S, as well as Wk and Ak. In order, therefore, to make the different flow types comparable, the range of finite deformations has been calculated at the same stretching rate factors for two separate scenarios: a 'low-strain' (Fig. 3, left column) and a 'high-strain' (Fig. 3, right column) case, where we set S = 1 and 3, respectively. As a starting example to assist the reading of our new diagram, we use the deformation in physical space of an initial unit square (as shown in the upper row of Fig. 3) where AAp = 20% and yp = 80% for our 'low' and for our 'high' strain scenarios. The unique position of this deformation in all the contour plots is indicated by a star. To allow us to follow the change in shape of our extruding slab, we show the contours for the normal-strain components of D (i.e. the pure-shear components) in the upper pair of contour plots in Figure 3, which give: (1) the stretch parallel to the slab, or more specifically, the stretch in the extrusion direction (x, shown by the unbroken lines); and (2) the stretch normal to the slab (y, the broken lines). The field that we suppose to be realistic for an extruding slab is shaded in grey for both the 'low' and the 'high' strain scenarios. This field delimits the y-direction stretch (i.e. changes in length normal to the slab, i.e. 'channel') to an upper boundary of stretch in y - - 1.0 and a lower boundary of stretch in y = 0.6. These respectively correspond to end-member scenarios where the slab does not change its width during deformation (stretch in y = 1.0) and where the slab thins/flattens with a maximum of 40% (stretch in y = 0.6). We do not delimit the stretch in x but the values that are present in our stretch-in-y-delimited field are not unreasonable (stretch in x < 2.4 for S = 1 and <3.5 S = 3). Note that the range of stretch-in-x

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Fig. 3. Contour diagrams for components of the finite-deformation tensor D graphed for AAe (percentage finitearea change) against "re (percentage contribution of the simple-shear component) at a constant stretching factor S for two scenarios. Two geologically relevant end-member scenarios are chosen: low (S = 1) and high (S = 3) strain (left and right columns, respectively), illustrated in uppermost (y versus x) plots. These show geographic space deformation of initial unit square with AAp = 20% and "re = 80% (these values are shown as stars in contour diagrams below). Upper pair of contour diagrams shows the normal-strain components of D: stretch in x and stretch in y, represented by solid and dashed lines, respectively. Grey field is our suggested limit of geologically realistic stretch values for an extruding slab (note sensitivities to area change at low "re, especially for S = 1). Lower pair of contour diagrams shows the effective shear strain component of D (note insensitivity to area change at low "rp). See main text for discussion. In all diagrams, stretch values are shown only for every second contour to limit clutter on diagrams.

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values, and therefore the intervals of our selected contours, is much broader in the high-strain scenario. One interesting observation, clearly apparent in the contour diagrams, is that for low strain (lefthand column of Fig. 3) where flow is pure-sheardominated (e.g. area of the diagram where yp < 20%), the normal strains of the finite deformation are mainly controlled by the area change (i.e. the contours for both the stretch in x and stretch in y are steep and are very sensitive to zL4p). However, for high strain (right-hand column of Fig. 3), the gradients of the contours are less steep and the normal-strain components of nearly all flow types, except those of strongly pure-sheardominated flows (i.e. where ye < 10%), are more sensitive to a change in the simple shear component. One way to examine this switching (or strong variation) in inter-parameter sensitivity is to fix one value and observe the variation in the other variables. For example, consider the case of 'channel flow' where there is no slab thickness change (i.e. no change in the y-direction). If different finite deformations, along the broken line stretch in y = 1 are compared (Fig. 3), an area increase is associated with an increase in the extrusion component (stretch in x), but a decrease in the simple-shear component. This effect is less pronounced at higher strains (compare shaded areas in upper pair of diagrams in Fig. 3), although still present. In the lower pair of contour plots in Figure 3, we show the effective shear-strain components (Tikoff & Fossen 1993) of D again for our 'low' and 'high' strain cases. It is obvious that the effective shearstrain component in pure-shear-dominated finite deformations (Aye < 50%) are insensitive to a change in area and are mainly a function of yp. This insensitivity is even greater at higher strains. Again, note that the intervals of our selected contours are broader in the high-strain scenario. This leads to the conclusion that the relative offset between the hanging wall and the footwall of the slab (i.e. the effective shear-strain component) is essentially not influenced by an area increase within the slab, given that deformation is pure-shear-dominated. Comparing simple-shear-dominated finite deformations (for example, comparing yp > 80% in the 'low' strain plot versus the 'high' strain plot), the effective shear-strain component is practically insensitive to a change in Ye but highly sensitive to an area change. It is highly counterintuitive, in fact, that an area (i.e. volume) increase within the slab has a major influence on the relative offset between hanging wall and footwall, if the overall flow geometry is simpleshear-dominated (Ye > 70%). We emphasize that in addition to their use in identifying which lesser-constrained parameters have the most/least degrees of freedom for subsequent interpretation/modelling, these plots can also be

used to test existing models. For example, it is impossible for an extruding slab that is experiencing thinning of its width at a low finite deformation to be accompanied by a high percentage of area change as well as a dominant simple-shear component! However, if the finite strain is high, even slabs which have been thinned can suffer a significant shear deformation, where the pure-shear components are nearly insensitive but the simple-shear components are highly sensitive to a change in area.

Homogeneous deformation with constant thickness Figure 4a shows contours for functions in the AAp and YP diagram for a range of different stretching rates S, for finite deformation tensors, where D22 = 1 and therefore (compare Equation 9a): W k = v~l - A 2

(12)

This holding constant of the Dee component at a value of 1 means that there is no shortening in the y-direction and therefore the material deforms in a slab with fixed width. Equation 12 means that the W1, is a direct function of A~ and is insensitive to S (which integrated over time gives the finite deformation). Note, however, that S is introduced again, when Ak is converted into zL4e (compare Equation 10). As can be seen, this specific range of flow types is geologically remarkable because they all have a specific orientation, along which material lines are neither stretching nor rotating, an attribute that, under isochoric flow, is characteristic for simple shear (Passchier 1991). For all the flow types plotted in Figure 4a, however, this orientation is not parallel to the shear zone boundary (in this case the shear zone boundary is stretching) but is in fact dipping into the shear direction at angles 0 < a < 90 degrees (this can be read off on the right-hand side of Fig. 4a). Note again that, as in the previous flow cases contoured in Figure 3, finite strains for pure-shear-dominated flows are sensitive to the area change, while for simple-shear-dominated flows finite strains are nearly insensitive to area change, being instead controlled by the contribution of the simple-shear component. An example of one such finite deformation (at constant thickness, D22 = 1) of a unit square in geographic space is given in Figure 4b. The two thick lines are the two orientations that are nonrotating. Note that the orientation parallel to the shear-zone boundary is stretched by 50% but that the line dipping at 45 ~ into the shear direction has not changed its length. This non-rotating and

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Fig. 4. Diagrams to illustrate effects of homogeneous deformation in a slab with fixed width (i.e. where D22 = 1). (a) Contour diagrams for a suite of finite-deformation tensors D at different stretching rate factors S, graphed for zL4p (percentage finite-area change) against TP (percentage contribution of the simple-shear component). Angle a ( 0 - 9 0 ~) is angle of dip into the shear direction of unique orientation of material lines neither stretching nor rotating (i.e. eigenvector). (b) Geographic space of initial unit square before and after deformation. Thick lines are orientations of the eigenvectors. Note that shear-zone boundary-parallel orientation (horizontal thick line) has not rotated but has stretched; orientation with a = 45 ~ has neither stretched nor rotated. (c) Mohr circle diagram for velocity gradient tensor L corresponding to deformation D shown in (b), showing simple relationship between key angles of interest.

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non-stretching orientation plots in the Mohr circle for L at the origin of the Cartesian coordinate system (Fig. 4c). The non-rotating orientations plot on the x-axis (i.e. the stretching rate axis). Typical for deformations with fixed slab width is that the second, non-rotating orientation (a2) also experiences no incremental stretch. It is furthermore typical that the sum of 2 a and 2/3 always equals 180 ~ which is a direct consequence of Equation 12. Because there is no shortening normal to the shear-zone boundary, the extrusion parallel to the 'channel' is merely due to the addition of material during deformation. As a result of this, for comparable finite strains (i.e. finite strains with similar stretching rate factors S), pure-shear-dominated flow in a channel consumes much more area increase than simpleshear-dominated flow in a channel. This is quite a specific deformation setting that is testable in the field and potentially recognizable by identification of suitably oriented material lines (e.g. sampling lines, dykes etc.) which are not rotated and not stretched during progressive extrusion.

Progressive deformation with constant thickness Figure 5 shows examples for area increase with progressive steady-state deformation over five increments. We investigate several flow types involving a constant slab thickness (D22 = 1) and a constant stretching-rate factor (S = 0.1) for a range of deformations with simple-shear contributions varying between yp = 0% (i.e. pure shear) and y p = 95% (highly simple-shear-dominated). Because of the exponential relationship in Equation 10, the finite-area increase accumulates in an exponential manner, as can be seen. Interestingly however, for similar finite deformations, simpleshear-dominated flow types (e.g. yp > 80%) show a much lower area increase than pure-sheardominated flows. This may be intuitive in view of the fact that, because of the fixed slab width, a pure-shear-type flow can only increase its finite strain by area increase, i.e. by extrusion of its material through the 'channel' (i.e. more material must be added to deform more). As an example,

Fig. 5. Examples of deformation at fixed stretching-rate factor, S = 0.1 in a slab of fixed width (i.e. D 2 2 = 1) with progressive incremental deformation steps (1 to 5) illustrating amounts of area increase for different percentage contribution of the simple-shear component (yp). The two end-member illustrations show that for comparable finitestrain ratios (R = 5, R = 4.3), a huge area increase (400%) is required for pure shear-only (yp = 0) deformation, all of which is expressed as extrusion (stretch in x = 5). See text for further discussion.

DILATANCY EFFECTS ON EXTRUSION the curve in Figure 5 for yp = 0% results in a finite strain ellipse with R = 5 (R is the ratio of the long and short axes of the finite-strain ellipse) after five increments experiencing simultaneously a shearzone-parallel 'extrusion' stretch of 5 and a finite-area increase of nearly 400%. A similar finite deformation (R = 4.3) results from a 95% simple-sheardominated flow. However, in this case, because of the shear-zone-parallel simple-shear strain of 1.7, the extrusion stretch and area increase are just 1.1 and 14%, respectively.

Discussion We recognize that not all geologists working in orogenic belts may choose to critically read through the above steps in our flow kinematics. This volume on 'channel' flow is, however, acknowledgment of the striking growth in interest in this topic throughout the geological community and we thereby undertake to beneficially focus discussion and suggestions in this vein. Embracing the positive growth in interest of 'channel flow' while recalling our above-noted cautions concerning a strict definition of the term, we illustrate what contributions different fieldwork disciplines can make in assessing exactly what flow types have operated, or are operating during extrusion of an orogen-scale metanaorphic slab. Recognizing, as also do the mid-crustal extrusion/channel flow models (e.g. Wu et al. 1998; Beaumont et al. 2002), that the Himalaya does indeed represent the proverbial 'natural laboratory', we base our discussion in this setting. The Himalayan orogen benefits from an excellent set of data constraints on orogenic depths, temperatures, and strain rates, types and magnitudes. All of this, together with an internal coherence and cylindrical geometry that give a notable continuity for more than 1000 km along orogenic strike, affords a dependable foundation for a 2D (plane-strain) consideration of flow. We reiterate that what is presented here is not a geological model, but a set of observations on the interdependence of attendant kinematic parameters during flow (allowing for the above-identified first-order assumptions of homogeneity and steady-state behaviour). These observations then provide a set of hypothetical rules that can allow a more specific identification of the overall flow type, using as input the basic data that are already at hand or can be readily collected. We discuss the data that can be collected in (or assessed from) the field, and what their values mean in terms of the observed trends in values of parameters calculated from our new deformation tensor and then critically consider published Himalayan field data on a case study basis. We

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consider the following phenomena: volume (area) change, kinematic vorticity numbers, and thinning of the slab. Volume change

Two simplified scenarios, Case 1 and Case 2, for volume increase related to melting can be envisioned for the extruding slab (Fig. 6). Case 1 is immediate in-situ melt generation (e.g. Searle et al. 1999) and/or minor melt ascent within the extruding slab, whereby all or most of the decompression melting (i.e. protracted predominantly decomposition melting of the HHC; cf. Whittington & Treloar 2002) and crystallization with cooling occur as part of the extrusion travel path to the topographic surface, i.e. parallel to and between the northward-dipping slab boundaries. This is schematically illustrated in Figure 6 as the lower of the two question-marked and dashed lines (labelled Case 1) that represent possible locations for a northward continuation of the MCT (or Main Himalayan Thrust, MHT; Zhao et al. 1993). In Case 1, there is no major addition of material to the finite area of deformation under consideration; the only volume change is the (relatively minor) volume increase due to phase changes on melting and residual water release upon crystallization (although the effect of this may still be significant if more than a few per cent of the slab area are under consideration - see below). In Case 2, the region of partial melting is well below the main extruding middle crustal layer of the MHT, and possibly even northwards of the MHT fault tip (labelled tip in Fig. 6), the point where MHT thrust displacement diminishes to zero. To date, no melting has been reported within or below the exposed MCT. This case is schematically illustrated as the upper dashed line in Figure 6 (labelled Case 2). It is located within the partialmelting zone (or alternatively the MHT may be restricted to the south, and so represented by only the solid line). In Case 2, all melt product that is seen today as the High Himalayan Leucogranite has ascended several tens of kilometres (Guillot et al. 1995a, b; Patifio-Douce & Harris 1998; Daniel et al. 2003; although see Scaillet & Searle 2006) into the material within the extruding region, thereby affording a massive positive volume change in the extruding region. Differentiating between these two cases is problematic and discrimination is probably limited to subsurface (geophysical) imaging (cf. Klemperer 2006), excluding direct observation in sufficiently well-exposed fossil orogens. Assessing volume change values, however, is still a clear target for fieldwork. In addition to mapping and 3D granite body interpretations, the spatial and temporal

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Fig. 6. Digitized and modified original unpublished sketch by D. Nelson (c. 1995; cf. fig. 6 of Wu et al. 1998). Added here are this study's proposed key alternatives (Case 1 and Case 2) in melt-production mechanism for a general middlecrust extrusion setting. Schematically indicated is steady-state, stepwise emplacement and crystallization of plutons with progressively intensifying general shear (Wk ~ 0.8) engendered via ascending, coalescing melt from shearingextruding anatectic region in middle crust. Amount of volume increase (both finite and incremental) in extruding slab can be large only if partially molten material can ascend into and supplement material in extruding region, requiring the lower boundary of slab (north-dipping thrust-MCT in case of Himalaya) to follow route of upper dashed line (labelled Case 2) or indeed to cease in front of main partially molten region (point labelled 'tip' which in this case would be MCT fault tip where thrust displacement = 0). If lower boundary of slab follows route of lower dashed line (line labelled Case 1), there is no major addition of material into system and volume effects are more minor but may still be significant over time. Both options may be applicable (i.e. have occurred) for different transects across an orogen (see discussion). HHC, Higher Himalayan Crystalline; India, Indian Shield and Palaeo-Proterozoic (meta)sediments; MCT, Main Central Thrust; MHT, Main Himalaya Thrust; STDS, Southern Tibet Detachment System; Tethyan Seq., low grade passive margin sedimentary rocks of Tethyan Sequence.

distribution of mass-balance calculations for mineral phase changes will also provide data. The issue of minor, but incremental volume addition over a large finite deformation (i.e. the total tectonometamorphic extrusion of the HHC) can still have significant influences on flow geometry due to the exponential effect of Equation 10. Figure 5 shows the potentially dramatic effects of incremental volume addition, and would predict a significantly simple-shear-dominated flow type (Wk close to 1) for channel flow with fixed width (i.e. D22 = 1). We suggest that (at least in certain cases) observed plutonism is related to a major increase in volume, as in Case 2. There is good evidence in specific transects across the HHC for pluton volumes that require reciprocal volumes of a partially molten source rock area (allowing melt

fraction of 1 0 - 1 5 % via moderate source rock muscovite content; Vielzeuf & Holloway 1988; PatifioDouce & Johnston 1991) at low temperatures (c. 750~ Montel 1993; Patifio-Douce & Harris 1998) to be much larger than, and therefore significant portions to be located outside of, the extruding HHC. Good candidates are thick ( > 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

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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

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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.

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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

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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

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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).

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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,

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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

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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

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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

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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

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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

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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 _<3 kbar (Fig. 6f). Further to the east, local 'F4' folds appear, tighten and are themselves transposed within planar, steeply dipping high-strain zones (Fig. 5c), with the preceding structural complexity only being apparent in local low strain windows associated with meso-scale F4 fold hinge zones (e.g. Fig. 5d). These planar high-strain panels define the structural geometry of the WPSZ and record sinistral transport parallel to a low-angle southward-plunging lineation. The high-strain shear fabric is associated with development of kyanite-staurolite _+ garnet-bearing assemblages that formed at the expense of syn-S3 andalusite (Fig. 6d). The kyanite-bearing assemblages formed at around 5 kbar, 600~ (Fig. 6f) and are overprinted by sillimanite-bearing fabric (Fig. 6e) in the core of the WPSZ.

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Fig. 4. Channelized extrusion in the Wing Pond Shear Zone (WPSZ). (a) Regional tectonostratigraphic map (DBG, Deadman's Bay Granite; HBGSZ, Hare Bay Gneiss Shear Zone, which contains numerous sheeted granites). (b) Combined metamorphic/structural map. High grade core of the zone is marked by biotite, andalusite and kyanite isograds. Structural transect along the Trans Canadian Highway shows the increase in structural complexity and localization of strain towards the centre of the zone.

Along the eastern margin of the shear zone, local F5 folds appear and the high-strain fabric becomes mylonitic. It is characterized by greenschist-facies assemblages, and appears to correspond to the eastern margin of the WPSZ which is exposed along-strike, SW of the Deadman's Bay Granite (Fig. 4a). The entire WPSZ domain is characterized by mainly steeply plunging mineral-stretching lineations, although there are local high-strain zones with shallow-plunging lineations. The F 3 - 5 folds are geometrically identical sheath folds and appear to reflect progressive deformation and shear localization within the WPSZ. The eastern margin of the WPSZ is flanked by a system of older shallow-dipping fabrics that resemble the flat-lying system west of the shear zone. Viewed

in this context the W P S Z has a simple geometry with a steeply dipping core flanked by domains of progressively lower strain.

Summary The geometry of the W P S Z and its flanking structure is mirrored by both the kinematic and metamorphic character of the structure. The western margin of the shear zone records sinistral kinematics (Fig. 5e) with a component of east-up displacement, while the eastern margin of the shear zone records west-up displacement with dextral component of movement (Fig. 5f). The gradual increase in strain towards the shear zone core suggests that this is a zone of diffuse crustal

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Fig. 5. Field evidence for channelized extrusion from the WPSZ. (a) Flat-lying low-grade psammites and pelites with tight ESE verging and facing fold pair (local F2), typical of the flat belts outside the WPSZ. Northern shore of Gander Lake. Notebook is 190 cm high. (b) Steeply dipping $3-$4 fabric typical of the WPSZ from the shores of Wing Pond. Note the abundant quartz segregations indicative of the higher metamorphic grades in these rocks. (c) Local intensification of $4-$5 fabric (in area to the left of the person) in the core region of the WPSZ; Square Pond quarry (east), just north of the Trans Canada Highway. (d) Typical lower strain augen in F4 fold-hinge zone from Mint Brook revealing underlying structural complexity in WPSZ. Close F4 folds refold tight F3 folds which themselves refold a spaced $2 solution fabric. Lens cap 55 mm diameter. (e) Plan view of sinistral asymmetric boudins of quartz segregations from a high strain zone close to the NW margin of the WPSZ, next to the Trans Canada Highway. ( f ) Plan view of dextral asymmetric boudins of quartz segregations from the SE margin of the WPSZ near Gambo.

extrusion (cf. Fig. 2b), rather than a zone of low strain flanked by high strain margins. There is also an apparent progressive increase in metamorphic pressure towards the central region of the shear zone that is compatible with the kinematics,

suggesting relative upward displacement of the shear zone core. On the flanks, andalusite-bearing assemblages formed at pressure of c. 3 kbar during progressive 'D3' deformation, while in the central region kyanite-staurolite-bearing

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Fig. 6. Petrological evidence for channelized extrusion from the WPSZ (long dimension of all photomicrographs = 7 mm). (a) Cordierite (now retrogressed) overprinted by $3 crenulations. (b) Partially retrogressed andalusite wrapped by $3 biotite-muscovite-bearing foliation. (c) $3 crenulations overgrown by andalusite. (d) Andalusite replaced by kyanite-staurolite-bearing 'D4' assemblage within the centre of the WPSZ. (e) Sillimanite-bearing 'D4' assemblage overprinting syn-D3 andalusite partially replaced by kyanite. ( f ) Metamorphic evolution of rocks in the WPSZ system depicted on a generalized KFMASH P-T pseudosection (§ muscovite-qtz) adapted from Reche et al. (1998) applicable to the aluminous metapelitic compositions that occur within the Gander Zone metasediments. On the flanks of the shear zone andalusite-bearing assemblages formed during progressive 'D3' deformation. Cordierite growth is generally later than andalusite, but still locally shows synkinematic relationships with '$3'. In the core of the shear zone, increasing pressure is implied by the replacement of andalusite by kyanite and staurolite. Subsequent growth of sillimanite is interpreted to reflect decompression associated with exhumation that differentially juxtaposed the now-exposed relatively high-P central region of the shear zone against the lower pressure flanking domains.

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assemblages formed at around 5 kbar on an apparent clockwise P-T evolution (Fig. 6f).

Conclusions Intense study of active orogenic belts such as the Himalayan-Tibetan system emphasizes the inherent difficulty in deriving unique geodynamic interpretations based on available geodetic, geophysical and field-based evidence. In studies of ancient orogenic systems, transient geodetic and geophysical data are generally not available or are less useful, placing greater reliance on field-based metamorphic, structural and geochronological evidence. Thermal-mechanical modelling can be extremely useful in testing the potential influence of various geodynamic boundary conditions, although many of the field characteristics predicted from modelling are common to general zones of crustal extrusion arising in other tectonic regimes, including transpression zones. Field examples from Nanga Parbat in northern Pakistan and the Wing Pond Shear Zone in Newfoundland display strong field evidence of significant crustal extrusion; however, there is a lack of evidence to show that extrusion is gravitationally driven or linked to subhorizontal mid-crustal channel flow. The main driving force of crustal extrusion in these areas is difficult to determine with certainty, but is at least as likely to be tectonically rather than gravitationally driven. Because of the inherent difficulties in identifying the driving mechanisms of lithospheric-scale geodynamic processes, we believe that a clear distinction should be made between primary observations that describe field characteristics of crustal extrusion (Fig. 2a,b), and geodynamic interpretation that tries to relate the data to specific largescale processes or driving mechanisms (Fig. 2c). In this respect, existing usage of the term 'channel flow' is already ambiguous, since it is used variously to describe any or all of the following: (i) a general type of flow condition (in relation to fluid dynamics); (ii) gravitationally driven mid-crustal flow, which if aided by high denudation might propagate upwards to cause crustal extrusion at the surface (in the context of recent thermalmechanical models); and (iii) generalized bulk crustal extrusion, which may extend downwards to connect to subhorizontal mid-crustal zones (in the context of a conceptual explanation for Himalayan field observations such as inverted metamorphism, high shear strains, and contemporaneous thrusting and normal faulting). Consequently, our recommendation is that the terminology used for field description and geometric/kinematic interpretation in orogenic belts should include

terms such as 'crustal extrusion', 'ductile extrusion' or 'channelized extrusion'. Such terms should be considered as broadly synonymous, and refer to zones that show inverted metamorphism, coeval marginal shear with opposing sense etc., without any implication of the far-field driving mechanism. Particularly in older orogenic belts, care should be taken to explain clearly the exact meaning of the term 'channel flow', to discuss any geodynamic implications in relation to the context in which the term is used, and to separate primary field observations from subsequent geodynamic interpretation. R. Law is thanked for a detailed and extremely constructive review of the paper, and extensive editorial advice. U. Ring reviewed an earlier version of the manuscript. R.R.J. thanks NPHM co-workers C. Bond, R. Butler, M. Casey, G. Lloyd, P. McDade and Z. Shipton, as well as M. Bott for input on channel flow. R.E.H. thanks R. D'Lemos and T. King for their many discussions concerning the geology of the Wing Pond Shear Zone. The invaluable logistical support of S. O'Brien, P. O'Neill and F. Blackwood of the Department of Mines and Energy is gratefully acknowledged, together with the assistance of E. Saunders during fieldwork in Newfoundland.

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Interpretation of deformation fabrics of infrastructure zone rocks in the context of channel flow and other tectonic models P. F. W I L L I A M S l, D. J I A N G 2 & S. L I N 3

1Department of Geology, University of New Brunswick, Fredericton, NB, Canada E3B 5A3 (e-mail: [email protected]) 2Department of Earth Sciences, University of Western Ontario, London, ON, Canada N6A 5B7 3Department of Earth Sciences, University of Waterloo, Waterloo, ON, Canada N2L 3G1 Abstract: Infrastructure zones are essentially horizontal to shallowly dipping crustal-scale zones of

non-coaxial flow, with two possible interpretations: (a) a crustal-scale shear zone, transporting upper crust over lower crust and/or mantle (transport flow); or (b) a zone of channel flow in which there is a flux of weak crust between relatively strong upper and lower crust and/or mantle, away from the centre of the orogen. Transport flow has a constant shear sense across the zone, whereas in channel flow the sense of shear reverses across the zone. Channel flow may be driven by extrusion (extrusive channel flow), due to the two channel walls approaching one another, or by a pressure gradient along the channd (normal channel flow), with no convergence of the walls necessary. Arguments based on strain compatibility and mechanics suggest that extrusive channel flow is unlikely. Kinematic vorticity numbers have been used to show that infrastructure and other shear zones have undergone flattening strains, but we show that the numbers are incompatible with extrusion. We also show that in addition to the problems inherent in determining kinematic vorticity numbers from fabric, the numbers cannot be related to bulk flow in mechanically heterogeneous zones, because of flow partitioning. Drag folds are a better indication, albeit qualitative, of whether a zone is thinning or not. They also give a conservative estimate of the minimum accumulated shear strains, and may inhibit strain localization. Like snowball garnets, they indicate shear strains that are so large that the pure-shear-thinning component for a steady-shear zone has to be small.

