Deformation Mechanisms, Rheology and Tectonics: From Minerals to the Lithosphere (Geological Society Special Publication No. 243)

Deformation Mechanisms, Rheology and Tectonics" from Minerals to the Lithosphere

Geological Society Special Publications Society Book Editors R. J. PANKHURST (CHIEF EDITOR) P. DOYLE F. J. GREGORY J. S. GRIFFITHS A. J. HARTLEY R. E. HOLDSWORTH J. A. HOWE P. T. LEAT A. C. MORTON N. S. ROBINS J. P. TURNER

Special Publication reviewing 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 has a team of Book Editors (listed above) who 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. Geological Society Special Publications are included in the ISI Index of Scientific Book Contents, but they do not have an impact factor, the latter being applicable only to journals. More information about submitting a proposal and producing a Special Publication can be found on the Society's 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: GAPmS, D., BRUN, J. P. & COBBOLD, P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243. KOEHN, D, ARNOLD, J. & PASSCHIER, C. W. 2005. Fracture and vein patterns as indicators of deformation history, a numerical study. In: GAPAIS, D., BRUN, J. P. & COBBOLD, P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 11-24.

G E O L O G I C A L SOCIETY SPECIAL PUBLICATION NO. 243

Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere EDITED BY

D.

GAPAIS,J. P. BRuN and P. R. COBBOLD Universite de Rennes, France

2005 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society's fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society's international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the American Geological Institute (AGI), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists' Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies' publications at a discount. The Society's online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London WlJ 0BG: Tel. +44 (0)20 7434 9944; Fax +44 (0)20 7439 8975; E-mail: enquiries @geolsoc.org.uk. For information about the Society's meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society's Corporate Affiliates Scheme, write to [email protected].

Published by The Geological Society from: The Geological Society Publishing House Unit 7, Brassmill Enterprise Centre Brassmill Lane Bath BAI 3JN, UK (Orders: Tel. +44 (0)1225 445046 Fax +44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. 9 The Geological Society of London 2005. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/05/$15.00.

British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. I S B N 1-86239-176-9 Typeset by Techset Composition, Salisbury, UK Printed by Alden Press, Osney Mead, UK

Distributors USA AAPG Bookstore PO Box 979 Tulsa OK 74101-0979 USA Orders: Tel. +1 918 584-2555 Fax +1 918 560-2652 E-mail [email protected] India Affiliated East-West Press Private Ltd Marketing Division G-l/16 Ansari Road, Darya Ganj New Delhi 110 002 India Orders: Tel. +91 11 2327-9113/2326-4180 Fax +91 11 2326-0538 E-mail affiliat @vsnl.com Japan Kanda Book Trading Company Cityhouse Tama 204 Tsurumaki 1-3-10 Tama-shi, Tokyo 206-0034 Japan Orders: Tel. +81 (0)423 57-7650 Fax +81 (0)423 57-7651 E-mail geokanda@ ma.kcom.ne.jp

Contents

Preface Dedication

REDDY, S. M. & BUCHAN,C. Constraining kinematic rotation axes in high-strain

vii ix 1

zones: a potential microstructural method?

KOEHN, D., ARNOLD,J. & PASSCHIER,C. W. Fracture and vein patterns as

11

indicators of deformation history: a numerical study

MIRABELLA,F., BOCCALI,V. & BARCHI,M. R. Segmentation and interaction of

25

normal faults within the Colfiorito fault system (central Italy)

AUSTIN, N. J. & KENNEDY,L. A. Textural controls on the brittle deformation of

37

dolomite: variations in peak strength AUSTIN, N. J., KENNEDY,L. A., LOGAN,J. M. & RODWAY,R. Textural controls on the brittle deformation of dolomite: the transition from brittle faulting to cataclastic flow

51

RENARD, F., ANDRI~ANI,M., BOULLIER,A.-M. & LABAUME,P. Crack-seal patterns: records of uncorrelated stress release variations in crustal rocks

67

ZUBTSOV, S., RENARD, F., GRATIER, J.-P., DYSTHE, D. K. & TRASKINE, V. Single-contact pressure solution creep on calcite monocrystals

81

BARATOUX, L., SCHULMANN,K., ULRICH, S. 8~ LEXA, O. Contrasting microstructures and deformation mechanisms in metagabbro mylonites contemporaneously deformed under different temperatures (c. 650 ~ and c. 750 ~