Regional high-grade metamorphic rocks (amphibolite facies and higher) are c o m m o n l y characterized by a shallowly dipping foliation or layering, except where overprinted by younger structures (cf. Fyson 1971; Mattauer 1973; Mattauer et al. 1981; Zwart 1979; Williams & Jiang 2005). The foliation is generally defined by compositional layering and is a transposition fabric that has developed by folding in a non-coaxial flow regime. The transposed folds are generally recumbent, commonly dismembered, and can occur at all scales from microscopic to regional. The larger folds may comprise fold nappes separated by discontinuities, and the nappe sheets m a y themselves be folded. Lower-grade rocks lying above, generally have a very different structure, with upright folds, more steeply dipping foliations, and reverse faults. This relationship of high-grade and low-grade rocks with the outlined structural characteristics has been referred to as the superstructureinfrastructure association (SIA) (Wegmann 1935; Fyson 1971). In some areas the axial surfaces of

upright folds of the superstructure are seen to bend into the transposition foliation of the infrastructure (Zwart 1979; Murphy 1987). This observation led to the suggestion that upright structures developed in response to crustal shortening, and were then modified by horizontal shear (Fig. 1) in the highgrade rocks of the infrastructure, while being preserved in their original form in the overlying low-grade rocks of the superstructure (Zwart 1979; Krabbendam et al. 1997; Rose & Harris 2000; Williams & Jiang 2005). There are two models of crustal deformation, and nappe or thrust sheet development during orogeny, that are compatible with the d e v e l o p m e n t of the SIA. One, which we refer to as transport flow, is a simple ductile shear zone between strong upper crust and lower crust a n d / o r mantle (e.g. Mattauer et aL 1981). The other is referred to as channel flow (e.g. Grujic et al. 1996; B e a u m o n t et al. 2001). It also has the characteristics o f a shear zone, but the sense of shear changes across the zone. A Rocky Mountain foreland fold and thrust

From: LAW,R. D., SEARLE,M. P. & GODIN,L. (eds) ChannelFlow, DuctileExtrusionand Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 221-235. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

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(a)

(b) superstructure i n f r ~

(c)

Fig. 1. Diagrammatic representation of the superstructure-infrastructure model. (a) Three rows of squares represent different levels of the lithosphere. The upper row represents cool strong crust, the middle row represents weak crust and the bottom row represents medium-strong crust (e.g. basement rocks) or mantle. (b) Crustal shortening produces upright folds in the upper level and the difference in shortening between the upper and lower levels is accommodated by the development of a transport infrastructure zone in the weak middle level. The shear converts the upright folds to overturned folds as they develop. Deformation in this infrastructure zone typically outlasts deformation in the upper superstructure zone. (c) Channel flow develops producing a shear zone in which the sense of shear changes between the top and the bottom, producing folds of opposite vergence. model (Rocky Mountain Model, RMM) has been applied to many infrastructure zones (e.g. Gee 1975a, b; Coward 1980; Boullier & Quenardel 1981; McClay & Coward 1981; Boyer & Elliott 1982; Butler 1982, 1983; Platt 1984; Duncan 1984; McDonough & Parrish 1991; Brown et al. 1992; Gibson et al. 1999). We consider this model, which was conceived in low-grade (generally greenschist or lower) supercrustal rocks, generally inappropriate when applied to high-grade metamorphic rocks of the infrastructure (Ramsay et al. 1988; Williams & Jiang 2005). The essential characteristic of the RMM is that deformation is concentrated around the margins of foliation- or layer-parallel sheets of comparatively

rigid rock. The sheets are locally folded by structures involving the whole sheet or several sheets, but are only weakly deformed internally. Shallowly dipping, approximately zone-parallel shear zones and/or faults are common in the infrastructure (e.g. Williams & Jiang 2005), but they do not have the geometry of the Rocky Mountain structures and are generally late relative to the flow folding (Talbot 1981; Casey & Dietrich 1997; Ramsay 1997; Williams & Jiang 2005). The main shortcoming of the model in this context is that it fails to explain the intense transposition fabric, which is characteristic of the infrastructure, occurring at all scales from regional to microscopic (cf. Williams & Jiang 2005). We therefore reject the RMM as an explanation of the infrastructure and do not discuss it further. Both transport flow and channel flow involve ductile penetrative non-coaxial flow and are capable of explaining the infrastructure fabric. The purpose of this paper is to critically consider some problems related to kinematics and fabric development that we believe to be relevant to understanding the process of deformation in infrastructure zones of transport flow and channel flow. We treat both as shear zones, and distinguish two types of channel flow: extrusive channel flow and normal channel flow. In both types, weak rock flows between two comparatively strong crustal slabs. The channel is generally, but not exclusively (e.g. Zeitler et al. 2001), approximately horizontal. We use the term 'extrusive channel flow' to describe the situation where the weak rock is squeezed from the channel by active convergence of the nearly rigid crustal slabs, i.e. convergence of the channel walls. Extrusive channel flow leads to an excess pressure (overpressure of Robin & Cruden 1994) build-up in the channel (cf. Jaeger 1964, pp. 140-143). We use the term 'normal channel flow' to refer to flow in response to a pressure gradient along the channel. In this case the distance between the channel walls may be constant through time. If the channel thickness changes, it is because of a homogeneous pure shear component of deformation acting on both the channel-bounding wall rocks and the channel itself, or there has to be volume-loss from the channel. The change in channel thickness, if any, does not produce excessive pressure in the channel. Thus normal channel flow is distinct from extrusive channel flow both kinematically and mechanically. Characterization

of the kinematics

of t h e m o d e l s All three models represent crustal-scale subhorizontal high-strain zones and are consistent therefore

INFRASTRUCTURE ZONES AND CHANNEL FLOW

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with penetrative flow, combining varying amounts of shear-zone-boundary-parallel simple shear and pure shear. We discuss each model in turn.

between channel slabs is solely due to a homogeneous pure-shear component resulting from crustal thinning or thickening.

Transport flow

The problem

An edifice develops at the margin of converging plates and important factors in determining its height and width are the rate of convergence, the rate of material flux into the suture, and the rate of erosion. Combined, these factors may result in a temporal change in the mass budget at the subduction zone. If the mass budget is positive due, for example, to an increase in the convergence rate and/or rate of material flux, a widening crustal zone will thicken to a level commensurate with the new mass budget. In this convergent situation a high-strain zone may develop between a shortening upper crust and a more slowly deforming lower crust or mantle below (Fig. lb), leading to a crustalscale subhorizontal shear zone with a constant sense of shear - a transport flow shear zone. Under conditions of crustal thickening the zone of weak (hot) rocks, and consequently the transport flow shear zone exploiting the weak rocks, might also thicken (there would be no reason for it to thin). Thus there would be a pure-shear component orthogonal with the shear-zone boundary as well as a simple-shear component parallel to the zone boundary, and in this situation the pure-shear component would result in thickening of the zone. Notwithstanding, the zone might at the same time be narrowing (widening and narrowing are used in the sense of Means (1995) to describe shear zones that are changing width by their boundaries migrating through the rock mass) by strain localization. However, thickening is probably not a diagnostic feature of transport flow since if the rate of plate convergence decreases or reverses due, for example, to plate rollback, the transport-flow shear zone may continue to be active, due to collapse of the edifice, until a new equilibrium is achieved. In this situation the transport zone might undergo thinning.

In an attempt to evaluate the models outlined above, it is necessary to show that they are capable of explaining the structures and fabric observed at all scales, in infrastructure zones. Further we would like to be able to use structure and fabric to tell us more about the processes involved. Correct interpretation of kinematics based on structure and fabric is of particular importance. For example, if the sense of shear can be determined across the infrastructure zone, it is possible to distinguish between transport and channel flow, and if a channel can be shown to be thickening, extrusive channel flow can be eliminated. However, this whole practice is fraught with difficulties, including the following. (1) Although shear sense is commonly derived from rocks (e.g. Simpson & Schmid 1983; Hanmer & Passchier 1991), it should be remembered that a specific set of fabric elements and structures may indicate sense of shear along a planar marker (e.g. layering), the sense of vorticity relative to the marker, or the sense of the non-coaxiality (reference frame independent) (e.g. Jiang & White 1995; Jiang 1999) and these have different kinematic significances. (2) Determination of kinematic vorticity numbers is predicated on the assumption of steady flow. (3) Flow partitioning resulting from continual strain localization and tectonic transposition make it hard, if not impossible, to derive the bulk kinematics of the infrastructure zone from local fabric analysis (Fig. 2). (4) If a zone is to be interpreted as a channel, it is necessary to show that not only geometry, but also the timing of deformation in different parts of the zone, are compatible with the interpretation. (5) The outcrop may not be representative of the whole zone; for example, if only the lower part of a channel remains due to erosion, it may be interpreted as a transport zone. (6) Complex overprinting relationships may result from migration of the zone or zone boundaries through the crust with time. (7) A zone may change character through time (e.g. from transport zone to channel). Problem (4) is discussed by Kuiper et al. (2006). Here we are primarily concerned with problem (3) - the significance of the kinematic determinations. Kinematic vorticity numbers have been determined from fabrics in infrastructure zones and the results have been used as a means of recognizing whether or not a zone is changing thickness compatible with extrusion (e.g. Grasemann et al. 1999; Xypolias & Koukouvelas 2001; Law et al. 2004;

Channel flow Since the driving force for extrusive channel flow is the convergence of the bounding slabs, the flow is necessarily in a thinning high-strain zone. Normal channel flow could conceivably occur under crustal-thickening or crustal-thinning conditions, depending on whether change in the mass budget is positive or negative. So at any given time the channel walls might be approaching, separating or remaining essentially the same distance from one another. As pointed out earlier, unlike extrusive channel flow, however, the change of distance

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P.F. WILLIAMS E T AL.

(a)

--ST D

(b)

c

~

~

s

Fig. 2. Possible partitioning of shear in a simple-shear zone (ABCD). (a) Foliation (ST) is inclined to the shear-zone boundaries. Shear occurs on the foliation which undergoes orthogonal stretching, and rotation as indicated by the offset marker at the left of the diagram and the various arrows. (b) An ST-parallel layer contains an S-foliation and stretching of the layer is achieved by slip on S, stretching orthogonal with S and rotation of S relative to the layer. Arrows indicate these various components of deformation within the layer.

Xypolias & Kokkalas 2006). We show that the numbers are probably not valid, and further, even if taken at face value, cannot be interpreted as representing the bulk flow. In fact they are all compatible with the infrastructure zone being a simple shear zone, in the case of transport flow, or two contiguous simple shear zones of opposite shear sense, in the case of channel flow. We also discuss dragfolding and its significance as a qualitative measure of thickness change, and of the magnitude of accumulated shear strain in the infrastructure zone.

Kinematic vorticity number and strain compatibility Extrusive channel flow zones must form under thinning conditions whereas normal channel flow and transport flow probably only experience thinning, if any, in their final stages. In general, the latter are more likely to be zones of constant thickness (bulk simple shear) or thickening zones. Thickening and thinning imply a pure shear strain orthogonal with the high-strain-zone boundary and therefore imply a general non-coaxial flow. The degree of non-coaxiality can be represented by the kinematic vorticity number (Means et al. 1980) which, particularly for plane-strain situations, can be represented by the ratio of pure-shear and simpleshear components of the flow (see also Tikoff & Fossen 1993). Thickening and thinning of shear zones implies a kinematic vorticity number less than 1. Several authors (Grasemann et al. 1999;

Xypolias & Koukouvelas 2001; Law et al. 2004) have used kinematic vorticity numbers determined for shear zones to support interpretation of the flow as having a thinning component. However, it must be realized that many problems exist with the methods and underlying assumptions involved in the determination of kinematic vorticity number for shear zones (Jiang & White 1995; Jiang & Williams 1999b), and other interpretations are possible. First, vorticity is an instantaneous quantity of the flow. Unless the flow characteristics are constant over time (steady flow), using a single kinematic vorticity number to represent the flow history is meaningless. In addition, where the flow history is non-steady, the methods used to determine the kinematic vorticity numbers become invalid. In zones where the deformation is heterogeneous, as is generally the case for both normal and extrusive channels, the flow is generally non-steady (Jiang 1994a; Jiang & Williams 1999b). Second, kinematic vorticity numbers determined from rock fabric are, by definition, internal kinematic vorticity numbers, which only record the vorticity component relative to the instantaneous stretching axes of the flow within a given domain (see Jiang 1999). In the process of tectonic transposition, ubiquitous in infrastructure zones (Williams & Jiang 2005), anisotropy and competence contrast between compositional layers lead to a significant component of spin of the instantaneous stretching axes of the flow (Lister & Williams 1979, 1983; Jiang 1994a, b). This spin-accommodated vorticity varies in space, and is not determinable from the

INFRASTRUCTURE ZONES AND CHANNEL FLOW kinematic vorticity numbers derived from fabrics. Thus the determined kinematic vorticity numbers are of questionable value as a representation of the flow in the domain from which the data were gathered; at best, they represent only the kinematics of the flow with respect to the local fabric attractor (Passchier 1997), not the kinematics of the infrastructure zone as a whole. The kinematic vorticity numbers have been determined for several shear zones (e.g. Simpson & De Paor 1997; Grasemann et al. 1999; Xypolias & Koukouvelas 2001; Bailey et al. 2004; Giorgis & Tikoff 2004; Law et al. 2004, Jessup et al. 2006) and have been assumed to represent the kinematics of the zone as a whole. There is wide variability in the numbers: 0.17-0.67 (Simpson & De Paor 1997); 0.4-0.9 (Grasemann et al. 1999); 0.140.99 (Xypolias & Koukouvelas 2001, table 1); 0.67-0.98 (Law et aL 2004); 0.65-0.98 (Jessup et al. 2006); 0.0-0.95 (Xypolias & Kokkalas 2006, fig. 10). Taking a kinematic vorticity number of 0.65 and assuming that the deformation is a plane-strain case (plane strain is an assumption on which many vorticity determination methods are based), the ratio of the pure-shear strain rate (which causes the thinning of the zone) to the simple-shear strain rate is 0.59 (for a relationship between the ratio of pure shear to simple shear and the kinematic vorticity number, see Jiang & White 1995, equation 3) and the convergence angle is 30 ~ (Jiang & White 1995, fig. 5). This represents a significant amount of pure shear. Similarly, on the basis of vorticity analysis, Bailey et al. (2004) proposed 70% and Giorgis and Tikoff (2004) proposed 92% flattening normal to the shear zones that they studied. Such a high component of pure shear poses significant strain compatibility problems which apply to all shear zones (cf. Hudleston 1999), including infrastructure zones, and can potentially be solved in the following ways: (1) by extrusive flow (Fig. 3); (2) by volume loss in the high-strain zone where the volume loss is sufficient to account for the flattening across the zone; (3) by applying the same pure-shear component to the country rock as to the high-strain zone (cf. Jiang & Williams 1998; Lin et al. 1998) so that the two differ only by the simple-shear component present in the high-strain zone (Fig. 4); and (4) any combination of the previous three. We evaluate each end-member situation below. Extrusive flow

If we consider a shear zone in which extrusion occurs in one direction from a point of origin (Fig. 3b), and the pure-shear component is significant at that point, then the strain rate and vorticity will vary progressively away from the origin (Fig. 3). Deformation

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throughout the zone will include a component of pure shear that is identical to the deformation at the origin. However, in addition a unit block at the point of origin has to extrude at a rate commensurate with the rate of pure shear (Fig. 3b). An adjacent block has to undergo the same deformation as well as an additional boundary-parallel shear to accommodate the extrusion of the first block. Therefore, as we move away from the origin, as pointed out by Robin & Cruden (1994) the simple-shear component increases relative to the constant pure-shear component. Figure 3a quantifies this point for a Newtonian material based on Jaeger (1964, pp. 140-143). It can be seen that assuming no shear component at the origin, only four times the instantaneous half zone width from the origin, the kinematic vorticity number, at the high-strain-zone boundary, is already in excess of c. 0.95 and increasing towards 1. It follows that it is highly improbable that the kinematic vorticity number of an extrusive zone would be significantly lower than 1, except close to the point of origin from which the extrusion emanates or near the centre of the zone away from the boundary. For a non-linear viscosity there would be even greater strain concentration near the margin of the zone, and the flow would become more 'plug-like'. In this case the non-coaxiality (and therefore the simple-shear component) would be even more concentrated near the channel margin with the kinematic vorticity number even closer to 1. Theoretically, assuming that the kinematic vorticity number for the zone as a whole can be determined everywhere along the boundary, it should be possible to recognize extrusive zones by the progressive change in kinematic vorticity number parallel to the zone boundary, in the direction of flow. In practice, however, there are additional problems. The point of origin may not be exposed and even if it is, since the change is rapid, the analysis has to be done in sufficient detail and precision and in precisely the right place. It follows that it is not reasonable to interpret the observed high ratios of pure-shear to simple-shear components in terms of extrusion, if the values are measured close to the high-strain-zone boundary, as in the case of the Himalayan measurements of Grasemann et al. (1999) and Law et al. (2004). There is also a mechanical argument against extrusive channel flow between undeformed wall rocks. In an extrusive channel flow, the pressure increase near the origin due to the mean stress (tectonic overpressure; Robin & Cruden 1994) should be so significant that it should be recorded in rock mineral assemblages. There is no evidence of such excess pressure in natural infrastructure zones.

226

P.F. WILLIAMS E T A L .

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INFRASTRUCTURE ZONES AND CHANNEL FLOW orj OJ

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Fig. 4. Rotation (a) and shear strain (b) in plane-strain general shear flow as functions of the amount of flattening across the zone for various values of the kinematic vorticity number (Wk). (a) Finite rotation (revolution of Jiang & Williams 2004) of spherical rigid inclusions. (b) Shear strain parallel to the zone boundary. (c) Diagrammatic representation of the model. As shown in (a), unless Wk is close to 1, very limited amounts of rotation can be achieved for plausible amounts of zone flattening (<50%). Thus the occurrence of snowball garnets in natural shear zones indicates that the kinematic vorticity number was close to 1. It can also be readily seen from (b) that unless Wk is close to 1, the shear zone deformation causes little shear strain along the shear-zone boundary direction for any plausible amount of flattening across the zone. We consider a flattening amount < 50% reasonable for natural shear zones based on arguments presented in the text.

Although Jaeger's (1964) model is based on Newtonian rheology between rigid and parallel walls with non-slipping boundary conditions, modelling based on non-linear theology a n d / o r slightly non-parallel walls (within the limit that tan(4,/2) ~ ~b/2 is valid, where ~b is the angle in radians between the walls) yields similar results (ICG 1999, pp. 2.52-2.57; D. Jiang unpublished data). 'Free'-slipping boundary conditions as alluded to by Sanderson & Marchini (1984) are not possible in nature. In the event of extremely localized slip (and more plug-like channel flow), either the boundary slip is g o v e m e d by a frictional law or there is a low-viscosity material near the boundaries. In either case, the above kinematic arguments do not change except that the excessive pressure build-up

due to mean stress may be reduced (see below). We discuss the deformable channel wall case later.

Volume loss V o l u m e loss could potentially a c c o m m o d a t e zone thinning, and the rate of pure shear might be controlled by the rate of v o l u m e loss. However, in order to explain some of the recorded kinematic vorticity numbers, the volume of material r e m o v e d would have to be so large that considerable compositional differences might be expected between the high-strain zone and the country rock. In our experience such variation is not generally observed between the infrastructure and superstructure, so this is not likely to be a solution to the

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P.F. WILLIAMS ETAL.

strain compatibility problem, although it may be a contributory factor together with one of the other possibilities. Distributed pure shear

Strain incompatibility is avoided if the pure-shear component is applied to the whole crust such that the infrastructure zone is a stretching zone (cf. stretching fault of Means 1990). However, there is a limit to the proportion of pure shear that is possible, because if the pure-shear component is very large relative to the simple-shear component (say 0.59, equivalent to a kinematic vorticity number of 0.65), the amount of shear that can be achieved at reasonable values of zone flattening is small (Fig. 4). Thus, if the pure shear is accommodated by the country rock as well as the high-strain zone and the pure-shear component is large, the only difference between the two will be a relatively small amount of accumulated shear in the highstrain zone. Therefore if in a given high-strain zone the accumulated shear is large, it follows that the pure-shear component must have been small. This point is demonstrated in Figure 4. It can be seen that for a kinematic vorticity number of 0.5 the zone will have been flattened by 80% before the shear strain reaches 15, or before a rotating spherical object such as garnet will have rotated 120 ~ Such a flattening would be expected to be apparent in the country rock as well as the highstrain zone. Further, the low kinematic vorticity number combined with the relatively small shear strain would not be expected to produce snowball garnets, and could not explain the repeated dragfolding (which, as discussed below, also implies high kinematic vorticity numbers and large magnitudes of accumulated shear) commonly observed in such zones. This would not be the typical shear zone. A large component of pure shear might be reasonable during the late stages of transport or channel flow where the final deformation may be driven by the collapse of the edifice and a local thinning of the upper crust and the infrastructure zone. In that situation most of the shear strain may have accumulated earlier when the pure-shear component was smaller, highlighting one of the problems of such studies - the fact that the deformation is not likely to be steady. Application of the pure-shear thinning might be reasonable under conditions of crustal extension. A thinning-normal high-strain zone might well develop under these conditions. We would expect it to be characterized by relatively low accumulated shear strain, by a lack of drag folds (see below), and by such evidence of thinning as asymmetrical shear bands (Williams & Price 1990). Such a shear zone

has been described from the Monashee complex of the Canadian Cordillera (Johnston et al. 2000) and it contrasts with an earlier infrastructure-zone fabric in surrounding rocks, which is characterized by abundant drag folds and is interpreted as a product of channel flow under thickening conditions (Williams & Jiang 2005). Discussion

If significant thinning across the infrastructure zone is unlikely, then the question arises: Do the kinematic vorticity number determinations have the significance attributed to them? We argue that they do not. As for any mechanically heterogeneous shear zone, they do not represent the kinematics of the infrastructure zone. As pointed out previously (Lister & Williams 1979, 1983; Jiang & White 1995), and earlier in this paper, the relationship between fabric and shear zone can be complex and is likely to lead to error. However, even if we assume that the reported kinematic vorticity numbers are correct, it is still improbable that they have the kinematic significance attributed to them, when determined in infrastructure zones. Shear zones generally have two foliations: one that appears parallel to the shear-zone boundary (C foliation), and a shape foliation (S) which is inclined at some angle, like 30 ~ to the C foliation. The latter, as originally described from a homogeneous granitoid (Berth6 et al. 1979), was defined by small-scale shear zones initiated parallel to, and remaining parallel to, the shear-zone boundary throughout deformation. In more heterogeneous rocks the C foliation, as originally defined, may or may not be present. What is generally observed is a layering that is approximately (within the limit of measurement) parallel to the zone boundary, and which can be shown to have developed by rotation and straining of various compositional domains. These domains may be older than the shear zone or may be dykes or veins (Fig. 5), that intruded the shear zone during its development. Such features are generally emplaced at a high angle to the maximum instantaneous stretching axis, and then rotated towards parallelism with the shear-zone boundary. They thus undergo initial shortening and then ongoing extension until at high shear strain they are close to parallel to the shear plane. As they approach parallelism they may be subjected to dragfolding and, so long as deformation continues, this cycle of transposition is repeated. Localization of shear within or between the various transposed layers is common, in keeping with the original definition of C. Thus a typical fabric in the type of shear zone described here comprises compositional domains at all scales. We refer to this foliation as the

INFRASTRUCTURE ZONES AND CHANNEL FLOW

Fig. 5. Sketch of a vertical outcrop surface from the Monashee complex. Leucocratic dykes (grey areas) intrude quartz feldspar gneiss and biotite sillimanite schist. Dyke a is interpreted as the youngest dyke, is undeformed and is believed to be in the approximate emplacement orientation. Dykes b and c are older and have been rotated by top-to-the-right shear and have been folded. Dyke c overprints d which is the oldest dyke. It has been rotated into the extensional field and is unfolded and boudinaged. See text for further discussion.

transposition foliation (ST). In heterogeneous rocks the S-foliation typically occurs in some but not all transposed compositional domains. Both Sa- and the S-foliation may be steady-state foliations in that, for steady conditions, they have a constant fabric and orientation relative to the shear-zone boundary. Whenever perturbed they return to the same state, so long as the deformation continues long enough. In foliated, compositionally heterogeneous rocks, flow partitioning is likely to be complex and to involve active shear parallel to existing foliations (including layering) rather than parallel to the shear-zone boundary. In fact Sx may be active at one scale and the S foliation may be active within individual layers of Sa- (Fig. 2). In this process of continual transposition, anisotropy and competence contrasts between layers lead to vorticity and strain rate partitioning (Lister & Williams 1979, 1983; Williams & Price 1990; Jiang 1994a, b). Using the shear-zone boundary as a reference frame, the vorticity in any compositional layer is the sum of an internal vorticity and a spin component, which is the rotation of the instantaneous stretching axes in the layer relative to the shear-zone boundary (see Jiang (1994b) for details of vorticity partitioning in a layered sequence subjected to imposed bulk flow). The fabric is related to the flow within the specified layer rather than the bulk flow. Further, since the flow is generally non-steady due to the spin component, the fabric will be undergoing continuous modification, and fabric memory of flow becomes an issue. In such a situation, if the overall zone is undergoing say a simple shear, flow reflected by fabrics will in general have a kinematic vorticity number between 0 and 1. The kinematic vorticity number determined from the fabric, by whatever method, will depend: (1) on

229

the flow partitioning the rock sample has experienced, and (2) on the memory of the fabric. To have any meaning at all the fabric has to be assumed to be a snapshot of a relatively steady flow period, and this is hard to justify. Since the transposition process is continuous, and fabrics record different stages of the history of a compositional domain during its transposition, we might expect the determined kinematic vorticity numbers to be widely variable between 0 and 1, despite the fact that the overall shear zone may have an insignificant thinning component.