97

GUEYDAN,F., MEHL, C. & PARRA,T. Stress-strain rate history of a midcrustal

127

shear zone and the onset of brittle deformation inferred from quartz recrystallized grain size DRURY, M. R. Dynamic recrystallization and strain softening of olivine aggregates in the laboratory and the lithosphere

143

RANALLI, G., MARTIN, S. & MAHATSENTE, R. Continental subduction and exhumation: an example from the Ulten Unit, Tonale Nappe, Eastern Austroalpine

159

RAIMBOURG,H., JOLIVET,L., LABROUSSE,L., LEROY,Y. & AVIGAD,D. Kinematics

175

of syn-eclogite deformation in the Bergen Arcs, Norway: implications for exhumation mechanisms BLENKINSOP, T. G. & KISTERS, A. F. M. Steep extrusion of late Archaean granulites in the Northern Marginal Zone, Zimbabwe: evidence for secular change in orogenic style

193

vi

CONTENTS

BROWN, M. Synergistic effects of melting and deformation: an example from the Variscan belt, western France

205

SPALLA,M. I., ZUCALI, M., DI PAOLA,S. t~ GOSSO, G. A critical assessment of the tectono-thermal memory of rocks and definition of tectono-metamorphic units: evidence from fabric and degree of metamorphic transformations

227

HANDY, M. R. BABIST, J., WAGNER, R., ROSENBERG, C. & KONRAD, M. Decoupling and its relation to strain partitioning in continental lithosphere: insight from the Periadriatic fault system (European Alps)

249

WILLINGSHOFER,E., SOKOUTIS,D. & BURG, J.-P. Lithospheric-scale analogue

277

modelling of collision zones with a pre-existing weak zone

DELACOU,B., SUE, C., CHAMPAGNAC,J.-D. & BURKHARD,M. Origin of the

295

current stress field in the western/central Alps: role of gravitational re-equilibration constrained by numerical modelling Index

311

Preface

This volume is derived from the 13th meeting on Deformation Mechanisms, Rheology and Tectonics (DRT2003). The meeting was held in St Malo (Brittany, France) in April 2003, and organized by an informal group from Gdosciences Rennes (UMR 6118 CNRS, Rennes University), including Michel Ballbvre, St6phane Bonnet, Arlette Falaise, Olivier Galland, Fr6d6ric Gueydan, Charles Gumiaux, Benjamin Le Bayon, Alain-Herv6 Le Gall, Monique Le Moigne, Sylvie Schueller, and C61ine Tirel. It was sponsored by the Centre National de la Recherche Scientifique, Rennes University, the city of Rennes, the Conseil G6n6ral d'Ille et Vilaine, and the R6gion Bretagne. Forty-eight reviewers have worked hard to improve the papers. We thank these persons and institutions for their contributions. The volume contains 18 papers that cover most of the topics and ideas presented and discussed during the meeting. The main approaches are experimental rock deformation, microstructural analysis, field studies, and analogue and numerical modelling. Several papers provide new insights on grainscale and aggregate-scale mechanisms, with new methodological implications (Gueydan et al., Reddy & Buchan, Zubstov et al.) and new advances on the knowledge of deformation processes and rock rheology (Austin & Kennedy, Austin et al., Baratoux et aL, Drury, Gueydan et al., Kiihn et al., Mirabella et al.,

Renard et al., Zubstov et al.) About half of the contributions provide new information or models for processes at crustal or lithospheric scales (Blenkinsop & Kisters, Brown, Delacou et al., Drury, Handy et al., Raimbourg et al., Ranalli et al., Spalla et al., Willingshofer et ai.), with particular emphasis on mantle rheology (Drury), thickening tectonics (Wiilingshofer et aL, Handy et al.) and various aspects of exhumation processes (Blenkinsop & Kisters, Delaeou et al., Raimbourg et al., Ranalli et al. ). The first meeting that led to the DRT series was in Leiden in 1976. Henk Zwart, Richard Lisle, Gordon Lister and Paul Williams were the organizers. The meeting was basically dedicated to Fabrics, Microtextures and Microtectonics. Since then, DRT meetings have accompanied international advances in the understanding of kinematics, mechanics and large-scale tectonic processes, which largely derived from important progress in the knowledge of rock rheology via experimental deformation, and in physical and numerical modelling. We hope that the present volume will further contribute to strengthening the links between studies of geological deformation processes at microscopic to lithospheric scales. Denis Gapais, Jean-Pierre Brun and Peter Cobbold Rennes, October 2004