Significance of drag folds The occurrence of drag folds provides another line of evidence that thinning in infrastructure zones must have been insignificant. Folds in shear zones can be divided into three groups: (a) inherited folds, (b) buckle folds, and (c) drag folds. The orientation of inherited folds relative to shear zones is unlimited. However, in the case of horizontal shear zones developing in response to crustal shortening, as discussed here, there are likely to be upright folds with axial planes steeply inclined to the shear-zone boundary (Fig. 1). Such folds develop an asymmetry related to the sense of shear (Fig. 1) and since both limbs rotate into the same quadrant of both the instantaneous and finite strain ellipses, they will not reverse their asymmetry with increasing strain (cf. Ramsay et al. 1983). Buckle folds develop where veins or dykes are emplaced perpendicular to the maximum instantaneous stretching axis, and rotate into the extensional field. During rotation they initially shorten and buckle. As they rotate into the extensional field they commonly boudinage and may unfold. Folds developed by this mechanism may be oriented such that the shear plane lies within the interlimb angle of the fold. In this situation, as shear strain accumulates the fold will change asymmetry (Ramsay et al. 1983). In our experience this does occur, but it is not common because the folds tend to be preserved at an early developmental stage in boudins, or become unfolded. Drag folds develop in ST, in response to the vorticity (Lister & Williams 1983) and since their axial planes are initially at a high angle to the shear plane, like the inherited folds discussed above, they will maintain their initial asymmetry as long as the sense of shear remains the same (cf. Mawer & Williams 1991) unless perturbed by some local heterogeneity such as refolding. Thus, generally speaking, the folds in shear zones are a good indicator of the sense of shear, especially if they mostly have the same asymmetry. However, once developed they rotate towards the shear direction

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(Hobbs et al. 1976, p. 287; Williams & Zwart 1977; Cobbold & Quinquis 1980; Skjernaa 1989; Mawer & Williams 1991; Jiang & Williams 1999a) and may evolve into sheath folds. In this situation initial S or Z folds will develop both asymmetries and it becomes necessary to recognize the total geometry in order to determine the sense of shear. Geometrically the development of drag folds is like the development of a breaking wave: they are initially symmetrical and become progressively asymmetrical (Mawer & Willians 1991; Lister & Williams 1983; Bons 1993). They may be initiated by rotation of competent domains such as, for example, porphyroblasts at the microscopic scale, boudins at a larger scale, and competent layers at any scale. If we consider the development of a drag fold in a competent marker (Fig. 6) in a zone of simple shear, where the marker is initially approximately parallel to the shear-zone boundary, as the fold develops the shear-zone-parallel subzone enveloping the fold has to thicken locally to accommodate the fold. This necessitates thinning elsewhere by displacement of material away from the crest in the top side of the layer, towards the troughs, and away from the troughs in the bottom side of the layer towards the crests (Fig. 6). The absolute limit on the shearzone-normal height of the crest relative to the trough is the thickness (perpendicular height) of the active shear zone. However, observation tells us that the height of the fold is normally less than the thickness of the zone. We cannot place any quantitative limit on the relative values, but have observed many folds at an early developmental stage which have a height equal to half the width of the zone defined by the nearest planar layers on either side of the fold. Smaller folds (in relative terms) have also been observed but none larger. This is consistent with results of the analogue experiment of Bons (1993). If a pure-shear component is added to the deformation, its effect on dragfolding will depend on whether it leads to thinning or thickening of the zone. If it leads to thickening, there will be a shortening parallel to the shear-zone boundary and thus approximately parallel to Sx. This shortening will initiate and/or amplify perturbations in ST thus

Fig. 6. Drag fold development showing flux of material necessitated by amplificationof the fold.

encouraging the development of drag folds. If on the other hand the shear zone is thinning, concomitant ST-parallel extension will tend to annihilate initial perturbations, thereby inhibiting drag fold development. Again we cannot place quantitative values on the cut-off value for thinning at which drag folds cease to develop. We know that they can develop in simple shear from the experiments of Bons (1993), but cannot limit the value further. However, if drag folds are common in shear zones we would expect that the zone was not undergoing significant thinning when the drag folds were forming. Some impression of the magnitude of bulk shear in a shear zone may also be given by drag folds. Fold hinges and axial surfaces can be treated as passive markers and the amount of strain required to transpose them and convert them to sheath folds can be calculated (e.g. Skjernaa 1989; Williams & Jiang 2005). However, experiments on analogues suggest that this method gives a low answer. In the simple shear experiments of Bons (1993), the development of a fold train is tracked through three stages during progressive simple shear in an annular deformation rig. In the first stage (bulk finite shear strain y = 111) one fold is well developed and two flanking folds are just beginning to develop (Fig. 7). The three folds have axialplanar dips, from left to right, of 73 ~ 30 ~ and 90 ~ At y = 116 the corresponding dips are 43 ~ 21 ~and 28 o, and at y = 122 dips are 27 ~ 20 ~ and 22 ~. The average shear of the axial planes during the first increment is 1.17 with a maximum value of 1.88. During the second increment it is 0.54 with a maximum of 0.89. These values are considerably below the actual incremental bulk shears of 5 and 6. The difference in bulk and local shear strain probably reflects the locally increased shear strength of the system, which is associated with the development of the drag fold (Bons 1993). This would result in a

Fig. 7. Development of drag folds by experimental deformation of analogue materials (camphor layers in octachloropropane matrix) in simple shear (after Bons 1993). See text for discussion.

INFRASTRUCTURE ZONES AND CHANNEL FLOW reduction in the shear-strain rate in the zone of influence of the drag fold. Again it is impossible to place meaningful values on the amount of accumulated shear strain represented by drag folds, but as a rule of thumb it seems reasonable to suggest that drag fold axial planes rotate considerably slower than a passive marker responding to the bulk shear. The rate of rotation will be influenced by such factors as difference between shear strength and normal strength, competency contrast and hinge migration. It will also be influenced by the ratio of pureshear to simple-shear components and the sign of the pure-shear component. Strongly thinning conditions will tend to inhibit drag fold development but will favour rotation of their axial planes, if they form. Strongly thickening conditions will favour drag fold development, but will slow axial-plane rotation and the final orientation of the axial plane will be inclined to the shear zone boundary. The same argument can be applied to the rotation of drag folds into the transport direction and the development of sheath folds. Since the folds result in local shear-strain-rate reduction, any calculation of the bulk shear strain based on their orientation will err on the low side, and judging from Bons' (1993) results the error could be as large as an order of magnitude. This argues for very large strains in shear zones characterized by several generations of drag folds. The most transposed fold in Boris' (1993) experiment (fold a in Fig. 7a) is already well developed in stage 1 (Fig. 7a). After a further bulk shear of 11 (Fig. 7c) it still needs to rotate c. 15 ~ to be within 5 ~ of the orientation of the shear zone, at which stage it might arbitrarily be considered fully transposed. If it continues to rotate at the same rate (i.e. 10 ~ for a bulk shear of 1 l) it will require a further bulk-shear-strain increment of c. 16. It took a bulk shear of 11 to convert an incipient fold (fold b in Fig. 7a) to a form similar to the initial form of fold a, and if that represents the early history of fold a the total bulk shear strain that it would have experienced by the time it is fully transposed would be c. 38. There are obviously many possible sources of error in this estimate, but it gives an indication of the magnitude of shear strain that may be represented by the development of one generation of drag folds. The use of drag folds to determine whether a zone is changing thickness or not, and as a measure of the magnitude of accumulated strain, is obviously imprecise and qualitative at best. However, it is probably more useful than kinematic vorticity determinations, which despite their apparent precision, at best, give only the internal kinematic vorticity of the sampled domain, a number which cannot be related to the bulk flow (see above), wherever there is evidence of flow

231

partitioning. Such evidence is common in infrastructure zones in the form of, for example, boudinage, shear bands and the folds themselves, so that despite the imprecision, the drag fold method is probably the more useful.

Discussion The tectonic models

Based primarily on mechanical grounds, we believe that extrusive channel flow, as the term is used here, is a highly improbable model. Additional evidence is based on kinematic vorticity number determinations. Numbers obtained in the Himalaya, from near to the channel boundary, are interpreted as indicating that the channel has thinned during flow. However, we believe that these numbers are probably inaccurate, and further, are incapable of being interpreted because of partitioning of flow. If a domain has been parallel to ST for a large part of the history and if the fabric memory is relatively short, the number may represent the bulk flow, but there is no way of recognizing this situation, if it even exists. In a continual transposition process, there is no measure of how long a given layer has been parallel to ST, nor do we know how much of the history is preserved in the fabric. Finally, even if the numbers do represent the bulk flow, if they are significantly less than one, and come from near the boundary of the channel, they are not consistent with thinning related to extrusive flow. This is because, as we have demonstrated, extrusive flow results in kinematic vorticity numbers approximating one, close to the channel boundary, except in the immediate vicinity of the origin of extrusion. Transport flow and normal channel flow are both compatible with the transposition foliation and other small-scale structures observed in infrastructure zones. Probably the best way to identify normal channel flow by use of small-scale structures, is to demonstrate the existence of an infrastructure zone in which the vergence of coeval structures changes across the width of the zone. Such a change is recognized in both the Himalaya and the Monashee complex of the Canadian Cordillera, and both have been interpreted in terms of extrusive channel flow (e.g. Burchfiel et al. 1992; Grujic et al. 1996; Grasemann et al. 1999; Johnston et al. 2000). In the Monashee complex the abundance of drag folds indicates that thinning cannot have been significant, supporting the conclusion that the zone is not extrusive, but is a normal channel (Williams & Jiang 2005).

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Strain localization

It might be expected that channel flow would lead rapidly to strain localization such that rocks in the centre of the zone would be experiencing the highest velocity with respect to markers outside the zone, but experiencing the slowest strain rate. In other words, a straight-line marker, initially perpendicular to the flow direction, would develop a parabola-like form. This would result in a gradient in finite strain from the centre of the zone to its margin. Such a gradient is reported in the Himalaya in the form of recognizable shear zones bounding the channel, but is not recognized in the Monashee complex (Williams & Jiang 2005), and may not be a general feature of channels. In the Himalaya the shear zones are young relative to penetrative deformation of the rocks in the channel (Searle et aL 2003) and we speculate that there may have been a tendency for flow to be more plug-like as the distal end of the channel approached the surface. Similar strain localization could have existed in the Monashee complex, but may have been erased by a greater depth of erosion. It may be that a gradient does exist in the Monashee complex, but cannot be recognized. Transposition leads to a steady-state foliation; perturbations develop in a foliation and give rise to drag folds which are transposed back into the foliation, for the foliation to be perturbed again with the start of another cycle of dragfolding. It follows that it is likely to be impossible to distinguish a shear strain of say 40 from a shear strain of 400. However, if this is the explanation, it implies extremely high strains, since we propose accumulated shear strains as high as 40 in the centre of the zone (cf. Williams & Jiang 2005). Another possibility is that the boundaries of a zone are migrating through the crust as the zone develops. This does not imply any change in the active thickness, but is compatible with such change. Thus the whole of the involved crust may have been at the margin of the zone at some stage in its history and therefore in the zone of strain concentration. It is not clear why strain-localization zones should migrate through the rock. In fact it is counterintuitive, since shearing is a geometrically weakening situation and the expected result of strain localization is further localization within the bounds of the initial strainlocalization zone. Changes in temperature resulting in changes in strength, if large enough, might well cause migration of boundaries. The other possibility is that there is in fact little large-scale strain localization in some channels. It could be that in heterogeneous crust the continual development of perturbations inhibits the development of strain-localization zones at the scale of

the perturbations. Strain localization is likely to be favoured by the existence of a planar transposition foliation (ST) approximately parallel to the shear plane. Development of a drag fold in ST renders ST less favourable for strain localization, within the confines of the zone occupied by the fold (cf. Bons 1993). In the Monashee complex, for example (Williams & Jiang 2005), the youngest folds frozen in the process of rotating towards the ST orientation have amplitudes in excess of 2 km so that we might expect no strain localization within a 2 km zone close to the fold. Such folds occur at all levels in the high-strain zone, and their spheres of influence (the volume within which ST is perturbed from its orientation parallel to the shear plane) overlap, so that development of strain-localization zones may be inhibited on an even larger scale, in fact on the scale of the whole high-strain zone. However, if the zone boundary is an interface between weak and relatively strong rock, as proposed, it is probable that drag fold development will be inhibited close to the boundary and this could lead to strain localization and ultimately to the channel flow becoming more of a plug-like flow, as appears to be the situation in the Himalaya. The geometrical surface separating positive from negative shear is not immune to folding so that it will also migrate back and forth through the rock body.

Conclusions Reported variability in kinematic vorticity numbers determined from fabrics in infrastructure zones is compatible with the zone approximating simple shear. This is because fabrics reflect the local flow which is a result of partitioning due to strain localization and tectonic transposition. Flow partitioning is likely to cause temporal variation in the vorticity number and such variation may be reflected in fabrics. Transport shear zone flow and normal channel flow are preferred models for the development of the infrastructure. Strain compatibility problems and mechanical arguments indicate that extrusive channel flow is unlikely. There is probably a natural progression, during orogeny, from transport flow to channel flow back to transport flow. However, channel flow conditions are probably not achieved in all orogens. Both infrastructural shear zone flow and normal channel flow are probably typically associated with thickening of the crust and of the zone, but thinning conditions, particularly in the final stages, are not precluded. Drag folds can develop in simple shear, are favoured by shear-zone-thickening conditions

INFRASTRUCTURE ZONES AND CHANNEL FLOW and are incompatible with significant shear-zone thinning. They also indicate high kinematic vorticity number values and large accumulated shear strains. This is considered to be a useful way of distinguishing shear zones that have developed under crustal-thinning situations (drag folds rare) from those that developed under simple-shear or crustalthickening situations (drag folds abundant). We acknowledge support from Discovery Grants from NSERC. Jiang also acknowledges funding from the Academic Development Funds from the University of Western Ontario. Lin acknowledges support from Lithoprobe. We thank R. Law for discussion and for inviting us to submit this paper, and acknowledge useful reviews by R. Jones, U. Ring and B. Tikoff.

References BAILEY, C. M., FRANCIS, B. E. & FAHRNEY, E. E. 2004. Strain and vorticity analysis of transpressional high-strain zones from Virginia Piedmont, USA. In: AESOP, G. I., HOLDSWORTH, R. E., MCCAFFREY, K. J. W. 8r HAND, M. (eds) Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224, 249-264. 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. BERTHt~, D., CHOUKROUNE, P. & JECOUZO, P. 1979. Orthogneiss, mylonite and non-coaxial deformation of granites: the example of the South Armorican shear zone. Journal of Structural Geology, 1, 31-42. BONS, P. D. 1993. Experimental deformation of polyphase rock analogues. Geologica Ultrajectina: Mededelingen van de Faculteit Aardwetenschappen der Universiteit Utrecht, 110, 1-207. BOUELIER, A.-M. & QUENARDEE, J.-M. 1981. The Caledonides of northern Norway: relation between preferred orientation of quartz lattice, strain and translation of the nappes. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Geological Society, London, Special Publications, 9, 185-195. BOYER, S. E. & ELLIOTT, D. 1982. Thrust systems. Bulletin of the American Association of Petroleum Geologists, 66, 1196-1230. 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. BURCHFIEL, B. C., CHEN, Z., HODGES, K. V., LIU, Y., ROYDEN, L. H., DENG, C. & Xu, J. 1992. The South Tibetan Detachment system: Extension contemporaneous with and parallel to shortening in a

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collisional mountain belt. Geological Society of America, Special Paper 269. BUTLER, R. W. H. 1982. A structural analysis of the Moine Thrust Zone between Loch Eriboll and Foinaven, NW Scotland. Journal of Structural Geology, 4, 19-29. BUTLER, R. W. H. 1983. Balanced cross-sections and their implications for deep structure of the northwest Alps. Journal of Structural Geology, 5, 125-138. CASEY, M. & DIETRICH, D. 1997. Overthrust shear in mountain building. In: SENGUPTA, S. (ed.) Evolution of Geological Structures in Micro- to Macroscales. Chapman & Hall, London, 119-142. COBBOLD, P. R. 8r QUINQUIS, H. 1980. Development of sheath folds in shear regimes. Journal of Structural Geology, 2, 119-126. COWARD, M. P. 1980. The Caledonian thrust and shear zones of N.W. Scotland. Journal of Structural Geology, 2, 11 - 17. DUNCAN, I. J. 1984. Structural evolution of the ThorOdin gneiss dome. Tectonophysics, 101, 87-130. FYSON, W. K. 1971. Fold attitudes in metamorphic rocks. American Journal of Science, 270, 373-382. GEE, D. G. 1975a. A tectonic model for the central part of the Scandinavian Caledonides. American Journal of Science, 275, 468-515. GEE, D. G. 1975b. A geotraverse through the Scandinavian Caledonides-Ostersund to Trondheim. Sveriges Geologiska Unders6kning C, Arsbok 69, no. 9. 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. GIORGIS, S. & TIKOFF, B. 2004. Constraints on kinematics and strain from feldspar porphyroclast populations. In: AESOP, 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, 265-285. 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., KONDIG, R., PAVLIS, T. & SCHMID, S. 1996. Ductile extrusion of the Higher Himalaya Crystalline in Bhutan: evidence from quartz microfabrics. Tectonophysics, 260, 21-43. HANMER, S. 8r PASSCHIER,C. 1991. Shear Sense Indicators: A Review. Geological Survey of Canada, Paper 90-17. HOBBS, B. E., MEANS, W. D. & WILLIAMS, P. F. 1976. An Outline of Structural Geology. Wiley, New York. HUDLESTON, P. 1999. Strain compatibility and shear zones: is there a problem? Journal of Structural Geology, 21, 923-932.

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ICG 1999. FLAC: Optional Features. Itasca Consulting Group, Inc., Minneapolis, USA. JAEGER, J. C. 1964. Elasticity, Fracture and Flow. Methuen, London. 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 channel flow and ductile extrusion of the middle crust. 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. 1994a. Vorticity determination, distribution, partitioning and the heterogeneity and nonsteadiness of natural deformations. Journal of Structural Geology, 16, 121-130. JIANG, D. 1994b. Flow variation in layered rocks subjected to bulk flow of various kinematic vorticities: theory and geological implications. Journal of Structural Geology, 16, 1159-1172. JIANG, D. 1999. Vorticity decomposition and its application to sectional flow characterization. Tectonophysics, 301, 243-259. JIANG, D. & WHITE, J. C. 1995. Kinematics of rock flow and the interpretation of geological structures, with particular reference to shear zones. Journal of Structural Geology, 17, 1249-1265. JIANG, D. & WILLIAMS, P. F. 1998. High-strain zones: A unified model. Journal of Structural Geology, 20, 1105-1120. JIANG, D. & WILLIAMS, P. F. 1999a. When do dragfolds not develop into sheath folds in shear zones? Journal of Structural Geology, 21, 577-583. J1ANG, D. & WILLIAMS, P. F. 1999b. A fundamental problem with the kinematic interpretation of geological structures. Journal of Structural Geology, 21, 933-937. JIANG, D. & WILLIAMS, P. F. 2004. Reference frame, angular momentum, and porphyroblast rotation. Journal of Structural Geology, 26, 2211-2224. 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. KRABBENDAM, M., LESLIE, A. G., CRANE, A. & GOODMAN, S. 1997. Generation of the Tay Nappe, Scotland, by large-scale SE-directed shearing. Journal of the Geological Society, London, 154, 15-24. 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. P. & GOD1N, L. (eds) Channel Flow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 589-611. 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. LIN, S., JIANG, D. & WILLIAMS, P. F. 1998. Transpression (or transtension) zones of triclinic symmetry: natural example and theoretical modelling. In: HOLDSWORTH, R. E., STRACHAN, R. A. & DEWEY, J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 41-58. LISTER, G. S. & WILLIAMS, P. F. 1979. Fabric development in shear zones: Theoretical controls and observed phenomena. Journal of Structural Geology, 1, 283-298. LISTER, G. S. & WILLIAMS,P. F. 1983. The partitioning of deformation in flowing rock masses. Tectonophysics, 92, 1-33. MCCLAY, K. R. & COWARD, M. P. 1981. The Moine Thrust Zone: an overview. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Geological Society, London, Special Publications, 9, 241 - 260. MCDONOUGH, M. R. & PARRISH, R. R. 1991. Proterozoic gneisses of the Malton Complex, near Valemount, British Columbia: U-Pb ages and Nd isotopic signatures. Canadian Journal of Earth Sciences, 28, 1202-1216. MATTAUER, M. 1973. Les ddformations des matdriaux de l'dcorce terrestre. Hermann, Paris. MATTAUER, M., FAURE, M. & MALAVIEILLE,J. 1981. Transverse lineation and large-scale structures related to Alpine obduction in Corsica. Journal of Structural Geology, 3, 401-409. MATER, C. K. & WILLIAMS, P. F. 1991. Progressive strain and foliation development in a sheared coticule-bearing phyllite. Journal of Structural Geology, 13, 539-555. MEANS, W. D. 1990. One-dimensional kinematics of stretching faults. Journal of Structural Geology, 12, 267-272. MEANS, W. D. 1995. Shear zones and rock history. Tectonophysics, 247, 157-160. MEANS, W. D., HOBBS, B. E., LISTER, G. S. & WILLIAMS, P. F. 1980. Vorticity and non-coaxiality in progressive deformations. Journal of Structural Geology, 2, 371-378. 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. PASSCHIER, C. W. 1997. The fabric attractor. Journal of Structural Geology, 19, 113-127. PLATY, J. P. 1984. Balanced cross-sections and their implications for deep structure of the northwest Alps: discussion. Journal of Structural Geology, 6, 603-606. 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 & Hall, London, 445-472. RAMSAY, J. G., CASEY, M. & KLIGFIELD, R. 1983. Role of shear in development of the Helvetic fold-thrust belt of Switzerland. Geology, 11, 439 -442.

INFRASTRUCTURE ZONES AND CHANNEL FLOW RAMSAY, J. G., BUTLER, R. W. H. & COWARD, M. P. 1988. General discussion: On the applicability of cross-section balancing techniques to the internal zones of mountain belts. Philosophical Transactions Royal Society of London, A 326, 321-325. ROBIN, P.-Y. F. & CRUDEN, A. R. 1994. Strain and vorticity patterns in ideally ductile transpressional zones. Journal of Structural Geology, 16, 447 -466. ROSE, P. T. S. & HARRIS, A. L. 2000. Evidence for the Lower Palaeozoic age of the Tay Nappe: the timing and nature of Grampian events in the Scottish Highland sector of the Laurentian Margin. Journal of the Geological Society, London, 157, 381-391. SANDERSON, D. J. & MARCHINI, W. R. D. 1984. Transpression. Journal of Structural Geology, 6, 449-458. 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. SIMPSON, C. & DE PAOR, D. G. 1997. Practical analysis of general shear zones using the porhyroclast hyperbolic distribution method: an example from the Scandinavian Caledonides. In: SENGUPTA, S. (ed.) Evolution of Geological Structures in Micro- to Macro-scales. Chapman & Hall, London, 69-184. SIMPSON, C. & SCHM~D, S. M. 1983. An evaluation of criteria to deduce the sense of movement in sheared rocks. Geological Society of America Bulletin, 94, 1281-1288. SKJERNAA, L. 1989. Tubular folds and sheath folds: definitions and conceptual models for their development, with examples from the Grapesvare area, northern Sweden. Journal of Structural Geology, 11, 689-703. TALBOT, C. J. 1981. Sliding and other deformation mechanisms in a glacier of salt, S Iran. In:

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MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Geological Society, London, Special Publications, 9, 173-183. TIKOFF, B. & FOSSEN, H. 1993. Simultaneous pure and simple shear: the unifying deformation matrix. Tectonophysics, 217, 267-283. WECMANN, C. E. 1935. Zur Deutung der Migmatite. Geologischen Rundschau, 26, 305-350. 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. WILLIAMS, P. F. & PRICE, G. P. 1990. Origin of kinkbands and shear-band cleavage in shear zones: an experimental study. Journal of Structural Geology, 12, 145-164. WILLIAMS, P. F. & ZWART, H. J. 1977. A model for the development of the Seve-Koli Caledonian nappe complex. In: SAXENA, S. K. & BHATTACHARJI, S. (eds) Energetics of Geological Processes. Springer-Veflag, Berlin, 169-187. XYPOLIAS P. 8r KOKKALAS, S. 2006. Heterogeneous ductile deformation along a mid-crustal extruding shear zone: an example from the External Hellenides (Greece). 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, 497-516. XYPOLIAS P. 8r KOUKOUVELAS,I. K. 2001. Kinematic vorficity numbers and strain rate patterns associated with ductile extrusion in the Chelmos shear zone (External Hellenides, Greece). Tectonophysics, 338, 59-77. ZEITLER, P. K., MELTZER, A. S., KOONS, P. O. ETAL. 2001. Erosion, Himalayan Geodynamics, mad the Geomorphology of Metamorphism. GSA Today 11, 4-9. ZWART, H. J. 1979. The geology of the central Pyrenees. Leidse Geologische Mededelingen, 50.