Dedication to Pierre Choukroune

Pierre Choukroune was born in Casablanca on 28 March 1943. He celebrated his 60th birthday a few weeks before the DRT2003 meeting in St Malo. Pierre received a PhD from Paris University in 1967 and a 'Doctorat d'Etat' from Montpellier University in 1974. The subject of both theses was the structural geology of the Pyrenees. The second thesis was published as Mrmoire 127 of the Socirt6 Grologique de France. Pierre started his professional career in 1967 as 'Assistant' (Assistant Lecturer) at Montpellier University. In 1975, he moved to Rennes as Professor. He stayed at Rennes until 1995, when he moved to the University of AixMarseille. Building on his early fieldwork in the Pyrenees, Pierre rapidly acquired a prominent and highly personal scientific profile. His interests have always spanned various scales, from detailed analysis of microstructures, through regional patterns of strain and displacement (e.g. Choukroune 1976) to plate tectonics (Choukroune 1992). Of the more than 100 papers that he has published in international journals over the last three decades, many have had major impacts on structural geology and tectonics. One of his first papers (Choukroune 1971) was on the analysis of pressure shadows around pyrite crystals, as a tool to understand the development of slaty cleavage. As pointed out by Ramsay

and Hubert (1983) in their textbook, 'This is an outstanding paper from the view point of descriptive excellence, the quality of the diagrams and photographs and the theoretical analysis of data'. This paper contains many of the ideas that Pierre further put forward, in particular the search for criteria of shear senses in deformed rocks. Soon after his arrival in Rennes, Pierre was taken to see the granite mylonites of the South Armorican Shear Zone. He immediately realized the significance of the typical fabric of the rock that results from the coeval development of a schistosity and of shear bands parallel to the bulk shearing plane, and consequently introduced the now famous concept of C/S fabrics. Pierre put every effort into bringing them to international attention (Bertha et al. 1979a, b). The discovery was to open more than a decade of research on shear criteria. Because of his wish to find a strong link between plate tectonics and structural geology, Pierre has been in close contact with a number of geophysicists. In 1974, Xavier Le Pichon asked him to join FAMOUS, the French-American project that studied fault patterns on the Mid Atlantic Ridge (Choukroune et al. 1978). He later took his expertise to the East Pacific Rise (Hekinian et al. 1983; Choukroune et al. 1985) and the Gulf of Aden, where he acted as Chief Scientist for the CYADEN cruise (Choukroune et al. 1988). When the French deep seismic program ECORS was launched in the early 1980s, Pierre acted as Chief Scientist for the first French-Spanish profile across the eastern Pyrenees. The result was a major contribution of deep seismic data to the understanding of a mountain belt (Choukroune et al. 1989). As a team leader, Pierre had a profound influence on the development of structural geology and tectonics at G~osciences Rennes. His initiatives in the Hercynides, the Alps (Choukroune et al. 1986), and several Archaean cratons (Choukroune et al. 1995; Choukroune & Ludden 1997) provided training grounds for a number of young structural geologists from Rennes and elsewhere. None of his PhD students and colleagues at Rennes escaped his thoughtful suggestions and his unusual scientific lucidity and inspiration. Even though he liked to say, 'we

x

DEDICATION TO PIERRE CHOUKROUNE

are a team', what some researchers call the 'Rennes school' bears his signature. As a professor, Pierre's teaching style - elegant and to the point - has inspired many careers in Earth Sciences. His book, 'D6formations et d6placements dans la crofite terrestre' (1995), summarizes much of his scientific philosophy. As well as an outstanding scientist and teacher, Pierre has been an efficient and distinguished academic administrator, for the University, the CNRS, and the Ministry of Education, where he has held a number of important positions. The DRT meeting in St Malo was shortly after his 60th birthday. W e therefore take the greatest pleasure in dedicating the proceedings to Pierre Choukroune, a talented, innovative and inspiring Earth scientist. Jean-Pierre Brun, Peter Cobbold and Denis Gapais

References BERTHt~, D., CHOUKROUNE,P. & Jt~GOUZO,P. 1979a. Orthogneiss, Mylonite and non-coaxial deformation of granites: the example of the South Armorican Shear Zone. Journal of Structural Geology, 1, 31-42. BERTHI~, D., CHOUKROUNE,P. &; GAPAIS, D. 1979b. Orientations pr~f&rentielles du Quartz et orthogneissification progressive en r~gime cisaillant: exemple du Cisaillement Sud-Armoricain. Bulletin de Mindralogie, 102, 265-272. CHOUKROUNE, P. 1971. Contribution ~ l'6tude des m6canismes de d6formation avec schistosit6 grfice aux cristallisations syncin6matiques dans les 'zones abrit6es' ('pressure shadows'). Bulletin de la Socidtd Gdologique de France, 13, 257-271. CHOUKROUNE, P. 1976. Strain patterns in the Pyrenean Chain. Philosophical Transactions of the Royal Society of London, 283, 271-280.