Did the Himalayan Crystallines extrude partially molten from beneath the Tibetan Plateau? T. M A R K H A R R I S O N

Institute of Geophysics and Planetary Physics and Department of Earth and Space Sciences, University of California, Los Angeles, CA 90095-1567, USA (e-mail: tmh @ oro. ess. ucla. edu) Research School of Earth Sciences, The Australian National University, Canberra, A C T 2601, Australia Abstract: The hypothesis that the Himalayan crystalline core originated by ductile channel flow

of partially molten mid-crust from beneath the Tibetan Plateau is critically reviewed. The proposal that widespread shallow anatexis exists beneath southern Tibet today is inconsistent with numerous observations (e.g. 'bright spots' restricted to a single rift and evidence that they represent aqueous fluids rather than molten silicate; the seismogenic southern Tibetan Moho; 3He/4He data indicating the presence of mantle heat and mass in the rift valley; the likelihood that any melt present is due to late Neogene calc-alkaline magmatism; the lack of Tertiary migmatites in the crustal section exposed in the uplifted rift flank of the Yangbajain graben; the lack of Gangdese zircon xenocrysts in the Greater Himalayan Crystallines (GHC); and the broadly coherent stratigraphy in the GHC). Evidence advanced in support of this model is equally or better explained as resulting from localized Neogene calc-alkaline magmatism. A recently developed rapid denudation/channel flow model does explain key petrogenetic and thermochronological features of the Himalaya, but is inconsistent with several geological constraints, most notably the small portion of the collision front over which focused erosion has localized exposure of the GHC. It is concluded that no evidence has yet been documented that requires the existence of partially molten crust flowing in a channel from beneath the Tibetan Plateau to form the Himalaya.

The Earth is an exceptionally complex system that preserves only a partial record of its evolution over a timescale that is difficult to conceive in its depth. Thus geologists tend to describe Earth history by interweaving quantitative modelling with storytelling. The advantage of this approach is obvious. For example, the development of the plate tectonic paradigm through 'geopoesy' permitted geologists to create a vision of planetary dynamics by temporarily overlooking quantitation of the underlying physical processes (Hess 1962; Wilson 1963; cf. Heirtzler & Le Pichon 1965). The disadvantage, however, is that much of what we observe and conclude exists in the form of words and thus is ambiguous or open to misinterpretation in a way that purely mathematical theories are only rarely subject to. The current controversy regarding the proposal that the Himalaya originated via crustal channel flow from beneath Tibet - leading to the Geological Society of London conference on 'Channel Flow, Ductile Extrusion and Exhumation of L o w e r - m i d Crust in Continental Collision Zones' - is in part fuelled by imperfect translations between observation and inference, and qualitative and quantitative

models. The base of this hypothesis rests on a series of factual observations, but some interpretations arising from these data have been poorly justified. This in turn has fed a growing number of offshoot models that may or may not subscribe to all the underlying assumptions of the root hypothesis. Deconstructing the myth that the crystalline core of the Himalaya formed by the extrusion of shallowly formed migmatites originating north of the Indus Tsangpo suture requires that the cross-pollination between qualitative and quantitative texts be separated and analysed in isolation. This review begins by summarizing the consensus view of the geological setting of the HimalayanTibetan orogen and then outlines the salient petrogenetic features of the range that any successful evolutionary model must satisfactorily explain. The qualitative shallow-Tibetan-anatexis model (Nelson et al. 1996) is described, and is followed by a discussion of observations that appear inconsistent with this hypothesis. The work of the Dalhousie group to numerically model the origin of the Himalayan core via crustal channel flow from beneath the Tibetan Plateau (Beaumont et al. 2001, 2004; Jamieson et al. 2004) is then critically examined.

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, 237-254. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

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Geological setting of the Himalaya and Tibet Background

Immediately prior to the onset of the Indo-Asian collision between 60 and 50 Ma (Yin & Harrison 2000; Zhu et al. 2005), the northern boundary of the Indian shield was almost certainly a thinned margin on which Proterozoic clastic sediments and the Cambrian-Eocene Tethyan shelf sequence were deposited (Brookfield 1993). South-directed thrusts in the central Himalaya, including the Main Central Thrust (MCT), Main Boundary Thrust and the Main Frontal Thrust (Fig. 1A; Le Fort 1996), appear to sole into a common decollement, the Main Himalayan Thrust (MHT; Fig. 1B; Zhao et al. 1993; Brown et al. 1996). In general, the MCT places high-grade gneisses of the Greater Himalayan Crystallines (GHC) on top of the Lesser Himalayan Formations (LHF), comprised largely of intermediate to low-grade schists, phyllites, carbonate and minor metavolcanics and gneisses (Fig. 1B). The protoliths of the Lesser Himalayan formations and Greater Himalayan Crystallines are interpreted, respectively, to be Middle and Late Proterozoic clastic rocks (Parrish & Hodges 1996). Geochronologic studies (e.g. Parrish & Hodges 1996; Vance & Harris 1999) suggest that high-grade metamorphism first affected the protolith of the Greater Himalayan Crystallines during an early Tertiary, or Eohimalayan, phase of crustal thickening (Le Fort 1996). The Main Boundary Thrust juxtaposes schists of the Lesser Himalayan Formations (and locally Carboniferous to Permian Gondwanan sequences) against unmetamorphosed Miocene-Pleistocene molasse (Siwalik Group), and the Main Frontal Thrust places Siwalik strata over Quaternary deposits of the Gangetic plain (Fig. 1). Estimates of the amount of slip along the MHT based on balanced cross-section reconstructions (Coward & Butler 1985; Srivastava & Mitra 1994; Hauck et al. 1998; DeCelles et al. 2001; Murphy & Yin 2003) are consistent with a displacement of greater than 400 km. The GHC are juxtaposed below lower-grade Tethyan shelf deposits by the South Tibetan Detachment System (STDS; fig. 1; Burchfiel et al. 1992). The timing of thrusting along the MCT is not well known but constrained by the knowledge that the GHC (i.e. the MCT hanging wall) was deforming at c. 22 Ma (e.g. Hodges et al. 1996; Coleman 1998) and that a broad shear zone beneath the GHC was active between about 8 and 4 Ma (e.g. Harrison et al. 1997). The STDS is known to have been active in several locations

between 17 and 11 Ma (e.g. Edwards & Harrison 1997; Murphy & Harrison 1999), and may have been active earlier. While it is often assumed that slip along the MCT was simultaneous with displacement on the STDS, this is established in only a few cases locally (e.g. Hodges et al. 1996) and the generality of this assumption remains unproven (Murphy & Harrison 1999). If, as originally assumed, the clastic package on the leading edge of India was metamorphosed via thrust imbrication to form the GHC, then it follows that exhumation of this package was via thrust-induced erosion (see review in Le Fort 1996). Following documentation that the STDS, which separates the GHC and Tethyan metasedimentary rocks, was a low-angle normal fault, it was speculated that the crystalline core of the Himalaya was exposed with no net horizontal extension between the STDS and MCT through gravity sliding (Burg et al. 1983; Burg & C h e n 1984), orogenic collapse (Dewey 1988), rigid wedge extrusion (Burchfiel & Royden 1985) or ductile wedge extrusion (Grujic et al. 1996; Vannay & Grassemann 2001; Vannay et al. 2004). More recently, Grujic et al. (2002) proposed that the GHC was extruded as a low-viscosity fluid channel between two parallel shear zones. Recent reviews of the current state of understanding of the evolution of the Himalayan-Tibetan orogen are given in Hodges (2000) and Yin & Harrison (2000), but two salient features of Himalayan geology stand out. Any successful model of the petrogenesis of the crystalline core of the Himalayan range must adequately explain the origin of the classic inverted metamorphic sequences and the paired leucogranite belts. Inverted

metamorphism

The juxtaposition of the GHC and LHF across the MCT is associated at most locations in the Himalaya with an increase in metamorphic grade with higher structural position (i.e. shallower depth; (Fig. 1B; e.g. Arita 1983; P~cher 1989). The GHC vary substantially in thickness across the Himalaya. For example, the MCT hanging wall thickness increases from about 2 km in the central Himalaya (84~ Colchen et al. 1986) to 20 km in Bhutan (89~ Grujic et al. 1996), due to variable initial thickness, the MCT cutting up-section at certain locations, and imbrication within the MCT hanging wall. Thermobarometric studies of the GHC indicate a general decrease in pressure and temperature with increasing distance up-section in the GHC (see review in Harrison et al. 1999). Typically, pressures of c. 8 kbar were achieved adjacent to the MCT (kyanite grade) during the early

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET?

239

;z~ ,.t:z

,...4,

:s2.'~

~d=l

240

T.M. HARRISON

Miocene, whereas peak pressures at the structurally highest levels were only about 3 - 4 kbar (sillimanite grade). Higher pressures (up to 10 kbar) detected locally are generally ascribed to an earlier Barrovian metamorphism termed the Eohimalayan phase (Le Fort 1996). The region approximately bounded by the garnet isograd in the Lesser Himalayan Formations and the MCT hanging wall gneisses of the GHC is typically characterized by a highly sheared, 4 - 8 km thick zone of distributed deformation with a top-to-the-south shear sense, referred to as the 'MCT Zone' (Fig. 1A; Hubbard 1996). Although a variety of models have been proposed linking early Miocene anatexis with the inverted metamorphic sequences, recent studies (Harrison et al. 1997, 1998; Catlos et al. 2001) showed that the dominant Tertiary recrystallization of elements of the MCT footwall largely occurred in the Late Miocene/Pliocene.

700-750~ (Montel 1993). The HHL belt varies in age from 24.0 to 17.2 Ma, but most of the large granite bodies constituting the majority of the leucogranite were emplaced between 23 and 19 Ma (Harrison et al. 1998) The North Himalayan granite belt trends parallel to, and c. 80 km to the north of, the HHL (Fig. 1). Granitoids of the northern belt appear in general to have an elliptical outcrop pattern (e.g. Lee et al. 2004). They differ from the HHL in their emplacement style (Fig. 1) and possibly higher melting temperatures (>750~ suggested by non-eutectic compositions and high light rare earth contents coupled with low monazite inheritance (Debon et al. 1986; Sch~irer et al. 1986; Montel, 1993; Harrison et al. 1997; Lee et al. 2004; cf. Zhang et al. 2004). With minor exception, crystallization ages of the North Himalayan belt range from 17.6 to 9.5 Ma (Harrison et al. 1997).

Paired

Tibetan rifts

leucogranite

belts

An apparently unique feature of the Himalayan range is the presence of two, roughly parallel, syncollisional granite belts, the High Himalayan leucogranites (HHL), which crop out along the crest of the range, and the North Himalayan granites (NHG; Fig. 1A). The HHL form a discontinuous chain of generally sill-like bodies adjacent to the STDS (Fig. 1B) emplaced at temperatures of c.

Although convergence between India and southern Asia continues today, the Tibetan Plateau is currently experiencing east-west extension (Molnar & Tapponnier 1978; Armijo et al. 1986; England & Houseman 1989; Yin 2000). In southern Tibet, extensional strain has been accommodated by a series of generally north-south trending rifts (Fig. 2; Armijo et al. 1986; Yin et al. 1994; Yin

Fig. 2. Neotectonic map of southern Tibet indicating location of the INDEPTH I and 1I seismic reflection surveys (thick grey line) within the Yadong-Gulu rift. NQTL indicates the location of the Nyainqentanghla massif, the uplifted rift flank of the Yangbajain graben. BS, Bangong-Nujiang suture; IS, Indus-Yarlung suture; JS, Jinsha suture. Modified from Kapp et al. (2005).

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET? 2000; Blisniuk et al. 2001; Taylor et al. 2003). Numerous mechanisms have been proposed to explain their development including: expansion of the Himalayan arc (Molnar & Lyon-Caen 1989; Ratschbacher et al. 1994), strain partitioning due to oblique convergence between India and southern Asia (Seeber & Armbruster, 1984; Armijo et al. 1986; McCaffrey & Nabelek 1998), convective removal of mantle lithosphere and associated plateau uplift (England & Houseman, 1989; Harrison et al. 1992; Molnar et al. 1993), gravitational collapse due to maximum sustainable elevation (Molnar & Tapponnier 1978; Armijo et al. 1986; Tapponnier et al. 1986; Dewey 1988), the influence of the Pacific margin causing east-west extension in east Asia (Yin 2000), and concentrated contraction along the central segment of the Himalayan arc (Kapp & Guynn 2004). There are relatively few constraints on the timing of rift initiation across the Tibetan Plateau (Yin et al. 1994; Coleman & Hodges 1995; Harrison et al. 1995; Blisniuk et al. 2001; Stockli et al. 2002; Taylor et al. 2003), but extension across the Yadong-Gulu rift (Fig. 2), the largest rift in southem Tibet, appears to have been underway by 9 Ma (Harrison et al. 1995; Stockli et al. 2002).

The shallow Tibetan anatexis model The Zhao and Morgan

hypothesis

Zhao & Morgan (1985, 1987) developed a model for the evolution of Tibet in which plateau uplift was driven hydraulically via a low-viscosity middle and lower crust beneath the Tibetan Plateau. This approach was a radical departure from previous models that assumed either rigid plate-like behaviour (Argand 1924; Tapponnier et al. 1986) or vertically homogeneous mechanical properties (Dewey & Burke 1973; England & Houseman 1988). The models of Zhao & Morgan (1985, 1987) were tuned to match an uplift history derived from palaeo-botanical results (e.g. Gut 1981) in which plateau growth was essentially a Quaternary phenomenon. With mounting observations that conflicted with this uplift history, and criticism of the method of translating plant fossil data into palaeoelevation information (e.g. Dewey et al. 1988; England & Houseman 1988), the Zhao & Morgan hypothesis was relegated to the ranks of lessfavoured models (e.g. Harrison et aL 1992).

INDEPTH

Between 1992 and 1995, project INDEPTH (INternational DEep Profiling of Tibet and the Himalaya) undertook a near-vertical incidence

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common-midpoint (CMP) reflection survey, as well as companion wide-angle reflection, broadband earthquake and magnetotelluric (MT) studies, along a roughly north-south profile within the YadongGulu rift of southern Tibet (Fig. 2; Zhao et al. 1993; Brown et al. 1996; Makovsky et aL 1996; Nelson et al. 1996; Wei et aL 2001; Xie et al. 2004). The CMP reflection profile revealed a set of prominent reflectors ('bright spots') at depths of 15 to 20 km, beginning in the south at the Indus-Yarlung suture and ending at the north end of the Yangbajain graben, which is the central portion of the YadongGulu rift (Fig. 2). The properties of these reflectors, and their coincidence with a low-velocity zone and electrically conductive crust, led Nelson et al. (1996) to suggest that they mark the top of a midcrustal partial melt layer. Passive seismic results of Li et al. (2003) suggested that the c. 20 km thick layer below that horizon was partial melt, albeit of low melt content. This interpretation was extended to suggest that a fluid, partially molten mid-crustal layer produced by crustal thickening exists throughout southern Tibet (Nelson et al. 1996). Nelson et al. (1996), noting the roots of their model in the work of Zhao & Morgan, proposed that Neogene underthrusting of Indian crust had acted as a plunger, displacing the molten middle crust to the north while at the same time contributing to this layer by melting and ductile flow. In this model, the region between the MCT and STDS is the earlier extruded equivalent of this partially molten region (Fig. 3). The northward younging of the Himalayan granite belts is thus interpreted to reflect a semi-continuous record of this partially molten, mid-crustal layer. Supporting

evidence

Given the unlikelihood of water-saturated anatexis (Clemens & Vielzeuf 1987; cf. Nelson et al. 1996), crustal melting at depths of 15 to 20km requires temperatures appropriate to vapour-absent melting reactions (i.e. >700-750~ Patifio Douce & Harris 1998). The INDEPTH team (Nelson et al. 1996) argued that the high heat flow measured adjacent to the Indus-Yarlung suture (Francheteau et al. 1984; Jaupart et al. 1985) was consistent with shallow Tibetan anatexis and later noted that the upper crustal residence of the Curie isotherm in southern Tibet implied a temperature of about 550~ at a depth of 15-20 km (Alsdorf & Nelson 1999). Gaillard et al. (2004) argued that the similarity between electrical conductivities inferred from MT measurements in the Yangbajain graben and that observed from experimental crystallization studies of leucogranites was evidence in support of the Nelson et al. (1996) hypothesis. Unsworth et al. (2005) interpreted MT data from

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Fig. 3. Interpretive lithosphere-scalecross-sectionof the Himalaya and southern Tibet illustrating the interpretationof underthrusting Indian crust acting as a plunger causing molten middle crust to be displaced southward toward the Himalaya. Thus the region between the MCT and STDS is the earlier-extruded equivalent of the presently partially molten region beneath southern Tibet. Fault abbreviations are given in Figure 1. From Searle (1999); modified from Nelson et al. (1996). several other transects across the Himalaya into southern Tibet in a similar fashion, and proposed a relationship between crustal viscosity and electrical resistivity that is consistent with the shallow Tibetan anatexis model. Mechie et al. (2004) inferred the presence of the a - f l quartz transition at 18 to 32 km depths in a transect across the Bangong-Nujiang suture (Fig. 2), implying temperatures of 700~ and 800~ respectively. During the mid-1990s, thermal models examining the effect of accreting highly radioactive material to the hanging wall of a continental collision under conditions of rapid erosion predicted the necessary high temperatures in the shallow crust (Royden 1993; Huerta et al. 1996; Henry et al. 1997). For example, assuming a convergence rate of 15 mm a 1, an erosion rate of I mm a- 1, and radioactive heat production of 2.5 IxW m -3, Henry et al. (1997) predicted that the 700~ isotherm under southern Tibet would reside at c. 15 km depth after c. 40 million years of convergence. In contrast to the highly contentious and longstanding nature of most debates regarding HimalayanTibetan tectonics (e.g. see Tapponnier et al. (1986) v i s a vis England and Houseman (1988)), the interpretation of Nelson et al. (1996) was quickly adopted by many influential Himalayan workers. Searle (1999, p. 239) wrote that 'similar processes o f . . . melting and leucogranite genesis are occurring today in this zone beneath the Tibetan Plateau as were occurring during the early and midMiocene along the High Himalaya' (Fig. 3). Hodges et al. (2001, p. 802) concluded that 'A channel of middle to lower crustal material extrudes southward from the central Tibetan Plateau between the [STD] and [MHT]' and that 'rocks currently exposed in the Greater Himalayan Zone ... represent the modern leading edge of this feeder channel' (p. 806).

Grujic et al. (2002, p. 178) similarly inferred that 'the [GHC] of Bhutan originated as an orogenic channel that projects for over 200 km to the lower crust of the Tibetan Plateau'. Despite this apparently high level of consensus, there are numerous lines of evidence that seriously challenge the shallow Tibetan anatexis model.

Critique of the shallow Tibetan anatexis model Is the t h e r m a l structure a n d f l u i d activity b e n e a t h the Y a d o n g - G u l u

rift

representative o f Tibet?

The INDEPTH I and II seismic reflection profiles, which imaged << 1% of the southern Tibetan crust, were undertaken along rift valleys whose existence is owed to crustal-scale (Masek et al. 1994) or lithospheric-scale (Yin 2000) extension. Extensive surveys north of the Yadong-Gulu rift (Haines et al. 2003; Zhao et al. 2001) did not image bright spots suggesting that they may be limited to southern Tibetan rift valleys. Based on similar electrical conductivities inferred from INDEPTH MT surveys to those observed in crystallization experiments, Gaillard et al. (2004) suggest that leucogranites are currently forming at shallow levels beneath the Yangbajain graben (Fig. 1). Their interpretation is non-unique and belied by the implausibility of us witnessing widespread anatexis across southern Tibet today. For example, leucogranites make up only about 3% of the present exposure of the Himalaya (Le Fort 1986). The HHL plutons are typically silllike bodies of 200-800 m thickness emplaced at

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET? 15-20 km depth (Scaillet & Searle 2004) and thus are expected to crystallize within c. 105 years of emplacement (Carslaw & Jaeger 1959). If Himalayan leucogranites are 'progressively younger, frozen, snapshots of the partially molten midcrustal layer' (Nelson et al. 1996, p. 1687) and can thus be taken to represent the melting history of the material extruded from beneath Tibet from 24 to 9 Ma (i.e. the known age range of Himalayan leucogranites), then the likelihood of us presently witnessing such an event in a Himalayan-sized portion of southern Tibet is less than one part in 5000 (i.e. 0.03 x (< 100 ka/15 million years)). Unsworth et al. (2005) collected additional MT data from transects to the west and east of the Yadong-Gulu rift which revealed high electrical conductivities at middle and lower crustal depths. While these data may well be characteristic of the subsurface electrical resistivity of the Himalaya, I note that these surveys were also undertaken within or adjacent to north-south trending rifts. Furthermore, their model relating crustal viscosity to electrical resisitivity requires multiple, nested assumptions - most importantly the untestable premise that the low resistivity is dominantly due to the presence of partial melt. The observation that only the upper portion of the GHC experienced Tertiary partial melting (Colchen et al. 1986; Inger & Harrison 1992, 1993) appears to be inconsistent with the Unsworth et al. (2005) interpretation. Because mantle-derived magmatism is commonly associated with continental rifts, it would first seem appropriate to assess the likelihood that geophysical signals of fluid and thermal activity in the Yadong-Gulu rift reflect the emplacement of such magmas before ascribing them to crustal thickening processes. In fact, geochemical and geological investigations described in later sections are consistent with the addition of mantle-derived heat and mass in the Yangbajain region throughout the Late Neogene. Thus it seems unlikely that the thermal structure beneath the Yadong-Gulu rift is representative of the Tibetan crust in general.

Is s h a l l o w anatexis consistent with the cold s o u t h e r n Tibetan M o h o ? Owens & Zandt (1997) found that the seismic velocity structure beneath southern Tibet was indicative of a generally cold crust. Indeed, the lower crust and/or upper mantle beneath southern Tibet, particularly in the region adjacent the YadongGulu rift, is seismogenic (Chen & Kao 1996; Jackson 2002) (Fig. 4). In order for this region to be seismogenic under the strain rates relevant to the Indo-Asian collision, the Moho temperature would have to be less than about 700~ (Ruppel

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Fig. 4. Cross-section along 90~ from Jackson (2002) showing earthquake focus depths (filled circles) and Moho depths (open squares). Earthquakes beneath southern Tibet are occurring at depths of 70-90 km implying a Moho temperature of less than c. 700~

& McNamara 1997; McKenzie et al. 2004; cf. Beaumont et al. 2004). Thus the shallow anatexis model requires an inverted geotherm between c. 15 and 90 km depth under southern Tibet. Mechie et al. (2004) observed P- but not S-wave arrivals along a transect from NW of the Nyainqentanghla massif to the south-central Qiangtang Block. From this they inferred the presence of the a - / 3 quartz transition at 18 to 32 km depths which implies high temperatures (700~ to 800~ respectively). While this interpretation is plausible, it is nonunique. Furthermore, their seismic lines lie in part in Late Cenozoic rifts (see Fig. 2) and thus may not be representative of Tibetan crust in general. Specifically, the highest inferred geotherm lies within the Shuang Hu rift with more southerly, cooler portions in small rifts which represent the terminations of conjugate strike-slip faults (Taylor et al. 2003). D o 'bright spots' represent melts rather than a q u e o u s fluids ? Makovsky & Klemperer (1999) concluded that the velocity properties of the 'bright spots' beneath the Yangbajain graben are best interpreted as porous regions containing about 10% saline aqueous fluids (or > 15%; Li et al. 2003). This conclusion is consistent with the high electrical conductivity of the mid-crust. The presence of active hydrothermal fields within the Yangbajain graben (Fig. 1; Cogan et al. 1998) tends to support this view, although it seems probable that rift-related magmatism is driving the hydrothermal system (see below). A r e 3 H e / 4 H e data consistent with s h a l l o w anatexis ? The isotopic composition of helium from geothermal springs in southern Tibet defines two domains (Yokoyama et al. 1999; Hoke et al. 2000) (Fig. 5).

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T.M. HARRISON et al. 1986; Xu 1990; Miller et al. 1999, 2000; Harrison et al. 2000; Williams et al. 2001; Chung et al. 2003; Kapp et al. 2005). Given the mantle signature in He isotopes from some southern Tibetan hot springs, it is reasonable to assume that this process continues to the present day (Hoke et al. 2000). Thus the locally high heat flow and hydrothermal activity in southern Tibet today (Francheteau et al. 1984) is more plausibly due to the continued emplacement of calc-alkaline magmas than crustal thickening. Is the g e o l o g y a n d g e o c h e m i s t r y o f the uplifted Yangbajain rift f l a n k consistent

Fig. 5, Plot of 3He/4He ratio (=R) of thermal spring waters from the Himalaya and southern Tibet relative to modern atmosphere (=RA) against the sample He/Ne ratio relative to that in air. Because of the negligible helium in air, the horizontal axis permits assessment of the fraction of 3He from a mantle source versus contamination by air. The data define a low 3He/4He 'crustal helium domain', largely restricted to the Tethyan Himalaya, and a high 3He/"He 'mantle helium domain', most notably from the Yangbajain geothermal field. A simple mixing model indicates a mantle contribution of between 1 and 5% 3He (from Hoke et al. 2000).

South of the Indus-Tsangpo suture, 3He/4He isotope ratios are typical of radiogenic helium production in the crust. North of the suture there is a resolvable 3He anomaly. 3He/4He ratios in hydrothermal fluids sampled directly above the Yangbajain bright spots contain the 3He anomaly indicative of a mantle contribution (Yokoyama et al. 1999; Hoke et al. 2000). Hoke et al. (2000) argued that this reflects degassing of volatiles from Quaternary mantle-derived melts intruded into the crust. Given the common relationship between mantle-derived melts and continental rift environments, and the fact that mafic magmas will tend to pond at a depth of 1 5 - 2 0 k m (Glazner & Ussler 1988), the helium isotope data are inconsistent with the notion that equilibrium anatexis brought on by crustal thickening is responsible for the bright spot and MT anomalies in the Yangbajain graben. D o e s crustal thickening explain Tibetan crust t h e r m a l a n o m a l i e s better than episodic calc-alkaline m a g m a t i s m ?

Calc-alkaline magmas are documented to have been emplaced semi-continuously within the Gangdese batholith (the Andean-type arc) between the closure of the Tethys ocean at 6 0 - 5 0 M a and c. 8 Ma (Honneger et al. 1982; Sch~irer & Allegre 1984; Xu et al. 1985; Coulon et al. 1986; Debon

with s h a l l o w anatexis ?

The Nyainqentanghla massif (Fig. 1), which bounds the western margin of the Yangbajain graben, was exposed by a SE-dipping detachment fault which, beginning at about 8 Ma, exhumed an oblique section of crust in its footwall (Harris et al. 1988a, b, c; Pan & Kidd 1992; Harrison et al. 1995; Kapp et al. 2005). Dating of footwall exposures reveals a collage of intrusions including 22 to 8 Ma calc-alkaline granitoids suggestive of continuous or episodic Miocene magmatism (Liu et al. 2004; Kapp et al. 2005). Geochemical and isotopic analyses show a Gangdese-arc affinity indicating significant mantle heat and mass transfer in their formation and are inconsistent with derivation from the Indian craton (Kapp et al. 2005). The undeformed nature of the footwall Cretaceous and Miocene granitoids suggests that the Mesozoic-Cenozoic Lhasa block experienced only upper-crustal penetrative deformation. Coupled with the lack of migmatites exposed in the massif, this fact indicates that the exposed crust was never a zone of anatexis nor involved in large-scale crustal flow (Kapp et al. 2005). Is the a b s e n c e o f G a n g d e s e zircons in the G H C consistent with the s h a l l o w anatexis m o d e l ?

The northern portion of the Yadong-Gulu rift is separated from the Himalaya by the CretaceousTertiary Gangdese batholith (see Fig. 1). Although the shallow anatexis model specified that partially molten (likely Tibetan) crust is being extruded southward (Nelson et al. 1996) from a region through which calc-alkaline magma continued to be injected, not one U-Pb zircon age of the > 1600 samples thus far measured from the GHC (including the North Himalayan Gneiss Domes) is younger than 500 Ma (Parrish and Hodges, 1996; DeCelles et al. 2000, 2004; Myrow et al. 2003). It seems unlikely that either Indian or Tibetan crust

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET? could have been extruded from beneath Tibet through a still-active magmatic zone without incorporating any zircons of Gangdese affinity.

Would the G H C extruded f r o m beneath Tibet contain a definable stratigraphy? The Greater Himalayan Crystallines are characterized by a broadly defined stratigraphy (Formations

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I to III; Colchen et al. 1986). Formation (FM) I, the basal unit, comprises metapsammitic gneisses; Fm II is directly above and dominated by calcsilicates; Fm III is a c. 490 Ma augen gneiss sheet that crops out continuously at the top of large segments of the GHC (Le Fort et al. 1986; Foster 2000; Miller et al. 2001; G. Gehrels pers. comm., 2003) (Fig. 6). This coherence provides a clear constraint on the level of stratigraphic disruption that can take place during extrusion from beneath

Fig. 6. Geological map of a portion of the central Nepal Himalaya showing the distribution of rock units in the Greater Himalayan Crystallines (from Colchen et al. 1986). The generally coherent stratigraphy, particularly that of Fm III, is taken as evidence that the GHC was unlikely to have been ductilely extruded from beneath the Tibetan Plateau.

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the Tibetan Plateau under ductile flow. Although difficult to assess quantitatively, it seems improbable that the regularity of the lower contact of Fm III could have been maintained during a 1000 km + round trip. Summary

Eight questions relating to the shallow anatexis hypothesis (Nelson et al. 1996) are raised in this section. The interpretation of shallow bright spot anomalies beneath the Yangbajain graben as due to in situ anatexis is highly non-unique. More likely they reflect the presence of saline aqueous fluids, observable at the surface as hot springs and geysers, driven by heat originating in the mantle. The location and nature of the seismic and MT anomalies under Yangbajain are consistent with the documented Neogene magmatic history, as recorded in the uplifted rift flank, and the active hydrothermal system in the graben. Helium isotope ratios from the Yangbajain geothermal field, and > 800~ peak magmatic temperatures of young granitoids in the Nyainqentanghla massif (Kapp et al. 2005), record a flux of juvenile heat and mass throughout the Neogene consistent with episodic emplacement of calc-alkaline magmas or lithospheric-scale rifting. The oblique crustal section exposed in the uplifted Yadong-Gulu rift flank reveals no evidence of shallow anatexis prior to rift initiation at c. 8 Ma. The lack of Gangdese-age zircon xenocrysts in the GHC is inconsistent with southward flow of crust across the suture. The coherence of lithostratigraphy within the GHC appears more consistent with thrust imbrication than ductile flow.

Accretion/rapid denudation models Thrust ramp models

During the period of INDEPTH deployment, numerical models were developed which indicated that a thermally mature continental collision zone could evolve a steady-state, near-isothermal structure where the mid-crust has been exposed by rapid denudation at the range front (e.g. Royden 1993; Huerta et al. 1996; Henry et al. 1997). These calculations appeared to provide support for shallow Tibetan anatexis. Harrison et al. (1999) pointed out that models involving the accretion of highly radioactive material to the hanging wall of a continental Himalayan collision under rapid erosion (e.g. Royden 1993; Huerta et al. 1996): (1) required several times more Palaeogene sediment and a far greater depth of Himalayan exposure than known; and (2)

were inconsistent with both synchroneity (at c. 23 Ma) of intrusion of HHL magmas along the 2000 km length of the collision front and the distinctive isotopic characteristics of the GHC and LHF. By emphasizing erosion over accretion, the model of Henry et al. (1997) circumvented criticisms of type (2), but required uniformly rapid erosion from a region equivalent to the c. 200 km wide zone between the trace of the Himalayan thrust system and the Indus-Tsangpo suture. In fact, the level of exposure throughout the vast majority of this region is at greenschist facies with only minor amounts (<15 km) of postOligocene exhumation indicated for much of this area (e.g. Ratschbacher et al. 1994). The deep crustal exposures are largely restricted to the GHC, which form a rather narrower (i.e. 5-100 km wide) aperture than required by the Henry et al. (1997) model. The fact that the Tethyan sediments have not been completely removed from atop the heatproducing-element-enriched Indian supracrustal section represents a severe shortcoming of this model. Indeed, once the Tethyan 'cap' is replaced and a thrust flat introduced into the model, the c. 700~ isotherm under southern Tibet drops from c. 15 km to c. 35 km (e.g. Henry & Copeland 1999). Thus, by the close of the twentieth century, there appeared little in the way of support for the shallow-Tibetan-anatexis model. The landscape changed dramatically with the publication of Beaumont et al. (2001). Focused denudation-induced

channel flow

Challenged by the model of Nelson et al. (1996), Beaumont et al. (2004, p. 28) developed a planestrain, coupled, thermomechanical model assuming a brittle-ductile crustal rheology with a stepped viscosity decrease at 700~ and applied it to the case of the Himalayan collision. By building a high plateau and permitting focused erosion over a narrow aperture at the southern edge of a plateau, the weak, partially molten Indian crust beneath Tibet is extruded along a channel between the MCT and STD to the topographic surface, thereby forming the GHC. The original article (Beaumont et at. 2001) was followed by companion papers (Beaumont et al. 2004; Jamieson et al. 2004) that provided additional documentation for the model and its application to Himalayan tectonics, metamorphism and melting. This framework was far more successful in reconciling model results with petrological, geochronological and geophysical observations of the Himalaya than earlier accretion/erosion treatments. For example, they showed that following c. 55 Ma of Indo-Asian convergence in the 'detached foreland thrust' mode of

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET? deformation, their model could reproduce: (1) the pressure at peak temperature within the GHC; (2) the age at which peak temperature was achieved in the GHC and LHF; (3) the broad ages of the paired granite belts; (4) the general form of pressure-temperature paths in the GHC; (5) the appropriate magnitude of Tertiary sediments shed from the range; and (6) anatexis within the midcrust. This success represented a substantial leap forward, but the model nonetheless faces numerous challenges.

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(see Murphy et al. 1997; Yin & Harrison 2000; Tapponnier et al. 2001; Yin et al. 2002). Although the requirement that Himalayan erosion be forestalled c. 30 million years in order to generate the detached foreland thrust mode is highly specific, it is consistent with the known spatial and temporal distribution of sediment shed from the Himalaya (Yin & Harrison 2000).

P r o b l e m s with p r e d i c t i o n s o f the c h a n n e l flow/rapid denudation model

Limitations o f the f o c u s e d d e n u d a t i o n induced channel flow model

In addition to being subject to many of the same limitations previously enumerated for the shallowTibetan-anatexis model, the focused denudationinduced channel flow model (Beaumont et al. 2001, 2004; Jamieson et al. 2004) is inherently restricted in making detailed comparisons between model results and the observed features of Himalayan geology by: (1) the 2D nature of the model which cannot accurately represent the 3D evolution of an orogenic system that includes, for example, substantial lateral extrusion (e.g. Tapponnier et al. 1986); (2) the lack of coupling of crust-mantle deformation in the model; (3) an initial model temperature structure at c. 55 Ma that is isothermal across the subduction zone (100 million years of Tethys subduction would have significantly depressed isotherms in the wedge); and (4) the assumption of a homogeneous crust ('models with homogeneous crust are unlikely to be representative of natural crustal composition'; Beaumont et al. 2004, p. 22). With regard to the latter, Beaumont et al. (2004, p. 22) emphasized that while 'channel flows can develop even where deformation of heterogeneous lower crust leads to complex middle and lower crust geometry and composition', natural channels resulting from heterogenous crust 'may be difficult to recognize using geophysical techniques'. Extrusion of lower crust by channel flow is driven by a high and extensive plateau and enabled by highly focused erosion at the range front. Thus, choice of elevation and erosion histories are key to determining which of the many modes of deformation will be activated in a particular model. Of concern is the need to initially build an 8 km high proto-plateau in order to stimulate the channel tunnelling mode (Beaumont et al. 2004). Furthermore, the form of the growth history of the plateau, spreading north and south from an initial mountain belt in central Tibet (fig. 11 in Beaumont et al. 2004), appears inconsistent with what is known about the evolution of Tibet

The synchronous pulse of large HHL plutons at 2 2 _ 1 Ma (see Harrison et al. (1997) for a review) across the collision front (with no older plutons of comparable size known) is difficult to reconcile with models in which the thermal budget is dominated by radiogenic heating (e.g. Royden 1993; Huerta et al. 1996; Henry et al. 1997; Beaumont et al. 2001, 2004). Given that several tens of millions of years are required by the radiogenic heating/rapid erosion mechanism to create conditions suitable for melting, small variations in heat generation and/or thermal properties would likely result in the diachronous appearance of melting across the collision front. Instead, the synchroneity of major melting appears more consistent with a mechanism that produces localized thermal anomalies (e.g. minor shear heating). Similarly, the 'detached foreland thrust' mode is subject to the same inconsistencies with large-scale geological observations as the shallow anatexis model. One prediction of a model invoking mid- or lower crustal flow is that crust-mantle deformation should be decoupled. Flesch et al. (2005) evaluated the nature of mechanical coupling through the Tibetan lithosphere by comparing present-day surface (from GPS and slip rates on active faults) and mantle (from SKS shear-wave splitting data assuming anisotropy is a mantle phenomenon) deformation fields. They found that both data-sets could be reconciled if Tibetan lithospheric deformation were vertically coherent (i.e. the maximum shear direction from surface deformation is parallel to the fast polarization direction of olivine). This would only be possible if both velocity boundary and gravity-induced stresses are transmitted to the mantle. While such strong crust-mantle coupling rules out weak southern Tibetan lower crust, the large misfit between surface (from GPS) and mantle (from shear-wave splitting) deformation fields in southeastern Tibet and Yunnan requires complete crust-mantle decoupling there (as suggested by Clark & Royden 2000). An underlying assumption of the Beaumont et al.-type model, and a requirement to create conditions amenable for extrusion by channel flow, is

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that Indian lithosphere is underthrust well beneath Tibet before its mantle lithosphere is subducted (i.e. advancing subduction). However, by comparing post-50 Ma reconstructions of block motions within Asia with tomographic imagery of subducted lithosphere, Replumaz et al. (2004) determined that the Indian plate has continuously overridden its own sinking mantle (i.e. there is no advancing subduction). Thus India does not appear to underthrust Tibet north of the Indus-Yarlung suture. Replumaz et al. (2004) concluded that their observation 'provides further evidence against models of plateau build-up involving Indian lithosphere'. Arguably the most significant criticism of the Beaumont et al. model is its requirement of highly focused erosion that limits the range front to a narrow zone of crystalline rock. In strong contrast, the GHC nappe extends up to 150 km south of the Himalayan front (Fig. 1B) over >70% of the range (Fig. 1A) (Upreti & Le Fort 1999).

U n i q u e n e s s o f p r e d i c t i o n s o f the c h a n n e l flow/rapid denudation model

As remarked earlier, the channel flow (Beaumont et al. 2001, 2004; Jamieson et al. 2004) model reconciles a variety of petrological, geochronological and geophysical observations of the Himalaya, notably thermobarometric data for the GHC and thermochronological results in the GHC and LHF. In fact, the model of Harrison et al. (1998) achieved much the same results using a thermal model in which it was assumed that thrust motion follows fault-bend-fold kinematics (Suppe 1983). Harrison et al. (1998) proposed that the origin of the inverted metamorphic sequences and paired granite belts was linked to minor shear heating on a continuously active thrust that cuts through Indian supracrustal rocks that had previously experienced low degrees of partial melting during a protracted thickening phase in the absence of significant denudation. Numerical simulations assuming a relatively low shear stress of 30 MPa on the shallow Himalayan decollement beginning at 25 Ma triggered partial melting reactions leading to formation of the HHL chain between 25 and 20 Ma and the NHG between 17 and 8 Ma. Late Miocene, out-of-sequence thrusting within the broad but steeply dipping shear zone beneath the MCT provides a mechanism to bring these rocks to the surface in their present location and explains how the inverted metamorphic sequences formed beneath the MCT. In fact, the particle paths of Indian crust created in the Himalayan orogenic wedge in the Beaumont et al. (2001) model are similar to those in Harrison et al. (1998). This underscores the need for any

quantitative model of Himalayan petrogenesis to invoke: (1) an early phase of crustal thickening without significant erosion; (2) advection of footwall rocks into the MCT hanging wall; and (3) a steep MCT ramp to explain the two granite belts and young recrystallization ages within the 'inverted' metamorphic sequences, and the thermobarometric results. It also illustrates that the key petrological, geochronological and geophysical constraints on the evolution of the Himalaya can be equally well or better explained by simple thrust kinematics without an appeal to channel flow.

Conclusions I conclude with reference to the summary statement in the announcement of the meeting on which this volume is based: 'In the Greater Himalayan ranges much discussion at recent meetings has centered around whether the middle or lower crust acts as a ductile, partially molten channel flowing out from beneath areas of over-thickened crust like the Tibetan plateau' (M.P. Searle, pers. comm. 2002). In evaluating the basis of several popular models that advance this hypothesis, both qualitatively and quantitatively, I conclude that there is no observational evidence that requires, or in many cases would even lead one to speculate on, the existence of partially molten middle crust flowing in a channel out from beneath the Tibetan Plateau. Thus the answer to the question posed in the title of this paper is: there is no evidence directly supporting extrusion of partially molten Tibetan crust into the Himalayan core, and there are several lines of contradictory evidence. The proposal of widespread shallow anatexis beneath southern Tibet (Nelson et al. 1996) is inconsistent with a wide variety of observations (restriction of bright spots to rifts and evidence that they represent aqueous fluids, a seismogenic south Tibetan Moho, 3He/4He data indicating mantle heat and mass, existence of late Neogene calc-alkaline magmatism, lack of migmatites in the uplifted rift flank of the Yangbajain graben, lack of Gangdese zircon xenocrysts in the GHC, coherence of the GHC stratigraphy). The success of the Beaumont et al. (2001) channel flow model in reproducing aspects of Himalayan petrogenesis underscores an emerging consensus that signature features of the Himalaya require an early phase of crustal thickening in the absence of significant erosion, advection of footwall rocks into the MCT hanging wall, and a steep, late-Neogene MCT ramp. While intellectually appealing, the Beaumont et al. (2001) model is inconsistent with several geological constraints, most notably the remarkably small portion of the

DID THE HIMALAYA EXTRUDE FROM BENEATH TIBET? collision front in which erosion localized exposure of the GHC to a narrow zone. Whether or not the arguments advanced in this paper convince the reader of the lack of evidence supporting the shallow-Tibetan-anatexis hypothesis or the underlying assumptions of the channel flow model, it is inarguable that ambiguities in interpretation (e.g. the potential for thermal interference between Neogene calc-alkaline magmatism and heating resulting from crustal thickening) prevent definitive selection from among the various proposed models. This raises the question: What kind of new observations are required to break out of the current debate? One advance that would have immediate impact is to undertake reflection profiling outside the rifts. Placing future, transects firmly in the flamework of late Cenozoic tectonics would permit us to assess the validity of the criticism that any reflection profile following crustal thinning features is unrepresentative of the Tibetan crust as a whole. The development of geophysical methods that could remotely observe directional flow in the mid-crust would directly test the hypothesis that the GHC represents the extrusion of anatectic material from shallow depths beneath Tibet. Refinement of tomographic imaging showing the fate of subducted Indian lithosphere would help clarify the general framework of numerical models attempting to reproduce the petrogenesis of the Himalayan crystalline core. I thank the conference organizers for inviting a sceptic into their midst, A. Yin, J. Celerier, A. Aikman, P. Kapp, J.-P. Avouac and P. DeCelles for discussions in which the ideas in this paper were developed, and R. Law, B. Kidd, J. Dewey and P. Treloar for reviews which greatly improved the manuscript. Support from the Australian Research Council is gratefully acknowledged. My greatest thanks to the late Doug Nelson whose intellectual stimulation led my group to return to the Nyainqentanghla to test his hypothesis. His probing mind, tremendous enthusiasm and keen advocacy skills are greatly missed from the community of Himalayan researchers.

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Exhumation of Greater Himalayan rock along the Main Central Thrust in Nepal: implications for channel flow D. M. R O B I N S O N 1 & O. N. P E A R S O N 2'3

tDepartment of Geological Sciences, University of Alabama, Box 870338, Tuscaloosa, Alabama 35487, USA (e-mail: [email protected]) 2Department of Geosciences, The University of Arizona, Tucson, Arizona 85721, USA 3Present address: US Geological Survey, Denver Federal Center, M S 939, PO Box 25046, Denver, Colorado 80225, USA Abstract: South-vergent channel flow from beneath the Tibetan Plateau may have played an important role in forming the Himalaya. The possibility that Greater Himalayan rocks currently exposed in the Himalayan Fold-Thrust Belt flowed at mid-crustal depths before being exhumed is intriguing, and may suggest a natural link between orogenic processes operating under the Tibetan Plateau and in the fold-thrust belt. Conceptual and numeric models for the HimalayanTibetan Orogen currently reported in the literature do an admirable job of replicating many of the observable primary geological features and relationships. However, detailed observations from Greater Himalayan rocks exposed in the fold-thrust belt's external klippen, and from Lesser Himalayan rocks in the proximal footwall of the Main Central Thrust, suggest that since Early Miocene time, it may be more appropriate to model the evolution of the fold-thrust belt using the critical taper paradigm. This does not exclude the possibility that channel flow and linked extrusion of Greater Himalayan rocks may have occurred, but it places important boundaries on a permissible time frame during which these processes may have operated.

The possibility of a linkage between the growth of the Himalayan Fold-Thrust Belt and the development of the Tibetan Plateau is a major topic of interest at the forefront of current research (e.g. Dewey et al. 1988; Harrison et al. 1992; Molnar et al. 1993; Matte et al. 1997; Yin & Harrison 2000; Tapponnier et al. 2001; DeCelles et al. 2002). Despite an obvious spatial link between the Tibetan Plateau and the Himalayan Fold-Thrust Belt (Fig. la), no consensus exists in the literature regarding how tectonic and structural processes active in the fold-thrust belt tie to processes operating on the plateau. Over the past decade, many workers have suggested that in some locations the crust beneath the Tibetan Plateau is unusually weak (see summary by Klemperer 2006), and may be flowing eastward or southward toward the Himalayan Fold-Thrust Belt (Bird 1991; Royden 1996; Royden et al. 1997; Clark & Royden 2000; Shen et al. 2001). Numerical models based upon this crustal flow hypothesis suggest that mid-crustal flow may indeed play an important role in the Himalayan-Tibetan Orogen (Beaumont et al. 2001, 2004; Jamieson et al. 2004, 2006). In these numerical models, channel flow is driven by differential pressure that results from the large topographic differences within the orogen, and operates on a thermally weakened middle crust. Material flowing within this channel is eventually

extruded within the Himalayan Fold-Thrust Belt between the Main Central Thrust and South Tibetan Detachment System, and is removed through focused surface denudation. If this southward flow within a mid-crustal channel and subsequent extrusion has in fact occurred (and possibly may still be occurring), it provides an obvious natural link that connects the structural evolution of the Himalayan Fold-Thrust Belt to the growth of the Tibetan Plateau. In this paper, we attempt to reconcile geological observations gained from detailed mapping in the Himalayan Fold-Thrust Belt of far-western and central Nepal with recently published numerical models (Beaumont et al. 2001, 2004; Jamieson et al. 2004) of channel flow. These numerical models do an admirable job of depicting many of the first-order geological relationships described for the Himalayan Fold-Thrust Belt and the Tibetan Plateau (including the presence of coeval thrust- and normal-sense faults, the presence of high-grade metamorphic material in the core of the fold-thrust belt, and the development of the North Himalayan Gneiss Domes) because the models are built to satisfy these first-order observations. However, as we describe in subsequent sections, some of the geological relationships developed in the numerical models do not match data from the field. In particular, issues that we

From: LAw, R. D., SEAgLE,M. P. & GODIN,L. (eds) ChannelFlow, Ductile Extrusion and Exhumation in Continental Collision Zones. Geological Society, London, Special Publications, 268, 255-267. 0305-8719106/$15.00

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Fig. 1. (a) Digital image of the Tibetan Plateau and Himalayan thrust belt (from the Global 30 Arc Second Elevation Data Set). Nepal is outlined in white for reference. (b) Tectonostratigraphic map of Nepal modified from Robinson et al. (2001). Geology of Kumaon, India (modifiedfrom Johnson 2005), showing correlation between the Dadeldhura/ Karnali Klippe (DKK) and the Almora Klippe (AK). Question marks denote correlation problems from Nepal into India.

address include: (1) whether all rocks from the Greater Himalayan tectonostratigraphic zone in the Nepal portion of the fold-thrust belt flowed within a mid-crustal channel; (2) constraints on the timing of extrusion of the Greater Himalayan zone rocks; and (3) how detailed structural observations in Lesser Himalayan zone rocks affect models for extrusion of Greater Himalayan rocks. Although we do not dismiss channel flow and ductile extrusion as important processes that may have operated in the Himalaya, we suggest that incorporating the field-based observations that we document into future numerical models will contribute to a greater understanding of the extrusion/emplacement of Greater Himalayan rocks, and aid the process of unravelling the geodynamic evolution of the Himalayan-Tibetan Orogen.

Background on channel flow in the Himalayan-Tibetan Orogen The southward flow of mid-crustal material from beneath the Tibetan Plateau was first suggested by Bird (1991). Since then, many workers have contributed to an understanding of the processes that

govem channel flow, and the possible application of various channel flow models to the HimalayanTibetan Orogen. Although it is beyond the scope of this paper to give a detailed review of the channel flow literature as it applies to the Himalaya, it is important to summarize a few key lines of evidence which suggest that channel flow may play an important role in the geodynamic evolution of the orogen. The primary drivers for channel flow in the Himalayan-Tibetan Orogen are the dual observations that: (1) unusually weak zones exist in the crust at different depths throughout the modern Tibetan Plateau (see summary in Klemperer 2006); and (2) that gravitational/pressure-driven flow could be a natural consequence of the topographic and crustal thickness variations between the Tibetan Plateau and the Indian foreland. Based upon results from the INDEPTH studies, Nelson et al. (1996) and Hauck e t al. (1998) proposed that crust on the leading edge of the subducting Indian plate partially melts below the Lhasa terrane, and flows southward toward the Himalayan orogenic front, where it is denudationally extruded between the coeval Main Central Thrust and South Tibetan Detachment System as Greater Himalayan zone rocks.