CHOUKROUNE, P. 1992. Tectonic evolution of the Pyr6nEes. Annual Revues of Earth and Planetary Sciences, 20, 143-158. CHOUKROUNE, P. 1995. Ddformations et Ddplacements dans la Cro~te Terrestre. Masson, Paris. CHOUKROUNE,P. & Ecors team. 1989. The Ecors deep seismic profile reflection data and the overall structure of an orogenic belt. Tectonics, 8, 23-39. CHOUKROUNE, P. & LUDDEN, J. N. 1997. Archaean Crustal Growth and Tectonic Processes. Geological Society, London, Special Publications, 121, 63 -98. CHOUKROUNE, P., FRANCHETEAU,J. & Le PICHON, X. 1978. In situ observations along Transform Fault 'A' near 37~ in the FAMOUS area, Mid-Atlantic Ridge. Geological Society of America Bulletin, 89, 1013-1029. CHOUKROUNE, P., FRANCHETEAU,J. & HEKINIAN,R. 1985. Carte g6ologique de la Ride Est Pacifique Bulletin de la Socidt~ Gdologique de 12~ France, 1, 145-148. CHOUKROUNE, P., FRANCHETEAU, J. et al. 1988. Tectonics of an incipient oceanic rift. Marine Geophysical Research, 9, 147-163. CHOUKROUNE, P., BALLI~VRE, M., COBBOLD, P. R., GAUTHIER, Y., MERLE, 0. & VUICHARD, J. P. 1986. Deformation and motion in the Western Alpine Arc. Tectonics, 5, 215-226. CHOUKROUNE, P., BOUHALLIER, H. & ARNDT, N. 1995. Soft Archean lithosphere during periods of crustal growth or reworking. Early Precambrian Processes. Geological Society, London, Special Publications, 95, 67-86. HEKINIAN, R., FRANCHETEAU, J., RENARD, V., BALLARD, R. D. & CHOUKROUNE, P. 1983. Intense hydrothermal activity at the axis of the East Pacific Rise near 13~ Submersible witnesses the growth of sulphide chimney. Marine Geophysical Research, 6, 1-14. RAMSAY, J. G. & HUBERT, M. I. 1983. The Techniques of Modern Structural Geology, Volume 1: Strain Analysis. Academic Press, London.

Constraining kinematic rotation axes in high-strain zones: a potential microstructural method? S T E V E N M. R E D D Y & C R A I G B U C H A N Tectonics Special Research Centre, Department of Applied Geology, Curtin University of Technology, GPO Box U1987, Perth, WA 6845, Australia (e-mail: S. Reddy @ cu rtin. edu. a u / C. Buchan @ curtin, edu. au ) Abstract: The correct determination of the kinematic rotation axis in high-strain zones is

essential to the study of the tectonic evolution of the Earth's crust. However, the common assumption that the kinematic rotation axis lies orthogonal to the XZ plane of the finite strain ellipse may be invalid in the case of general shear. Orientation data obtained by electron backscatter diffraction from calcite, deformed in the high-strain Gressoney Shear Zone of the Western Alps, has been investigated using orientation maps, bulk sample crystallographic orientation and misorientation analyses, and detailed intragrain misorientation and crystallographic dispersion analysis. The results demonstrate a strong geometrical coincidence amongst (1) the bulk macroscopic kinematic rotation axis, (2) the orientation of misorientation axes associated with low-angle boundaries, and (3) rotation axes associated with crystallographic dispersion at the intragrain scale. This coincidence is interpreted to reflect a geometric control of the kinematic framework of the high-strain zone on the activity of crystal slip systems. It is proposed that this relationship may be exploited as a new microstructural tool to determine the orientation of bulk kinematic rotation axes in high-strain zones without assuming a geometric link between kinematic rotation and XZ sections. Although further testing is required, application of the approach may lead to a significant advance in our understanding of natural general shear deformation.