GREATER HIMALAYAN EXHUMATION Indeed, several lines of evidence gained directly from Greater Himalayan rocks currently exposed in the High Himalaya support this general concept. Mineral assemblages and thermobarometric datasets, rock fabrics, and the widespread presence of migmatites and leucogranites in most Greater Himalayan rocks (Searle & Szulc 2005 and references therein) support the idea that the rocks may have flowed at mid-crustal depths in a channel which was subsequently exhumed. Additionally, the normal-sense South Tibetan Detachment System and the thrust-sense Main Central Thrust provide logical structures that may serve as the upper and lower bounding surfaces for channel flow and subsequent extrusion/exhumation (e.g. Grujic et al. 1996, 2002; Wu et al. 1998; Vannay & Grasemann 2001). Initial timing of motion on the Main Central Thrust is dated at c. 20-22 Ma (Hubbard & Harrison 1989; Hodges et al. 1996; Johnson et al. 2001), although it may be as young as c. 16 Ma (Catlos et al. 2002; Kohn et al. 2004, 2005). Some workers present evidence to suggest that the fault may currently still be active (Hodges et al. 2004). The South Tibetan Detachment System was active by Miocene time and may have continued in some areas to the Pleistocene (see summary in Hodges 2000). In summary, a general consensus in the literature is that motion on the faults was broadly coeval (see review by Godin et al. 2006), which is one of the requirements of channel flow models currently articulated in the literature. Channel flow models for the Himalayan-Tibetan Plateau orogenic system are particularly intriguing because of the proposed link between channel flow and ductile extrusion driven by focused denudation at the Himalayan topographic front (Beaumont et al. 2001). This link between flow and extrusion in the Himalaya is critical in both conceptual and numerical models currently articulated in the literature. In this paper, we argue that several aspects of these channel flow models are problematic. In subsequent sections, we discuss evidence from the external crystalline klippen, Greater Himalayan rocks, and Lesser Himalayan rocks in the proximal footwall of the Main Central Thrust.

Insights from the External Greater Himalayan klippen Across Nepal, Greater Himalayan zone rocks are currently exposed in two distinct settings (Fig. lb). In the northern part of the country, Greater and Tibetan Himalayan zone rocks form the peaks of the 'High Himalaya'. To the south of the High Himalaya, Greater Himalayan rocks are exposed in klippen that rest upon Lesser Himalayan

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zone rocks. These klippen are commonly known as the 'Lesser Himalayan Crystalline Nappes', and include the Almora, Dadeldhura, Jajarkot and Kathmandu klippen (Figs lb & 2). Although there has been debate in the literature regarding whether rocks in the klippen are of Greater Himalayan zone affinity, recent geochronological and Nd isotope studies (Robinson et al. 2001; Gehrels et al. 2003) show that the metamorphic rocks in the Kathmandu Klippe are indeed of Greater Himalayan zone affinity. Likewise, there is a debate surrounding the structural manner of klippen emplacement (Johnson 2005). The klippen are either part of a major thrust sheet that lies structurally below the Main Central Thrust (R~ii et al. 1998; Upreti 1999; Upreti & Le Fort 1999), or part of a major thrust sheet that lies structurally above the Main Central Thrust (e.g. Langtang thrust; Kohn et al. 2004), or part of the Main Central Thrust sheet itself (Johnson et al. 2001; Johnson 2005; Pearson & DeCelles 2005). In the latter case, the fault that carries the klippen is the southern continuation of the Main Central Thrust. For reasons detailed in Johnson et al. (2001), Johnson (2005) and Pearson & DeCelles (2005), we view the fault that carries the klippen as the southern continuation of the Main Central Thrust. Greater Himalayan zone rocks in the external klippen impact the channel flow debate for the following reason. Current channel flow models purport to explain the presence of all Greater Himalayan zone rocks currently exposed in the fold-thrust belt. Just as the models explain many of the characteristics of the Greater Himalayan zone rocks in the High Himalaya, they must also explain observed geological relationships in the klippen. In this section we focus on rocks of the Kathmandu Klippe in central Nepal (Fig. lb), as more structural, stratigraphic and thermobarometric data exist for these rocks than for those in any of the other klippen. As previously mentioned, the Kathmandu Klippe is underlain by a fault that we regard as the southern continuation of the Main Central Thrust. The fault is locally known as the Mahabharat Thrust (Sttcklin 1980; Sttcklin & Bhattarai 1982; Upreti & Le Fort 1999), and we map the fault in the same location as St6cklin & Bhattarai (1982). Detailed geological maps showing the position of the Main Central Thrust below the Kathmandu Klippe can be found in Pearson & DeCelles (2005, figs 9, 10 & 11). The Main Central Thrust juxtaposes hanging wall rock that consists of the Greater Himalayan zone Bhimphedi Group (St6cklin 1980; St6cklin & Bhattarai 1982) against lower Lesser Himalayan zone rocks carried by the Ramgarh Thrust (Pearson & DeCelles 2005). In this sense,

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the structural contact between Greater and Lesser Himalayan zone rocks is identical to the contact observed at the base of the High Himalaya across Nepal, where Greater Himalayan zone rocks overlie Lesser Himalayan zone rocks carried by the Ramgarh Thrust (Pearson & DeCelles 2005). Thus the Main Central Thrust below the Kathmandu Klippe could potentially serve as the lower shear zone boundary for a channel that was subsequently extruded/exhumed. Although a potential lower boundary exists in the Kathmandu Klippe, there is no evidence to support the presence of a fault or shear zone as an upper boundary (an equivalent of the South Tibetan Detachment System) that separates Greater Himalayan from Tibetan Himalayan rocks. Within the klippe, Tibetan Himalayan zone rocks belong to the Phulchauki Group (St6cklin 1980). St6cklin (1980) described the contact between rocks of the Bhimphedi and Phulchauki Groups as either transitional or as a minor unconformity. More recently, Gehrels et al. (2006) used ages from detrital zircons to revise the stratigraphy of the Bhimphedi (Greater Himalayan) and Phulchauki (Tibetan Himalayan) Groups, and suggest that the contact is a disconformity. The lack of an upper bounding shear zone clearly suggests that existing channel flow models that describe the emplacement of Greater Himalayan rocks in the High Himalaya may not apply to the Kathmandu Klippe. Stratigraphic data for the Kathmandu Klippe also suggest that channel flow and extrusion models applied to the High Himalaya may not be applicable for the klippen. Gross stratigraphic relationships within the well-bedded Bhimphedi Group appear to be intact (St6cklin 1980; Gehrels et al. 2006), and in the more quartz-rich lithologies, primary sedimentary structures such as cross-beds and ripple marks can be observed (St6cklin & Bhattarai 1982). If rocks of the Bhimphedi Group flowed within a channel, it is highly likely that original stratigraphic relationships and sedimentary structures would have been obscured. Equally important is the lack of both migmatites and leucogranites (as compared to the High Himalaya) from the stratigraphy of the Kathmandu Klippe. Searle & Szulc (2005) suggest that the widespread presence of these rock types in the High Himalaya lubricated flow within a mid-crustal channel, and that the presence of these lithologies can be regarded as a boundary condition for the orogenic channel flow model. It is possible to argue that the lack of migmatites and leucogranites in the Kathmandu Klippe is merely a function of the fact that klippe rocks may have occupied a higher structural position than Greater Himalayan rocks exposed in the High Himalaya. Indeed, metamorphic grades in

the klippen are generally less than in the High Himalaya (see summary in Hodges 2000). Perhaps the Bhimphedi Group rocks occupied a position similar to points G3 and G4 in model HT1 of Jamieson et al. (2004), and simply represent crust that flowed outward with the underlying channel. This is an awkward solution, however, for the following three reasons: (1) the absence of an underlying hot channel ( s e n s u Jamieson et al. 2004) would have to be explained; (2) the absence of an upper bounding shear zone is problematic; and (3) metamorphic conditions in the lower part of the Kathmandu Klippe reached as high as c. 700~ and c. 900MPa (Johnson & Rogers 1997; RS_i et al. 1998; Johnson et al. 2001; Johnson 2003). These temperatures and pressures are significantly higher than those attained in Model HT1 (Jamieson et al. 2004), where maximum temperatures and pressures for points G3 and G4 are 605~ kb and 441~ kb, respectively. The preponderance of evidence available for the Kathmandu Klippe suggests that Greater Himalayan rocks within the klippe were not involved in channel flow, s e n s u Beaumont et al. (2001, 2004) and Jamieson et al. (2004). If this is the case, could channel flow and subsequent extrusion/exhumation have occurred for the High Himalaya but not for the external klippen? It may be possible to argue that Greater Himalayan rocks of the klippen originally occupied a position just south of the southernmost Greater Himalayan rocks that participated in channel flow. If this is the case, then the following two interesting possibilities must be considered: either (1) the amount of shortening accommodated by the Main Central Thrust at the base of the High Himalaya must be an order of magnitude less than the c. 950 km estimated by Jamieson et al. (2004) in order to allow a link between the High Himalaya and the klippen on a single thrust sheet; or (2) the external klippen must not be part of the Main Central Thrust sheet. Although it is impossible to accurately quantify the amount of shortening accommodated by the Main Central Thrust, as the leading edge of the thrust sheet is eroded (Fig. 2), it is unlikely that the fault accommodated < 100 km of displacement (see balanced cross-sections in DeCelles et al. 2001; Robinson 2001; Pearson 2002; Robinson et al. 2006). As already mentioned, previous workers (R~ii et al. 1998; Upreti 1999; Upreti & Le Fort 1999) suggest the possibility of a two-thrust model for the Kathmandu Klippe. Johnson et al. (2001) and Pearson (2002) argued against such a model, but the possibility remains that supporting evidence could be discovered in the relatively unknown Gosainkund region to the north of Kathmandu. Regardless of the viability of either of these two solutions, or the advent of others, it is clear that current

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Fig. 2. (a) Simplified cross-section through far-westem Nepal; location shown by line A-A' in Figure lb. This cross-section is modified from Robinson (2001) and is simplifiedfrom Robinson et al. (2006; this cross-sectionis B-B' in their figs 3 & 4). Lithologic patterns and abbreviations are the same as in Figure 1 (LHD, Lesser Himalayan Duplex). (b, c, d) Cartoons showing emplacement of thrust sheets detailed in the cross-section. The active fault in each cartoon is in bold. Numbers 1-5 correspond to both the sequence of faulting and to the following faults: 1, MCT; 2, RT; 3-5, faults within the LHD. (b) Greater Himalayan rock emplaced along the MCT. (c) Emplacement of the RT sheet 2. (d) Development of the LHD.

channel flow models need to be modified so that they can reproduce fundamental geological relationships present in the klippen.

Timing of Greater Himalayan rock extrusion The numerical channel flow models (Beaumont et al. 2001, 2004; Jamieson et al. 2004, 2006) make specific predictions about when Greater Himalayan rocks are extruded/exhumed from the channel and eroded. These predictions can be validated by provenance data-sets from synorogenic sediments, and by data-sets of mineral cooling ages within Greater Himalayan rocks. Jamieson et al. (2004) recognize that model cooling ages for Greater Himalayan rocks are too young, and state that future models need to be modified to address this discrepancy. In this section, we summarize evidence which suggests that Greater Himalayan rocks were largely emplaced prior to the Middle Miocene. This evidence is important because, if correct, it helps to constrain the permissible timing for extrusion/exhumation and cooling of Greater Himalayan rocks in future numerical channel flow models. Studies of synorogenic sediments (e.g. DeCelles et al. 1998, 2004; Robinson et al. 2001; Najman et al. 2005; Najman 2006) suggest the following sequence and timing for erosion of sediment from the rising fold-thrust belt: (1) in Nepal, the Eocene Bhainskati Formation records the first

influx of sediments derived from Tibetan Himalayan rocks and from the Indus-Yarlung suture zone; (2) the first sediments with a Greater Himalayan affinity appear at c. 15 Ma in the uppermost Dumri Formation and the lowermost Siwalik Group; (3) synorogenic sediments bearing a Lesser Himalayan isotopic signature first appear at c. l l - 1 2 M a . Although numerical models (Beaumont et al. 2001, 2004; Jamieson et al. 2004, 2006) do a reasonable job of conforming to this sequence, some significant differences exist: in model HT1 and model HT111, both Tibetan and Greater Himalayan detritus are shed from the rising fold-thrust belt as early as 24 Ma, and Lesser Himalayan rocks begin to be eroded at c. 15 Ma. The significance of these differences in timing will be discussed below. Mineral cooling ages also help to constrain the timing of exhumation of Greater Himalayan rocks. The 40 Ar/ 39 Ar data from Greater Himalayan muscovites generally show that the Greater Himalayan sequence began to cool through the c. 350~ blocking isotherm during the Early Miocene; cooling continued through the Middle and Late Miocene, and Pliocene dates are reported from rocks proximal to the trace of the Main Central Thrust in the High Himalaya (see summary in Jamieson et al. (2004) and references therein). This range of ages is consistent with an interpretation that suggests that the bulk of Greater Himalayan slab exhumation occurred during the Miocene. By contrast, cooling below the c. 350~ isotherm of Greater Himalayan rocks in the numerical models (Beaumont et al.

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2001, 2004; Jamieson et al. 2004, 2006) occurs mostly during the Pliocene (Jamieson et al. 2004). Based upon provenance and cooling data, as well as structural observations from Lesser Himalayan and Subhimalayan rocks, Robinson et al. (2003) proposed a kinematic model for development of the Himalayan Fold-Thrust Belt in Nepal. This kinematic model honours the provenance and cooling data discussed above. Although the kinematic model developed by Robinson et al. (2003) is a relatively conventional critical taper-type model, it does not exclude the possibility that channel flow and subsequent extrusion/exhumation may have played an important role in early development of the fold-thrust belt. However, the kinematic model does place some important constraints on the mode and timing of Greater Himalayan rock extrusion/exhumation; therefore, we briefly summarize the main points. (1)

(2)

(3)

(4)

In Early Miocene time, an immense slab of Greater and Tibetan Himalayan rock was emplaced along the Main Central Thrust above a long footwall flat of lower Lesser Himalayan rock (Fig. 2b). During the Middle Miocene, the Ramgarh Thrust, which carried lower Lesser Himalayan rock in addition to its overburden (Greater and Tibetan Himalayan rock), was emplaced as a hanging wall flat over other Lesser Himalayan rocks in its footwall (Fig. 2c). In western Nepal, major slip on the Ramgarh Thrust occurred between c. 15 and 11 Ma (DeCelles et al. 2001; Robinson 2001; Pearson & DeCelles 2005), whereas in central Nepal, slip occurred between c. 11 and 9 Ma (Kohn et al. 2004). Beginning in the Late Miocene, slip was transferred to other Lesser Himalayan thrusts which began to form a large duplex (the 'Lesser Himalayan Duplex'), for which the Ramgarh Thrust served as the roof thrust (Fig. 2d). Growth of the Lesser Himalayan Duplex uplifted and arched rocks structurally above the Ramgarh Thrust, including Greater and Tibetan Himalayan rocks. Growth of the Lesser Himalayan Duplex ended by c. 5 Ma, when slip was transferred to the Main Boundary Thrust system; however, some minor out-of-sequence thrusting in the hinterland is completely consistent with the growth of a forward-propagating thrust system (Davis et al. 1983; Thiede et al. 2004).

The kinematic model's main constraint on the timing of Greater Himalayan rock extrusion is based upon the fact that across Nepal, Greater Himalayan rocks carried by the Main Central Thrust always rest upon lowermost Lesser Himalayan rocks carried by the Ramgarh Thrust (Pearson &

DeCelles 2005). The Ramgarh Thrust is the oldest structure contained entirely within Lesser Himalayan rocks (as it is deformed by younger structures), and the fault is present at the base of the High Himalaya, beneath the klippen, and in the proximal hanging wall of the Main Boundar-y Thrust (Figs lb & 2; Pearson & DeCelles 2005). This means that, barring major out-of-sequence thrusting, Greater Himalayan rocks in the hanging wall of the Ramgarh Thrust were emplaced prior to the onset of motion along the Ramgarh Thrust in the Middle Miocene. Although some workers have argued that geochronologic and thermochronologic data from central Nepal (Harrison et al. 1997, 1998; Catlos et al. 2001, 2002) support major out-ofsequence displacement on the Main Central Thrust, Robinson et al. (2003) showed that the data can also be explained by growth of the Lesser Himalayan duplex; therefore major late Miocene reactivation of the fault is not required. Participation by currently exposed Greater Himalayan rocks in channel flow and subsequent extrusion/exhumation (sensu Beaumont et al. 2001, 2004; Jamieson et al. 2004) therefore must have ended by the Middle Miocene before slip was transferred to the Ramgarh Thrust and into the Lesser Himalayan Duplex. By contrast, the numerical models show extrusion beginning in the Middle Miocene and continuing to the present day. An additional issue related to the timing of Greater Himalayan rock extrusion/exhumation concerns the location and timing of a surface denudation point, which is the location through time of the erosional porthole through which Greater Himalayan rocks were extruded/exhumed. Data incorporated into the kinematic model developed by Robinson et al. (2003) suggest that the Main Central Thrust sheet was emplaced prior to the onset of displacement on the Ramgarh Thrust at c. 15 Ma. During emplacement of the Main Central Thrust sheet, the topographic front of the growing range (which we assume to be the most likely location for focused denudation) migrated southwards, with respect to the footwall cut-off of the Main Central Thrust (Fig. 3a). Beginning in the Middle Miocene, slip was transferred to the Ramgarh Thrust (Pearson & DeCelles 2005), which also carried structurally higher Greater and Tibetan Himalayan rock. During emplacement of the Ramgarh Thrust sheet, the surface denudation point continued its southern migration, relative to its footwall cut-off (Fig. 3b). In the Late Miocene, growth of the Lesser Himalayan Duplex resulted in a large culmination that arched rocks carried by both the Ramgarh and Main Central thrusts. Provenance data from synorogenic sediments (DeCelles et al. 1998; Robinson et al. 2001) indicate that in western Nepal, the duplex was erosionally breached by c. 11-12 Ma.

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Fig. 3. Schematic cartoon showing evolution of the fold-thrust belt in western Nepal since Early Miocene time. Each time sequence is hung on the RT ramp illustrating the relative magnitude of southern propagation of the faults. Active fault is in bold. Note approximate horizontal scale. Lithologic patterns and abbreviations are the same as in figs 1 & 2, with the addition of the Lesser Himalayan Bhainskati and Dulnri formations. (a) Development of the Main Central Thrust with emplacement of the Greater Himalayan rock over Lower Lesser Himalayan rock. Note the location of the erosional porthole for the 'channel', marked as the surface denudation point of the Greater Himalayan rock. (b) The Ramgarh Thrust emplaced lower Lesser Himalayan rock over Upper Lesser Himalayan rock including the Dumri Formation. Note the location of the surface denudation point far south of original location. (e) Growth of the Lesser Himalayan Duplex. Overburden consisting of the Ramgarh Thrust sheet, Greater Himalayan and Tibetan Himalayan rock is passively uplifted and tilted northward. Erosional porthole is located north of the Middle Miocene location. (d) Emplacement of more thrust sheets within the Lesser Himalayan Duplex and formation of the Main Boundary Thrust and Main Frontal Thrust. Note location of erosional porthole of the 'channel', the northward tilt of rock above and below the Main Central Thrust, and presence of the external klippe.

At this time, the surface denudation point migrated rapidly northward, relative to the footwall cut-off of the Ramgarh Thrust (Fig. 3c). As slip was subsequently transferred to the Main Boundary and Main Frontal Thrust systems, the Greater Himalayan surface denudation point remained north of the klippen at the present location of the High Himalaya (Fig. 3d). If channel flow was active from the Late Miocene to the present, this is where the flow would have been e x t r u d e d / e x h u m e d , although ductile fabrics in Greater Himalayan rocks are not present after c. 17 Ma (Searle et aL 2003), which suggests that linked channel flow and extrusion (sensu B e a u m o n t et al. 2001, 2004; Jamieson et al. 2004) was not active at that time. Whereas this possible Miocene migration of the surface denudation point clearly affects models for extrusion of

Greater Himalayan rock, it does not bear directly upon channel flow, assuming that flow of rocks currently exposed ended in the Early Miocene. Therefore, channel flow models that portray a static denudation point (with respect to the base of the channel) throughout the Miocene m a y be unrealistic. Furthermore, surface denudation through a static erosional porthole since the Miocene cannot be invoked as a driver for post-Early Miocene crustal flow.

Detailed structural observations from Lesser Himalayan rocks Over the course of many field seasons in Nepal, we have m a p p e d the contact between Greater and

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Lesser Himalayan rocks in numerous locations (DeCelles et al. 2001; Robinson 2001; Pearson 2002; Robinson et al. 2003, 2006; Pearson & DeCelles 2005). As a result of our work and that of others (e.g. Parrish & Hodges 1996; Kohn et al. 2004; Martin et al. 2005), we view the Main Central Thrust as a relatively discrete structure which places Greater Himalayan rock above Lesser Himalayan rock. We recognize that this definition of the Main Central Thrust contrasts with the popular notion that the fault is a broad, crustal-scale shear zone composed of a tectonic m61ange or an imbricated stack of Greater and Lesser Himalayan rocks (see summary in Hodges 2000). Some workers have suggested that upper and lower strands of the Main Central Thrust can be mapped (e.g. Arita 1983; Catlos et al. 2001); however, for reasons detailed in Pearson & DeCelles (2005), our work suggests that the lower strands are contained entirely within Lesser Himalayan zone rocks, and therefore cannot be part of the Main Central Thrust system (sensu Heim & Gansser 1939). Regardless of how one chooses to define and locate the Main Central Thrust, unequivocal Lesser Himalayan rocks that lie structurally below Greater Himalayan rocks exhibit some important geological relationships that directly relate to the wider debate on channel flow and ductile extrusion/exhumation for the Himalaya. Across Nepal, Greater Himalayan rocks carried by the Main Central Thrust always rest upon lowermost Lesser Himalayan rocks of Early Proterozoic age (Fig. 4c; Pearson & DeCelles 2005), which include the Kushma and Ranimata formations and Ulleri Augen Gneiss, or their lateral equivalents (dark grey unit in Figs lb, 2a & 3). In the Kathmandu area of central Nepal, the lateral equivalents are the Robang Formation and associated Dunga Quartzite beds of St6cklin (1980). The Kushma Formation is a quartzite with local muscovite partings (Fig. 4a), and the Ranimata Formation is a chloritic phyllite/schist (Fig. 4b). More detailed lithological descriptions are available in Upreti (1996) and DeCelles et al. (2001). The Kushma and Ranimata formations are carried by the Ramgarh Thrust, which places them above younger Lesser Himalayan rocks (DeCelles et al. 2001; Robinson 2001; Pearson 2002; Robinson et al. 2003, 2006; Pearson & DeCelles 2005). The Rarngarh Thrust sheet is mapped across Nepal; it is present beneath the High Himalaya, on both the north and south flanks of the external klippen, and in the proximal hanging wall of the Main Boundary Thrust (Fig. 4c; Pearson & DeCelles 2005). As previously mentioned, our observations across Nepal suggest the contact between Greater and Lesser Himalayan rocks is relatively discrete. This is supported by recent studies that used both

whole-rock Nd isotopes and U - P b ages of detrital zircons to explicitly recognize Greater and Lesser Himalayan rocks; these studies involve transects across the Main Central Thrust in the Langtang area of central Nepal (Parrish & Hodges 1996; Pearson 2002), near Nanga Parbat in Pakistan (Whittington et al. 1999; Argles et al. 2003), in Kumaon India (Ahmad et al. 2000) and in the Annapurna area of central Nepal (Martin et al. 2005). In these transects, there is no evidence to suggest that extensive mixing of Greater and Lesser Himalayan rocks occurred. Similarly, we are unaware of any locations where Greater Himalayan rocks lie structurally below Lesser Himalayan rocks. These fundamental observations preclude channel flow and extrusion/exhumation models where Lesser Himalayan rocks become incorporated into the channel during flow or mixed with Greater Himalayan rocks during extrusion (e.g. model HT 111, Jamieson et al. 2006). Our observation that the Main Central Thrust is a relatively discrete ductile shear zone is consistent with studies of other large displacement shear zones, including the S/irv thrust in Norway (Gilotti & Kumpulainen 1986), the Moine thrust in Scotland (Law et al. 1986; Law 1998) and the Willard thrust in Utah, western United States (Yonkee 1997). This geological relationship is also supported by other evidence from Lesser Himalayan rocks in the proximal footwall of the Main Central Thrust. First and foremost, in all mapped locations of the Ramgarh Thrust, Lesser Himalayan rocks carried by the fault are not overturned and are not transposed (Pearson & DeCelles 2005). Primary sedimentary structures, including planar and trough cross-beds, are commonly preserved in the Kushma Formation (Fig. 4a) and in quartzose lithologies within the Ranimata Formation, and attest to the units being right-way-up. The contact between the Kushma and Ranimata formations is sharp and conformable (Fig. 4c; Pearson & DeCelles 2005), and does not appear to be obscured by transposition. Additionally, with a few minor exceptions, we do not observe rootless isoclinal folds in the Ramgarh Thrust sheet. Based upon microstructural observations, Martin et al. (2005) conclude that in the Annapurna area of central Nepal, Main Central Thrust-related deformation of Lesser Himalyan rocks extends only < 1 km below the fault. These geological observations of Lesser Himalayan rocks from the proximal footwall of the Main Central Thrust stand in stark contrast to those developed in the numerical models (Beaumont et al. 2001, 2004; Jamieson et al. 2004, 2006), where Lesser Himalayan rocks are overturned (e.g. Model HT1, Jamieson et al. 2006) and strained to the extent that stratigraphy becomes incoherent. It

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Fig. 4.

(a) The Kushma Formation; bedding is outlined by dashed white lines, cross-bedding outlined by solid white lines, large bed is c. 1 m thick. (b) The Ranimata Formation, phyllite with quartz angen; pen is c. 15 cm long. (c) Surface appearance and conformity of dip of rock units above and below the MCT in western Nepal near the village of Dhuli. Above are the resistant Greater Himalayan rocks (GH) and below are the less resistant phyllites of the Ranimata Formation (Rn) and the more resistant quartzites of the Kushma Formation (K). View looks toward the NW.

may be easy to dismiss these observations as irrelevant, as they reflect later, more brittle deformation, and not channel flow. However, the tight integration between channel flow and ductile extrusion/ exhumation in models for the Himalaya suggest that our understanding of channel flow (sensu stricto) may be enhanced by more accurate modelling of the emplacement of Greater Himalayan rocks.

Summary Models for flow and extrusion/exhumation of Greater Himalayan rocks currently reported in the literature accurately portray many of the firstorder geological relationships present in the orogen. However, careful examination of key relationships in the footwall of the Main Central Thrust and in the external klippen, combined with constraints on the timing of Greater Himalayan

rock extrusion, suggest that significant modifications are needed in current channel flow and extrusion models. In particular, future models need to account for the following six observations. (1)

(2)

(3)

The external klippen contain Greater Himalayan rocks. Primary stratigraphic relationships and sedimentary structures are commonly preserved in these rocks, and they do not appear to have flowed under mid-crustal conditions. The contact between Greater and Tibetan Himalayan rocks in the Kathmandu Klippe is a disconformity (Gehrels et al. 2006). No large normal faults or normal-sense shear zones akin to the South Tibetan Detachment System are described in any of the external klippen in Nepal. Across Nepal, Greater Himalayan rocks carried by the Main Central Thrust always

264

(4)

(5)

(6)

D.M. ROBINSON & O. N. PEARSON rest structurally above Lesser Himalayan rocks carried by the Ramgarh Thrust. Movement on the Ramgarh Thrust began in the Middle Miocene (Kohn et al. 2004; Pearson & DeCelles 2005); therefore, emplacement of Greater Himalayan rocks currently exposed in Nepal occurred prior to the Middle Miocene. Throughout the Miocene, the surface denudation point has not remained in a static location, with respect to the base of a potential channel. Therefore, surface denudation through a static erosional porthole cannot be invoked as a driver for post-Early Miocene crustal flow. Isotopic studies combined with field mapping suggest that the Main Central Thrust can be viewed as a relatively discrete ductile shear zone that juxtaposes Greater and Lesser Himalayan rocks. There is no evidence for a broad zone of mixing, and Greater Himalayan rocks do not lie structurally below Lesser Himalayan rocks. This observation precludes models where Lesser Himalayan rocks become mixed with Greater Himalayan rocks in channel flow and/or subsequent extrusion/ exhumation. In the proximal footwall of the Main Central Thrust, Lesser Himalayan rocks carried by the Ramgarh Thrust are stratigraphically coherent and upright. Models must be modified to account for this fact.

In our view, the observations listed above raise doubts about the applicability of channel flow models, as currently reported in the literature, for Nepal. This does not preclude channel flow and ductile extrusion of Greater Himalayan rocks from playing an important role in formation of the Himalaya and Tibet. However, our data suggest that since the early Miocene, the Himalayan FoldThrust Belt has behaved in a manner consistent with that documented for other fold-thrust belts, and that our current field-based observations are more easily reconciled with a conventional critical taper model. The observations we discuss in this paper will aid the development of future models for channel flow and extrusion of Greater Himalayan rocks, and will contribute to a greater understanding of the geodynamic evolution of the orogen.

Partial funding for this research was provided by The University of Alabama, Department of Geological Sciences and College of Arts and Sciences. We thank P. DeCelles for his editorial assistance. We are grateful to the editors, and for constructive reviews by D. Brown, L. Godin, R. Law, M. Searle and an anonymous reviewer that helped improve this manuscript.

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in Langtang gneisses: implications for chronology of the Himalayan metamorphic core. Journal of Metamorphic Petrology, 23, 399-406. LAW, R. D. 1998. Quartz mylonites from the Moine thrust zone at the Stack of Glencoul, Northwest Scotland. In: SNOKE, A. W., TULLIS, J. & TODD, V. R. (eds) Fault-Related Rocks: A Photographic Atlas. Princeton University Press, Princeton, New Jersey, 490-493. LAW, R. D., CASEY, M. & KNIPE, R. J. 1986. Kinematic and tectonic significance of microstructures and crystallographic fabrics within quartz mylonites from the Assynt and Eriboll regions of the Moine thrust zone, NW Scotland. Transactions of the Royal Society of Edinburgh: Earth Sciences, 77, 99-126. MARTIN, A. J., DECELLES, P. G., GEHRELS, P. G., PATCHETT, P. J. & ISACHSEN, C. 2005. Isotopic and structural constraints on the location of the Main Central Thrust in the Annapurna Range, central Nepal Himalaya. Geological Society of America Bulletin, 117, 926-944. MATTE, P., MATTAUER, M., OLIVET, J. M. & GRIOT, D. A. 1997. Continental subduction beneath Tibet and the Himalayan orogeny: a review. Terra Nova, 9, 264-270. MOLNAR, P., ENGLAND, P. & MARTINOD, J. 1993. Mantle dynamics, uplift of the Tibetan Plateau, and the Indian Monsoon. Review of Geophysics, 31, 357-396. NAJMAN, Y. 2006. The detrital record of orogenesis: a review of approaches and techniques used in the Himalayan sedimentary basins. Earth Science Reviews, 74, 1-72. NAJMAN, Y., CARTER, A., OLIVER, G. & GARZANTI, E. 2005. Provenance of Eocene foreland basin sediments, Nepal: Constraints to the timing and diachroneity of early Himalayan orogenesis. Geology, 33, 309-312. NELSON, K. D., ZHAO, W., BROWN, L. D. ETAL. 1996. Partially molten middle crust beneath southern Tibet: A synthesis of Project INDEPTH results. Science, 274, 1684-1688. 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. PEARSON, 0. N. 2002. Structural evolution of the central Nepal fold-thrust belt and regional tectonic and structural significance of the Ramgarh Thrust. PhD thesis, University of Arizona. 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, TC4008. DOI: 10.1029/ 2003TC001617. RAI, S. M., GU1LLOT, S., LE FORT, P. & UPRET1, B. N. 1998. Pressure-temperature evolution in the Kathmandu and Gosainkund regions, central Nepal. Journal of Asian Earth Sciences, 107, 1381-1397. ROBINSON, D. M. 2001. Structural and Nd-isotopic evidence for the tectonic evolution of the Himalayan

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Locking of southward extrusion in favour of rapid crustal-scale buckling of the Greater Himalayan sequence, Nar valley, central Nepal L. G O D I N 1, T. P. G L E E S O N 2, M. P. S E A R L E 3, T. D. U L L R I C H 4 & R. R. P A R R I S H 5

aDepartment of Geological Sciences and Geological Engineering, Queen's University, Kingston, Ontario, K7L 3N6 Canada (e-mail: godin@ geol.queensu.ca) 2Department of Civil Engineering, Queen's University, Kingston, Ontario, K7L 3N6, Canada 3Department of Earth Sciences, Oxford University, Parks Road, Oxford, OX1 3PR, UK 4Pacific Centre for Isotopic and Geochemical Research, Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, British Columbia, V6T 1Z4, Canada 5NERC Isotope Geosciences Laboratory, Keyworth, Nottingham, NG12 5GG, UK Abstract: The South Tibetan detachment system (STDS) bounds the upper limit of the Greater

Himalayan sequence (GHS), which consists of the exhumed middle crust of the Himalaya. In the Annapurna range of central Nepal, the GHS comprises a sequence of amphibolite-grade augen gneisses with a 3.5 km thick leucogranite at the higher structural levels (Manaslu granite). Two major low-angle normal-sense shear zones have been mapped. The Chame detachment has similar grade metamorphic rocks above and below and is interpreted as a ductile shear zone wholly within the GHS. The Phu detachment is a ductile-brittle normal fault which wraps around the top of the Manaslu leucogranite and defines the uppermost, youngest strand of the STDS, placing folded unmetamorphosed Palaeozoic rocks of the Tethyan sedimentary sequence above the GHS. Our data indicate that ductile flow and southward extrusion of the GHS terminated with cessation of movement on the brittle upper strand of the Phu detachment at c. 19 Ma, which was followed almost immediately by crustal-scale buckling. Argon thermochronology reveals that the bulk of the metamorphic rocks and lower portions of the Tethyan sedimentary sequence in the Nat valley cooled through the hornblende, biotite and muscovite closure temperatures at c. 16 Ma, suggesting very rapid cooling rates. The thermochronology results indicate that this cooling occurred 2-3 million years earlier than in the frontal part of the extruded GHS. Although the extrusion in the frontal part of the GHS must have locked at the same time as in the Nar valley, the exhumation there was slower, and most probably only assisted by erosion, rather than by rapid folding as is the case in the Nar valley. This buckling indicates a step northward in deformation within the Himalayan belt, which may be a response to a lower deforming taper geometry in the foreland.

Since the first orogen-parallel synconvergent normal faults were recognized in the Himalayas (Caby et al. 1983; Burg, 1983; Burchfiel & Royden 1985), there have been numerous studies attempting to understand the mechanics of normal faults which are synchronous and parallel with structurally lower thrust faults (e.g. Hodges et al. 1992; Searle et al. 1997, 2003; Vannay & Grasemann 2001). In the Miocene, two major parallel yet opposite-sense shear zones, namely the Main Central thrust (MCT), and the South Tibetan detachment system (STDS), were active and broadly coeval (Hubbard & Harrison 1989; Searle & Rex 1989; Hodges et al. 1992, 1996; Godin et al. 2006). However, both the MCT and

STDS zones display complex multiple strands that might have operated at different times and under different mechanical conditions (plastic to brittle). For example, the MCT zone is often subdivided into a younger, structurally lower brittle thrust, and an older, structurally higher ductile shear zone (e.g. Arita 1983; review in Godin et al. 2006). The present-day position of the Greater Himalayan sequence, the exhumed metamorphic core of the Himalaya, is interpreted to have resulted from predominantly ductile extrusion of midcrustal material between these two opposite-sense shear zones. Several extrusion mechanisms have been proposed, ranging from wedge-type geometries (Grujic et al. 1996; Grasemann et al. 1999),

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, 269-292. 0305-8719/06/$15.00

9 The Geological Society of London 2006.

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to general shear distributed throughout the Greater Himalayan sequence (Vannay & Grasemann 2001), to channel flow within the core of the Greater Himalayan sequence (Beaumont et al. 2001, 2004; Grujic et al. 2002). The STDS marks the structural and metamorphic upper limit of the Greater Himalayan sequence, separating 'unmetamorphosed to weakly metamorphosed [folded] strata of the hanging wall (Tethyan sedimentary sequence) from upperamphibolite facies gneisses and leucogranites of the Greater Himalayan sequence footwall' (Hodges 2000, p. 338). It is a system of ductile shear zones and brittle normal faults recording a complex strain history. In various regions of the Himalayan orogen, it has been proposed that the STDS may have had an early thrust component (Weismayr & Grasemann 2002), overprinted by several phases of normal- and thrust-sense motion (Godin et al. 1999), and possibly also by dextral strike-slip motion (P~cher 1991). More recently, Yin (2003) suggested that the STDS may represent a passive roof thrust within the MCT system. Certainly, there is general agreement that the fault system acted as a major ductile top-to-thenorth normal-sense fault in the Miocene, coeval to the structurally lower MCT (Vannay & Hodges 1996; Godin et al. 2001; Godin et al. 2006), although some would argue that there is no evidence for synchronous motion (Murphy & Harrison 1999). The portion of the STDS history coupled with the MCT, that we term the MCT-STDS coupling period, has led many to believe that the metamorphic core in the High Himalaya was extruded southward from its mid-crustal depths between these two kinematically and dynamically linked opposite-sense faults (Burchfiel & Royden 1985; Hodges et al. 1992; Royden 1996; Grujic et al. 1996, 2002; Beaumont et al. 2001). Both fault systems may have had, prior to or after the extrusive phase of the Greater Himalayan sequence, periods of independent motion, such as renewed late brittle (normal and/or thrust) reactivation, which are not associated with extrusion of the Greater Himalayan sequence. Structural, physiographic and geochronological data suggest segments of the MCT and STDS may have been active after the MCT-STDS coupling period (Harrison et al. 1997; Hurtado et al. 2001; Catlos et al. 2002). Nevertheless, although both structures may have been reactivated in recent times, these reactivations may be local and sporadic, without a dynamic coeval link between the two fault systems. Central Nepal (Fig. 1a) is an ideal region to study the metamorphic core and its bounding structures. The area has a long history of geological investigations (Bordet et al. 1971; P~cher 1975;

Colchen et al. 1981). Furthermore, the Greater Himalayan sequence in the Annapurna range is exposed in two contrasting structural settings: in the well-studied 'frontal' north-dipping homoclinal slab (Le Fort et al. 1986; Hodges et al. 1996; Vannay & Hodges 1996; Godin et al. 2001), and in the structural windows provided by apparent domes north of the Annapurna range (Border et al. 1975; Le Fort et al. 1999; Godin 2001; Gleeson & Godin 2006). These two contrasting, yet spatially close, settings provide the opportunity to study and compare exhumation processes at various structural positions within the Greater Himalayan sequence. The objectives of this paper are to present structural and geochronological constraints on the timing of cooling of the folded Greater Himalayan sequence exposed north of the Annapurna range in the Nar valley, and to compare the geometry and cooling ages from the Nar valley with previously published results from the Greater Himalayan sequence of the frontal homoclinal segment in the Kali Gandaki valley (e.g. Vannay & Hodges 1996; Godin et al. 2001). The ultimate goal of the paper is to document the geometry and timing of cessation of the MCT-STDS coupling, and make suggestions as to why this might have happened.

Geology of north-central Nepal The geology of north-central Nepal (Fig. l a) is defined by three main lithotectonic units broadly striking east-west. The northernmost unit is the Tethyan sedimentary sequence, consisting of a c. 12 km thick passive margin succession deposited on the northern margin of the pataeo-Indian continent. In central Nepal, these sedimentary rocks range in age from Ordovician to Lower Cretaceous (Bordet et al. 1975; Colchen et al. 1981; Gradstein et al. 1992; Garzanti 1999). The lowermost Tethyan sedimentary sequence is metamorphosed to zeolite-lowest greenschist grade with a foliation typically defined by muscovite. The metamorphic grade decreases upwards to the epizone-anchizone boundary (Garzanti et al. 1994). The Tethyan sedimentary sequence is in normal fault contact with the underlying Greater Himalayan sequence along the Miocene-age STDS (Burchfiel et al. 1992). In central Nepal, the STDS includes the structurally lower ductile level of the Annapurna detachment zone (Brown & Nazarchuk 1993; Godin et al. 1999), which correlates with the Deurali detachment in the Annapurna Sanctuary (Hodges et al. 1996), and with the Chame detachment in the Marsyandi valley (Coleman 1996). The structurally higher

BUCKLING OF THE GREATER HIMALAYAN SEQUENCE

271

Fig. 1. (a) Regional geology map modified from Colchen et al. (1981) and Searle & Godin (2003). Inset shows the study area within the Himalayan orogen. Greater Himalayan sequence in grey; important structures noted are Main Frontal thrust (MPT), Main Boundary thrust (MBT), Main Central thrust (MCT), and South Tibetan detachment system (STDS). The various strands of the STDS are indicated by numbers. 1, Annapurna detachment; 2, Deurali detachment; 3, Macchapuchare detachment; 4, Chame detachment; 5, Phu detachment. Section line 6, indicates the study area of Vannay & Hodges (1996) and Godin et aL (2001) in the Kali Gandaki valley. (b) Map of Manang region, Marsyandi valley, showing area mapped in detail (box outline) and axial surface traces of the fold system affecting the Greater Himalayan sequence. brittle level of the Annapurna detachment coincides with the Machapuchare detachment in the Annapurna Sanctuary (Hodges e t al. 1996), and with the Phu detachment in the M a r s y a n d i - N a r

region (Searle & Godin 2003). Most of the surface trace of the STDS is exposed south of the Annapurna range (Fig. la). However, just east of Annapurna II, the fault system takes on a northerly

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trend, wraps around the Miocene Manaslu leucogranite, and then resumes its easterly course (Fig. la, b). This change in orientation is due to a series of upright, shallow, west-plunging folds which, combined with topography, account for these two apparent orthogonal bends (Bordet et al. 1975; Searle & Godin 2003). In the Marsyandi valley, the STDS consists of two detachment faults (Fig. lb): the lower ductile Chame detachment (Coleman 1996), and the upper ductile-brittle Phu detachment (Searle & Godin 2003). The Greater Himalayan sequence, representing the exhumed metamorphic core of the Himalaya, is located in the footwall of the STDS. It consists of amphibolite-facies Proterozoic gneisses intensely affected by Oligocene and Miocene thermal events (Coleman 1996; Vannay & Hodges 1996; Godin et al. 2001). The structurally lowest part of the Greater Himalayan sequence is in thrust contact with the Lesser Himalayan sequence along the north-dipping, south-vergent Miocene-age MCT zone (Bouchez & P~cher 1976; Brunel 1986). The metasedimentary rocks in the immediate hanging wall of the Chame detachment were initially assigned to the Cambro-Ordovician levels of the Tethyan sedimentary sequence (Bordet et al. 1971, 1975; Colchen et al. 1981, 1986). Although this stratigraphic correlation may prove to be correct, their metamorphic degree (amphibolite facies) and highly sheared structural style suggest these rocks share a common tectonic history with the Greater Himalayan sequence (Searle & Godin 2003; Gleeson & Godin 2006).

Geology of the Nar valley Greater

Himalayan

sequence

The Greater Himalayan sequence is typically described in terms of three distinct units, Unit I, Unit II and Unit III (Formations I, II, III of Bordet et al. 1971). The structurally lowest unit (Unit I), located in the hanging wall of the MCT, is dominated by a pelitic kyanite-sillimanite-garnetbiotite schist. The central and upper parts of the Greater Himalayan sequence are dominated by a hornblende-biotite gneiss (Unit II), 'interlayered' with a distinctive Lower Ordovician granitic augen gneiss (Unit III; Bordet et al. 1975; Colchen et al. 1981). Recent work in the upper part of the Greater Himalayan sequence in the Nar valley reveals a more complex lithological sequence with significant previously unrecognized marble and pelite layers, metamorphosed to amphibolite facies (Searle & Godin 2003; Gleeson & Godin 2006). In the Nar valley, the Greater Himalayan sequence is separated into two distinct structural

levels. The structurally lower level (Level 1), cropping out in the core of the Chako antiform along the Phu Khola (Figs 2 & 3), comprises hornblendebiotite gneiss (calc-silicate gneiss of Bordet et al. 1971), and a thin (50m thick) sheet of highly strained biotite schist (schiste gt p l a q u e t t e of Bordet et al. 1971). Locally, this unit contains lenses of granitic augen gneiss that are interpreted to represent lower strained equivalents of the biotite schist. These two lower level units are correlated with Unit II and Unit III, exposed further south in the Marsyandi valley (Searle & Godin 2003; Gleeson & Godin 2006). The rocks of the upper level (Level 2) consist of a biotite-bearing marble (Unit IV) and a garnetbearing pelitic schist (Unit V). Petrography and garnet-biotite thermometry of the pelitic schist indicate peak metamorphic temperatures of 500-650~ consistent with amphibolite-facies metamorphism (Gleeson & Godin 2006). These two upper level units are interpreted as previously undescribed components of the Greater Himalayan sequence, found structurally above the Chame detachment (Searle & Godin 2003; Gleeson & Godin 2006). The Manaslu leucogranite forms in its eastern portion a 3 km thick sill of tourmaline-muscovite leucogranite located along the top of the Greater Himalayan sequence beneath the Phu detachment (Searle & Godin 2003). It is a peraluminous granite with extremely high Sr initial ratios, and is interpreted as a crustal minimum-melt granite derived from a protolith source similar to the sillimanite gneiss of Unit 1 (Le Fort 1981; Vidal et al. 1982). Earlier workers described the Manaslu leucogranite as intruding across the STDS into unmetamorphosed sediments of the Tethyan sedimentary sequence (e.g. Le Fort 1975, 1981" Colchen et al. 1986; Guillot et al. 1993). However, pressure-temperature conditions of rocks along the Nar valley and the upper Marysandi valley show that the carbonate rocks in which the leucogranite was intruded are at high metamorphic grade and do not exhibit contact metamorphic relations (Schneider & Masch 1993). Searle & Godin (2003) mapped a new upper strand of the STDS, called the Phu detachment, wrapping around the top of the Manaslu leucogranite; they place the Manaslu leucogranite wholly in the Greater Himalayan sequence, in common with all other Himalayan leucogranites. The Th-Pb ion microprobe ages of monazites from the Manaslu leucogranite define two major pulses of magmatism at 22.9 _+ 0.6 Ma and 19.3 +__0.3 Ma (Harrison et al. 1999). Structural studies along the Pangre glacier show that the Phu detachment cuts the contact of the Manaslu granite with the overlying Greater Himalayan sequence rocks (Fig. 2),

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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

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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

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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

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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.

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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.

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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

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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

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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).

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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).

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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

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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

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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

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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.

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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.

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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.

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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 (<30%) and the limited thickness of the melted portion of the crust (5-10 km), both demand that lateral draining and collection occur in order to produce lenses several kilometres in size. We will come back to this aspect later. Shearing

To estimate the potential role of shearing (simple shear) on magma segregation in the GHS, we use the scaling laws of Rabinowicz & Vigneresse (2004). These authors derived the set of equations that govern melt motion within a matrix deformed by simple shear. What is calculated below primarily concerns melt segregation within the molten crust (GHS). The geometrical and timescale characteristics we derive are for in-situ melts segregated into veins within a shear zone. These veins are not to be confused with dykes through which melt ascends towards the upper levels of the crust, although thick veins may eventually develop in this way due to the buoyancy of the magma they contain. The basic equations that allow us to calculate magma segregation values are as follows. The compaction length L (m) is now defined as (Stevenson 1989): L---- ('rlsko/'rlmdi~o) 1/2

(7)

The compaction velocity Va (m s -1) is given by: Vd =- ( arlgko ) / rlm

(8)

The scaling for time ~-is given by: 'r = L@o/Vd

(9)

Note that the parameters L and ~-, although identical in their physical meaning to those defined by McKenzie (1985), differ in that they consider reaction progress via the porosity term ~o. Thus the calculated numbers are lower than those obtained from McKenzie (1985). In addition, two dimensionless numbers are introduced. The first, called the velocity number Nv, is calculated as the ratio between the shear and compaction velocities: Nv ----coLCbo/2Vd

(10)

The other, termed the shear number N~, is the ratio between the stress and buoyancy gradients: Ns = ~qsCo~o/Lt~qg

(11)

in which the term Co is the shear velocity, calculated as: Co = V / h

(12)

where h is the thickness of the shear zone (m) and V is its tangential velocity (m s-l). In this work, unless otherwise stated, we used a constant value for V of 2 cm a- ~, but some calculations were performed with V = 5 cm a-~ which corresponds to the present-day Himalayan convergence rate (Bilham et al. 1997). Rabinowicz & Vigneresse (2004) performed a series of numerical simulations to evaluate the relative role of shearing versus compaction in a constantly sheared two-phase system. The results suggest that the characteristic time T~ of melt segregation into veins due to simple shear is: T~ = 6"r/N~

(13)

while that corresponding to compaction during shearing, To, is given by: Tc = 50r

(14)

In addition, their numerical experiments show that the spacing of veins is close to the compaction length scale L, while their width is equal to q%L. We have calculated the characteristic times for melt channelling for shear zone thicknesses varying from 0.1 up to 10 km (Fig. 5). For these calculations we have used a constant melt viscosity of 104.5 Pa s, and a grain size of 4 mm. We have considered melt fractions in the source of 0.1 and 0.2. This gives compaction lengths of 80 and 160 m respectively. Under these conditions the

MAGMATIC PROCESSES IN THE HIMALAYA

~4

-

-

_o 0 0.1

i 1

i 10

i 100

i 1000

10000

S h e a r zone t h i c k n e s s ( m )

Fig. 5. Relation between characteristic timescales needed to segregate anatectic melt into veins (via shearing (T~) and compaction (To)) and shear zone thickness, using the scaling equations of Rabinowicz & Vigneresse (2004). Calculations have been performed using a melt viscosity of 10 4.5 Pa s, a grain size of 4 ram, a shear rate of 2 cm a- 1 and melt fractions f of 0.1 and 0.2. See text for details on assumptions and calculations.

characteristic time of magma compaction during shearing (Tc) ranges from 250 to 500 ka (Fig. 5). Figure 5 shows the variations of the characteristic time of melt channelling (T~) with shear zone thickness. It can be seen that shear zones thicker than 1 km need at least 1 million years to segregate melt into veins, the results not being very sensitive to melt fraction. Shear zone thicknesses for which Ts > To imply that compaction is a more efficient mechanism than shearing for segregating melt, and occurs for shear zone thicknesses higher than 1 0 0 m (Fig. 5). For example, a 10 km thick crustal shear zone, which would be equivalent to the assumed thickness of the main part of GHS if sheared homogeneously during crustal convergence, would segregate melt into dykes in more than 10 million years, which is similar to the accepted time interval of active magmatism in the central part of the Himalaya, and much longer than To. On this basis we conclude that shear-assisted melt segregation in the GHS requires that the thickness of active shear zones at any given time be significantly smaller than the full thickness of GHS. Shear zones having thicknesses in the range 1 0 - 1 0 0 m are able to collect melt over a timescale range of 10 to 100 ka. Such timescales are comparable to estimates of magma residence time at source based on trace element behaviour (Ayres et al. 1997; Harris et al. 2000). For such shear zone thicknesses, veins filled by melt will have a vertical extent similar to the thickness of the shear zone, and a width varying between 8 and 32 m. Such vein thicknesses are comparable to those measured in the field (e.g. Inger & Harris 1993; Scaillet et al. 1996). These calculations suggest therefore that extraction of melt via simple

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shear could occur with a recurrence time of 10100 ka. The Manaslu leucogranite body is at least 5 km thick and was emplaced during the time interval 2 4 - 1 9 Ma (Harrison et al. 1999). Assuming discontinuous but regular magma supply via sills 100 m thick, as suggested by isotopic constraints and thermal modelling (Deniel et al. 1987; Annen et al. 2006), this gives an average time recurrence for magma injection of c. 100 ka. For 20 m thick sills, the time interval between two magma supply periods drops to 20 ka. This broadly fits with previous estimates (see also Cruden & McCaffrey (2001) for a general discussion on this aspect). It must be noted, however, that the above figures can be affected if we adopt a different set of input parameters. For instance, we have performed a maximum calculation in which all critical parameters are tuned to speed up melt collection. This requires using the lowest permissible viscosity for HHL melts (104 Pa s at 820~ on the verge of muscovite thermal stability), an average grain size of 4 mm, a rate of shearing of 5 cm a-1 and a melt fraction of 0.2 (0.3 is the theoretical maximum but not all the melt will be drained out towards the vein). Under those conditions the characteristic time of melt segregation into veins is less than 1 million years, even for a shear zone 1 km thick. The process generates kilometre-long dykes whose maximum likely thickness would be c. 600 m. In view of the high rate of shear, such thick dykes will be deflected towards the shear plane. Such flat-lying magma dyke-sheets, if indeed present in the Himalaya, thus probably correspond to particularly low-viscosity HHL melts, possibly the hottest, although we stress that the change in viscosity relative to the standard value adopted here remains small, only half a log unit lower. The important point is that using field- and laboratory-constrained input parameters relevant to the Himalayan context, the calculations show that shear-assisted melt extraction is: (1) conceivable in the GHS; and (2) occurs on timescales comparable to those obtained from independent methods. One possible reason for the active shear zone thickness being comparatively much smaller than the bulk GHS is that heat propagates upwards across the metamorphic pile, so that the melting wave also moves upwards. We suggest that whenever the partial melt zone reaches a critical thickness of 1 0 - 1 0 0 m, it becomes a favourable locus for melt segregation via shearing. Shearing and concomitant melt segregation possibly occur at a smaller scale, yet the segregated melt does not escape far from its source. Conversely, as soon as the sheared partial melt zones become 10-100 m thick, the veins generated by shearing may have the potential to self-propagate upwards, eventually feeding laccoliths (e.g. Petford 1995).