Understanding the kinematic evolution of highstrain zones (HSZ) is essential for interpreting the tectonic evolution and palaeogeographic reconstruction of modern and ancient orogens. Commonly HSZ are assumed to have developed by simple shear, enabling mineral lineation orientations to be geometrically related to the transport direction (Ramsay 1980; Hanmer & Passchier 1991). However, theoretical modelling of 'general shear', involving simultaneous components of coaxial and non-coaxial shear, predicts patterns of deformation that are far more complex than those produced by end member pure and simple shear alone (Sanderson & Marchini 1984; Fossen & Tikoff 1993, 1998; Robin & Cruden 1994; Tikoff & Teyssier 1994; Jones et al. 1997; Tikoff & Greene 1997; Jiang & Williams 1998; Lin et al. 1998; Passchier 1998). One important complexity is that the kinematic rotation axis (or vorticity vector) of the deformation may not lie orthogonal to the orientation of maximum finite stretch indicated by mineral stretching lineations. Consequently, the XZ surfaces of HSZ may not be the correct surface from which to gather kinematic (shear sense) data.

Despite the importance of constraining tectonic transport direction, there are currently few criteria with which to quantify the orientation of the kinematic rotation axis that are independent of the assumption that mineral stretching lineations lie orthogonal to the kinematic rotation axis. This paper presents quantitative microstructural data obtained by electron backscatter diffraction (EBSD) of polycrystals defornaed in an HSZ. The data demonstrate a geometric correlation between macroscopic kinematic rotation axes and both misorientation and dispersion axes associated with intragrain low-angle boundaries at the single and multiple grain scale. These observations indicate a potential means of constraining kinematic rotation axes that is independent of the orientation of mineral stretching lineations. The data presented here come from a single calcschist sample ($3-74 located at N45~ ' E7~ ') taken from the Gressoney Shear Zone of the Western Italian Alps (Fig. 1), an HSZ documented in some detail in previous studies (Reddy et al. 1999, 2003; Wheeler et al. 2001a). The Gressoney Shear Zone is dominated by bulk top-to-SE extensional deformation that

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 1-10. 0305-8719/05/$15.00 9 The Geological Societv of London 2005.

2

S.M. REDDY & C. BUCHAN

Fig. 1. Geological maps of the internal zones of the Alps (modified after Reddy et aL 2003). (a) Simplified geological map of the Western Alps; location of (b) marked by rectangular box. (b) Geological map of the internal zones of the Western Alps north of the Aosta Fault showing location of sample $3-74. Arrows indicate sense of shear with relative displacement of the hangingwall marked by the arrowhead. EMC, Eclogitic Micaschist Complex; GMC, Gneiss Minuti Complex; IIDK, Seconda Zona Diorito Kinzigitica;MR, Monte Rosa; AS, Arolla Schist; VP, Valpelline; PK, Pillonet Klippe. The heavy line indicates the contact between oceanic eclogite- and greenschist-facies rocks that mark the base of the Gressoney Shear Zone. was responsible for exhumation of oceanic eclogite facies rocks from lower to mid-crustal levels between 45 and 36 Ma ago (Reddy et al. 1999, 2003). Conservative estimates of bulk strain indicate "g > 50 across the 1 km wide HSZ (Wheeler et al. 2001a). However this strain was spatially and temporally localized (Reddy et aI. 2003) and in detail the Gressoney Shear Zone has a complex deformation history that may include kinematic partitioning (Reddy et al. 1999). The analysed sample contains well-developed kinematic indicators, including shear bands and mica fish structures that could be observed in three dimensions in the field. At both the macro- and microscopic scales, mineral stretching lineations can be traced continuously into shear bands suggesting that the shear bands are a true reflection of the overall sense of shear (Reddy et al. 1999, 2003). The symmetry plane of these monoclinic fabric elements contains the mineral elongation, and the pole of this plane is interpreted to represent the kinematic rotation axis. This geometry indicates either deformation approximating simple shear or general shear deformation in which the coaxial stretching direction was subparallel to the

non-coaxial maximum stretch component (Type E transpression of Fossen & Tikoff 1998). Since Gressoney Shear Zone deformation is not consistent with significant constrictional strain (Reddy et al. 1999, 2003) the former option is more likely. To summarize for sample $3-74, the kinematic rotation axis lies in the foliation plane, close to 90 ~ from the mineral lineation and therefore subparallel to Y (Fig. 2), and the deformation is close to end-member simple shear.