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Final stage of HHL building The above lines of evidence show that compaction and shear-assisted melt segregation are both viable mechanisms, yet each cannot account on its own for the building of laccoliths whose thicknesses largely exceed 1 kin. To be efficient in generating sheets several hundred metres thick, compaction requires that a substantial proportion of the GHS be melted at any single time. Our calculations suggest thicknesses in the range of 5 to 10 kin. We note that such thicknesses of melted crust are comparable to those inferred to exist presently beneath southern Tibet (e.g. Nelson et al. 1996; Li et al. 2003) and it is thus conceivable that they were also produced during Miocene times (e.g. Searle et al. 2003; Gaillard et al. 2004). Numerical simulations (Beaumont et al. 2004; Jamieson et al. 2004) also suggest that the process is a viable one. Similarly, shear-controlled extraction cannot produce sills much thicker than 1 km and, assuming that it does, such a process will be faced with the lack of convection, a constraint needed to produce isotopically heterogeneous magma bodies at the hundred-metre scale. Simple fluid dynamical calculations show that a sill filled by leucogranite melt will vigorously convect once its thickness exceeds a few metres, which will lead to homogenization of the melt, a feature not supported by the strong isotopic heterogeneity of HHL (see next section and Annen et al. 2006). Therefore, we conclude that the production of large-scale HHL bodies, such as the Manaslu pluton in Nepal, or the Monlakarchung-Pasalum pluton in Bhutan, requires an additional step of magma accumulation beyond that provided by compaction and shearing mechanisms alone. In this respect, such HHL bodies may not reflect the conditions of magma extraction corresponding to the emplacement of sills with thicknesses in the range 1-1000 m. Although the exact proportion of numerous pods and sheets of this size intruding the GHS is difficult to quantify exactly, they may represent a considerable, if not dominant, portion of melted crust within the whole Himalaya. For obvious reasons, geological studies of HHL have focused on the thickest bodies cropping out in the range. At this point, we recognize that the processes leading to emplacement of the thickest HHL still remain obscure or, at the least, not well characterized. Building of large plutons or laccoliths requires crustal melt draining over a volume or surface one order of magnitude higher than the pluton itself (e.g. Brown & Solar 1998b; Cruden & MacCaffrey 2001). Only for the Manaslu (Harrison et al. 1999), Shisha Pangma (Searle et al. 1997) and EverestRongbuk (Murphy & Harrison 1999; Simpson et al. 2000; Searle et al. 2003) leucogranites is the

time interval of laccolith growth known with some precision and indicates that the average melt input rate is not significantly higher than that required to produce smaller sills or dykes. The lack of conspicuous magmatic fabric in these large plutons suggests that they did not grow within a stress field characterized by a strong shear component. Thus the thickest HHL may correspond to local zones of the Himalaya range where compaction-driven extraction (followed by dykes, as suggested in Fig. 1), predominates over the shear-controlled melt extraction mechanism (i.e. zones for which T, >> To), perhaps also reflecting a local enhanced source fertility relative to other parts of the range. The less competitive character of shearing relative to compaction may correspond to the requirement that shearing is distributed evenly across the entire GHS metamorphic pile (Fig. 5). Alternatively, it may also correspond to a lower rate of shearing. For instance, with a shear velocity of 0.5 cm a - l , instead of 2 cm a -1, any shear zone thicker than 10 m will be less efficient than compaction in extracting melt. We therefore suggest that thick HHL may have been fed from a melt layer on top of a migmatite source unit that was generated by compaction (Fig. 1). This layer was subsequently drained by sills and dykes that carried leucosome melts out of the melting zone. For an average thickness hc --- 100-300 m of the compacted layer, the radial zone z required to be withdrawn to build up a laccolith of volume V1 can be calculated using the following equation, which assumes a circular geometry of the tapped area: z = 2(V~/(hcI])) 1/2

(15)

Given volumes of 2000-3000 km 3 as estimated for the largest HHL plutons (e.g. Badrinath, Manaslu, Everest, Monlakarchung Pasalum), this corresponds to a lateral distance of melt transport of 50-100 km (Fig. 2). This distance broadly corresponds to the average spacing observed between the largest laccoliths, at least in the central part of the Himalayan range. Our model is similar to that proposed by Cruden & McCaffrey (2001 ), except that it considers the existence of an intermediate melt layer. The causes of dyke initiation may be intrinsic to evolution of the melt source layer, such as a local increase of its internal pressure due to volatile oversaturation during crystallization. Alternatively, dyke initiation could be triggered by a change in the regional stress field. A crustal ramp could produce extension in the hanging wall of the entire melted layer thus favouring initiation of a dyking mechanism in the hanging wall to the ramp (Fig. 1). Note that in view of the laminar

MAGMATIC PROCESSES IN THE HIMALAYA regime calculated for magma flow within dykes (Scaillet et al. 1996), repeated tapping of a compacted melt source layer having a thickness of 100-200 m will still preserve the isotopic heterogeneity of the source. If sills generated by compaction provide a reservoir to feed overlying thick laccoliths, then an important aspect for such a compaction-dyke model is that the layer must either remain molten to allow dyke propagation, or the sill must periodically be reinflated. To evaluate the time needed for complete solidification of this layer, we have performed numerical simulations of conductive heat loss of a 300-1000 m thick stagnant magma sill emplaced in host rock at various temperatures (see Appendix). It can be seen that the time for solidification strongly increases with host rock temperature (Fig. 6), in agreement with previous findings (Davidson et al. 1992). A 3 0 0 - 1 0 0 0 m thick sill emplaced at 750~ and 15 km depth takes between 1.6 to 2.2 million years to crystallize, which indicates that the melt layer may survive for a significant time after extraction from the migmatite source layer, while sitting on top of it, and before dyke transport upwards. In contrast, it takes only 10 000 years for a 1 km thick sill to fully crystallize if emplaced in cold upper crust initially at 350~ (Fig. 6). Calculations of magma ascent rates performed using observed dyke thicknesses (10-20 m) and inferred HHL viscosities show that laccolith building can be achieved in a few hundred years (Scaillet et al. 1996). The time needed to build a complex leucogranite sheet network like that observed in the Everest area (Searle 2003) is calculated to be on the order of 1000-2000 years, if melt supply is continuous. Therefore, it appears that the integrated time needed for magma segregation in the Himalaya can be significantly less than 1 million years, approaching 100 000 years for a favourable combination of controlling parameters (low melt

i0000000 C= 0

:'~ 2 ~~

~oooooo ~ ioooo I000

~-

1oo 0

200

400

600

800

1000

1200

Thickness o f sill ( m )

Fig. 6; Effect of host rock temperature on the time for full crystallization of leucogranite sills of various thicknesses.

303

viscosity, large grain size, high melt fraction), and being essentially controlled by the compaction and shearing timescales. However, for the large plutons, we stress that not all processes need to operate in one single step. Instead, it may be discontinuous, depending on crustal-scale tectonics. Thus, the age of leucogranite crystallization/ emplacement might not be strictly equivalent to the age of melt initiation, a common assumption made when dating large magma bodies, since a compacted melt layer 100-300 m thick may incubate on top of the migmatite source unit for several million years (Fig. 6). Our calculations show that a difference of 2 - 4 million years may exist between the onset of melting and final emplacement of the leucogranite body at higher levels.

Timescales of magma cooling So far we have ignored one essential aspect of HHL petrogenesis, that is their strong isotopic heterogeneity (e.g. Deniel et al. 1987). The fact that HHL are heterogeneous down to a scale of c. 100 m implies that bodies thicker than a few hundred metres never convected whilst molten. This imposes an important constraint on the input rate of individual melt batches, which should be low enough to allow formerly injected magma batches to crystallize before the next one is intruded (see Annen et al. 2006). This aspect must be reconciled with the recorded cooling patterns of HHL. Dating using several complementary geochronometers indicates that HHL are almost always characterized by fast cooling rates soon after emplacement (e.g. Copeland et al. 1990; Searle et al. 1997, 2003; Harrison et al. 1999). The simulations performed by Jamieson et al. (2004) within the context of the channel flow model were unable to capture the HHL time-temperature path. As pointed out by Jamieson et al. (2004), fast cooling can be achieved via two main mechanisms: either the laccolith is emplaced in host rock sufficiently cold to quench it, or HHL cooling is accelerated by normal faulting on its top. For those HHL emplaced into hot metamorphic host rocks such as sillimanite gneisses (e.g. Searle et al. 1993, 1997, 2003; Searle & Szulc 2005), rapid quenching as a mechanism to explain the rapid cooling path at 20-17 Ma detected in the Bhagirathi, Shisha Pangma and Everest leucogranites can be ruled out, unless the entire package, host sillimanite gneisses and leucogranite sills and dykes, were exhumed rapidly together. Instead, isotopic dating suggests that normal-sense motion of the overlying South Tibetan Detachment (STD) in these areas was active during this time as footwall rocks were extruded rapidly upwards along the footwall of the STD.

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The above simulations show that a 1 km thick HHL emplaced in host rock at 300-350~ instantaneously relaxes to ambient temperature, such that crystallization ages will be only slightly older than cooling ages as recorded by muscovite. Incorporating the lack of convection constraint shows that building a 5 km thick laccolith without convection requires the accumulation of 20-100 m thick sills over a 5 million years time period with the top intrusion lying at 10-12 km depth (Annen et al. 2006) (note that shallower depths of intrusion, which would allow a shorter time period of magma intrusion, are not allowed by the requirement of magmatic muscovite in HHL that demands a minimum pressure depth of c. 3 kbar). This is comparable to the time interval estimated for emplacement of the Manaslu pluton (Harrison et al. 1999). Such an estimate, however, is based on laccolith growth with a static roof. Assuming that the roof of the growing magma body is permanently removed due to normal-sense motion on the overlying STD, then the time interval necessary to build a pluton 5 km thick drops to 1.3 million years. Various other possibilities and combinations were tested by Annen & Scaillet (2006), including minimum depth constraints imposed by the presence of magmatic muscovite, and exhumation associated with normal-sense motion on the STD accelerating cooling. If all sills having thicknesses on the order of 1 km are as isotopically heterogeneous as the thickest bodies, their emplacement either via compaction or via shearing processes into sillimanite-grade gneisses requires normal-sense motion on the overlying STD to be active so as to accelerate cooling and inhibit convection. On the basis of the simulations of Annen et al. (2006) and Annen & Scaillet (2006) we infer that a 1 km thick laccolith may be emplaced in c. 250 ka if intruded coeval with normal-sense motion along the overlying STD.

Some possible consequences of melt extraction for localized crustal flow Our calculations confirm the view that the ascent of granitic magma within continental crust can occur rapidly (Petford et al. 2000; Jackson et al. 2003), even instantaneously for the last, dyke-controlled, part of magma ascent (Scaillet et al. 1996). Besides the physical aspects of magma transfer in continental crust, our model also has implications for the rheological behaviour of the crust and the way it is affected by the presence of melt (e.g. Hollister & Crawford 1986; Davidson et al. 1994). As long as the melt fraction remains trapped within its protolith, it may induce a mechanical weakening of the crust, which has been shown to be critical in

recent models of extrusion tectonics coupled to channel flow (Beaumont et al. 2001, 2004). However, our calculations show that melt draining via compaction/shearing is inescapable even on a short timescale. This will severely change the mechanical response of thickened crust during long-term crustal convergence. Once the lower to middle crust has been drained of its melt fraction, it may become mechanically stronger and thus strain accumulation is likely to migrate towards higher structural levels, localizing where melt is concentrated. Qualitatively this predicts that structurally lower parts of a metamorphic pile may undergo less horizontal transport than structurally higher levels of the metamorphic pile; however, this will be counteracted by the increasingly shorter crystallization times of the magma as it rises to higher structural levels (Fig. 6). Overall, the process of magma migration towards the surface should result in a rapid ( < 1 million years) shift from a plastic to a brittle mechanical response of the deforming crustal block. This may have profound consequences on crustal-scale tectonics or melting events within the whole zone of convergence. In particular, such enbrittlement of the crust could be responsible for the initiation of new thrust faults, which should be preferentially located at the interface between plastic and brittle crust. Our model thus predicts that the width of single crustal blocks bounded by large faults should be similar to the horizontal dimensions of zones tapped to build laccoliths at higher structural levels, that is between 50 and 100 km.

Conclusion Compaction and shearing of migmatite zones are clearly both of importance in melt extraction during crustal convergence, their relative roles being determined by the overall rate of convergence and its variation in time and space. Periods during which, or areas where, shear rates are low or nil will correspond to times of preferred accumulation of melt in the source region via compaction. The crustal levels where melt coalescence has proceeded to some extent will possibly be those levels first drained once imposed shear rates exceed some threshold value. Melting and shearing can trigger each other and are thus in feedback relationships, as proposed by Hollister (1993) and Brown & Solar (1998a). We suggest that the time needed to build a complex leucogranite sheet network like that observed in the Everest area (Searle 2003) is on the order of 1000-2000 years, if melt supply is continuous. The Himalayan range displays in a unique and dramatic way the various steps leading to the assembly of granite bodies in the crust. Does the

MAGMATIC PROCESSES IN THE HIMALAYA Himalaya, however, represent a reference model for granitic m a g m a ascent in the crust that could be valid in older orogens? The great sensitivity of the timescales of vertical m a g m a transfer to various parameters (melt fraction, grain size, melt viscosity, shear rate, existence of a crustal ramp) implies that m a g m a availability on top of a migmatite source unit, and its coalescence in the upper-mid-crust, may be inhibited or promoted to various extents along-strike in a single orogen, or among different orogens. Therefore, before comparing Himalayan m a g m a t i s m to that of other orogens, all factors controlling melt transfer processes should be carefully evaluated. Finally, with respect to the application of the channel flow model to the Himalaya, we suggest that a necessary i m p r o v e m e n t for such models of flow of mid-crustal rocks is to consider the possibility of melt redistribution within the crust, since this may profoundly affect the long-term structural and metamorphic evolution of the orogen. We are grateful to P. Le Fort for many discussions on Himalayan leucogranites, and D. Nelson who inspired comparisons of geophysical evidence for melting beneath Tibet with formation of Himalayan granites. Discussions on physical aspects of magma extraction and emplacement with C. Annen and S. Sparks have been helpful. M.P.S. used NERC grant NER/K/S/2000/951 during this work. The detailed reviews of S. Cruden and M. Edwards are gratefully acknowledged, as well as the thorough and very helpful editorial handling of R. Law.

Appendix We performed one-dimensional numerical simulations of the thermal evolution of leucogranite sills overlying thickened partially molten crust. The initial non-linear geotherm is calculated using a second-order polynomial curve through to temperatures of 0, 350 and 800~ produced at 0, 15 and 30 km depth, respectively. We explored similar initial geotherms by adopting temperatures of 650 and 750~ at 15 km. At t = 0 a leucogranite sill of a given thickness is emplaced at 15 km depth and the subsequent thermal evolution is calculated by solving the one-dimensional equation for conductive heat transfer (Davidson et al. 1992), using an explicit difference method, with a node spacing of 100 m, and a time step of 32 years. The initial temperature of the magma is set to 800~ and the latent heat release is calculated by empirically fitted curves on the relationships between crystallinity and temperature indicated by experimental phase equilibria on Himalayan leucogranites (see Annen et al. 2005). BoundarT conditions were constant temperatures of 0 and 800~ at 0 and 30 km depth, respectively. The latter condition is valid as long as the lower crust remains at a temperature necessary for partial melting and so our results apply for conditions during and soon after granite melt production. Other physical parameters are as described by Harrison et al. (1998).

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Thermal evolution of leucogranites in extensional faults: implications for Miocene denudation rates in the Himalaya C. A N N E N 1 & B. S C A I L L E T 2

1Section des Sciences de la Terre, Universitd de Genkve, 13 rue des Marafchers, 1205 Genkve, Switzerland (e-mail: [email protected]) 2Institut des Sciences de la Terre d'Orldans, UMR 6113 CNRS-UO, 1A rue de la Fdrollerie, 45 071 Orldans Cedex2, France Abstract: The crustally derived High Himalayan leucogranites (HHL) are characterized by strong isotopic heterogeneity and occurrence of magmatic muscovite. Such attributes indicate that the HHL were non-convecting magma bodies and crystallized at pressure-equivalent depths of more than 8.5 kin. We have performed one-dimensional thermal modelling in order to simulate the process of incremental growth of a laccolith whose roof is tectonically removed during intrusion, in a context of crustal exhumation due to channel flow. The objective is to define under what conditions HHL laccoliths emplaced close to active normal faults may be built without convecting while crystallizing muscovite. The results indicate that for a HHL thickness in the range 5-10 km, denudation rates cannot be higher than 4 mm a-1, and are more likely below 3 mm a- 1. At such denudation rates, the intrusion process needs to start at depths of c. 22 km, except when the final laccolith thickness is 10 km, in which case the depth of first-emplaced magmas cannot exceed 18 kin. Thick HHL laccoliths (>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.

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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

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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

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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,

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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.

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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.

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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

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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

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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

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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

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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.

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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.

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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

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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

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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

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the detachment, therefore no mineral was separated for isotopic dating.

Detachment fault surface

Sampling characteristics and distribution In order to document cooling age patterns across the detachment fault zone, four kinds of samples were collected from three structural positions: (1) mylonite on the surface of the detachment fault; (2) mylonite, metamorphic rocks and undeformed granite from the footwall of the detachment fault; and (3) metamorphic rocks, mylonite and leucogranite from the structurally deeper south-central part of the Greater Himalayan metamorphic sequence. All of the mica and K-feldspar separates were obtained from rocks along the detachment fault or within its footwall, with the exception of samples NL-25 and NL-29 collected from the south-central Greater Himalaya zone, 30 km south of the detachment fault in the Nyalam area (Figs 2 & 3). Multiple rock types from each location were chosen to maximize the number of datable minerals, allowing for better reconstruction of the cooling history for each area. No metamorphic rocks are found above

Five samples of mylonite were collected on the detachment fault surface (Fig. 3). Abundant muscovite grains are presented in these samples. The separated biotite and muscovite grains and K-feldspar porphyroclasts are set in a foliation of K-feldspar (c. 0.1-1cm) and recrystallized quartz (see Fig. 5 d - f ) . Samples NL-15, 36 and NL-39 were collected from nearby sites separated by north-south trending valleys, but at the same elevation. Samples NL-48 and NL-49 are from another valley, 5 km away from NL-15, but within the same detachment fault surface (see Fig. 3).

Footwall of the detachment fault Eight samples were collected from the footwall of the detachment fault (Fig. 3). Samples NL-2, NL-3, NL-4 and NL-7 were collected along a transect from elevation 3 9 0 0 m to 4800 m. These samples were collected from mylonite containing

332

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north-verging asymmetric folds (see Fig. 5b). Samples NL-2 and NL-3 are deformed metasedimentary rocks with an assemblage of quartz, plagioclase, muscovite, biotite, K-feldspar and hornblende. These samples have domains of topto-the-north microstructures, and coarse-grained recrystallized quartz, which may be associated with normal-sense faulting (see Fig. 5g-j). Samples NL-4 and NL-7 are near the detachment fault plane and are associated with a NNE-trending mylonitic lineation and north-dipping foliation.

Samples NL-37 and NL-38 were collected from the same position in the footwall of the detachment fault. No protomylonitic texture was observed in thin section and no metamorphic recrystallized quartz and biotite grains were identified. These samples are plagioclase-K-feldspar granite that intrude into metasedimentary rocks and mylonite. These undeformed granitic rocks contain biotite, hornblende, quartz and K-feldspar. Sample NL-53 was collected at a position less than 1.5 km south of the Nyalam detachment fault

SOUTH TIBETAN DETACHMENT SYSTEM in Ruji, north of Nyalam. This is a mylonitic rock, whose mineral assemblage includes biotite, muscovite and recrystallized quartz. It exhibits a north-dipping foliation and north-south trending lineation. Sample NL-55 is from an undeformed

333

leucogranite located 5 km to the south of sample NL-53. This sample consists mainly of muscovite, biotite, quartz and K-feldspar. No deformation and no recrystallized quartz were observed in thin section.

Fig. 5. Examples of microstructures along the detachment surface and in footwall. (a) Relict feldspar with asymmetric tails indicates top to north (normal) sense of shear. (b) Abundantly developed asymmetric folds and boudinage in marble unit below Nyalam detachment indicating normal sense of shear. (c) Deformed leucogranite on surface of detachment; leucogranite body is 0.1-2 m thick in field. (d) Muscovite grain parallel to lineation. (e) Recrystallized quartz around amphibole in mylonitized leucogranite. (f) Feldspar clast, 'mica-fish' and shear bands indicate normal sense of shear. (g) Deformed muscovite grain. (h) Shear bands around a feldspar clast and 'mica-fish'. (i) Recrystallized quartz tails around feldspar clast. (j) Mica, feldspar and recrystallized quartz indicate normal sense of shear. (k) Recrystallized quartz and biotite orientated parallel to foliation and lineation, developed in metasedimentary rocks. (1) Biotite and quartz anastomose around undeformed garnet clast in footwall to Nyalam detachment indicating weak effect from the overlying detachment. Abbreviations: Am, amphibole; Bit, biotite; Feld, K-feldspar; Gt, garnet; Mus, muscovite; P1, plagioclase; Qz, quartz; Recry. Qz, recrystallized quartz.

334

Y. WANG ETAL.

Fig. 5. (Continued)

South-central Greater Himalayan metamorphic sequence Two samples were collected south of the detachment area (Fig. 2), within the underlying Greater Himalayan metamorphic sequence. Samples NL-25 and NL-29 were collected from south of Nyalam, and are located in the central segment of the Greater Himalaya zone. Sample NL-25 has similar petrologic features to NL-53, but with a south-dipping foliation and north-south trending

lineation; the samples were collected from fold limbs (see Fig. 2b). Sample NL-29 a biotite gneiss, collected between and Zhangmu in the lower part of the Himalayan metamorphic sequence.

adjacent is from Nyalam Greater

Analytical methods Seventeen mica and three K-feldspar separates from mylonite, metamorphic rocks, leucogranite and granite were analysed for 4~

SOUTH TIBETAN DETACHMENT SYSTEM thermochronology. Three apatite separates from the same samples used for K-feldspar dating were also dated by the fission-track method (Wang et al. 2001). Potassium feldspar multidiffusion domain modelling data and a portion of the muscovite and biotite ages were collected using a MM-1200 micromass-spectrometer of the Institute of Geology, China Seismological Bureau. The remaining 4~ muscovite and biotite analyses were completed on a MM-5400 micromassspectrometer of China University of Geosciences. Twenty mineral separates of K-feldspar, muscovite and biotite were obtained and purified using a Frantz magnetic separator and conventional heavy organic liquid techniques. Individual grains were then selected under a binocular microscope. The mineral separates, and Fish Canyon Tuff sanidine and ZBH biotite (standard sample in China) flux monitors were irradiated in an atomic reactor belonging to the Research Institute of Atomic Energy of China and set