Analytical procedure An oriented petrographic thin section of the sample, cut parallel to the XZ section (XZ = 130.58NE, X = 16.321) was polished with progressively finer grades of diamond paste down to a 0.25 txm diamond polish. The section was then polished using a 0.06 txm colloidal silica suspension in a NaOH solution of pH 10 on a Buehler Vibromet II T M polisher. The sample was studied optically before being analysed in the scanning electron microscope (SEM). For EBSD analysis, the thin section was mounted in a Phillips XL30 SEM (operating conditions were 20 kV and spot size c. 0.7 txm) and tilted

CONSTRAINING KINEMATIC ROTATION AXES IN HIGH-STRAIN ZONES Kinematic Rotation

~ ........

!

' |

g ~

Fig. 2. Idealized HSZ coordinate reference frame used throughout this paper. The XY plane marks the shear plane of the HSZ with X the shear direction. In an HSZ deforming by simple shear the kinematic rotation axis corresponds to Y. Z marks the pole to the HSZ.

to 70 ~ Electron backscatter diffraction patterns (EBSPs) were obtained using an HKL Technology Nordlys detector and data were collected and processed using HKL Technology's Channel 5 software. Each EBSP was automatically matched by the Channel 5 software to theoretical EBSPs for calcite to give orientation information for each grid node. The closeness of fit between observed and theoretical patterns was good (average mean angular deviation = 0.62). Data were noise reduced using the Channel 5 'wildspike' correction and a four-neighbour zero solution extrapolation. For this study a single orientation map was constructed by automatically collecting EBSPs from a user-defined grid of 120 x 70 points at 4 txm spacing (map area = 480 x 280 txm). The average grain diameter of the mapped area was 17 ~m (range = 5-108 ~m) with mean grain area of 347 txm2. Zero solutions making up 18% of the map mostly reflect minor quartz and white mica in the sample. Channel 5 software was used to produce orientation maps, and calculate the minimum misorientation angle and misorientation axis geometry. The crystal symmetry of calcite means that a number of possible misorientation angles and axes may be calculated between two differently oriented crystals or parts of crystals. However, following several other workers, the angle/axis pair corresponding to the minimum misorientation angle was used for misorientation analysis (Pospiech et al. 1986; Mainprice et al. 1993; Lloyd et al. 1997; Wheeler et al. 2001b). This is justified for intragrain low-angle misorientation analysis.

3

Results The results from a single orientation map (Fig. 3a) indicate a strong crystallographic preferred orientation (CPO) for most of the calcite grains analysed (cf. Fig. 3a, d). The orientation of the c-axis cluster is consistent with macroscopic kinematic indicators that record topto-SE shearing (Fig. 3d). The only volumetrically significant exception to the CPO is grain 1 which has an a-pole parallel to the dominant c-direction in the sample (Fig. 3d). The uncorrelated misorientation angle distribution (Fig. 4a) indicates a smaller number of high misorientation angles than that predicted for a random distribution and is consistent with the presence of a CPO. The absence of a misorientation angle peak in the correlated data at 78 ~ indicates that few of the calcite grains are twinned (e.g. Bestmann & Prior 2003). However, twinning is present in other parts of the sample. The correlated misorientation angle distribution form the whole mapped area indicates numerous (c. 2000) low-angle boundaries between 2 ~ and 10 ~ (Figs 3a and 4a). These have orientations that cluster around the mesoscopic kinematic rotation axis (Y direction) of the sample (Fig. 4b). Although there are significant errors (up to 30 ~) associated with the determination of misorientations axes for small numbers of low-angle misorientations, these can be minimized by using a large dataset (Prior 1999). The clustering of c. 2000 low-angle misorientation axes around the centre of the pole figure (Fig. 4b) is statistically significant and is likely to reflect a geological process. Misorientation axes associated with correlated misorientation angles > 1 0 ~ have a random distribution (not shown), indicating no systematic relationship between the orientations of different calcite grains. The qualitative distribution of low-angle boundaries (Fig. 3a) is shown in more detail for two grains (1 & 2) in which the orientation data has been coloured to demonstrate crystallographic orientation variations of 15 ~ and 12 ~ respectively from the (blue) centre of each grain (Fig. 3b, c). These data indicate that the variations in orientation across these grains are systematic, forming discrete bands of constant relative misorientation towards the grain edges (Fig. 3b, c). The overall bending of the crystal lattice is cumulative and there are no individual angular misorientations greater than 10 ~ within either of the two grains. The systematic reorientation of crystal axes by low-angle boundaries is also apparent in pole figures of crystallographic orientation for grains 1 and 2 (Fig. 5) in which colour variations in the pole figures correspond

4

S.M. REDDY & C. BUCHAN

Fig. 3. (a) Automated EBSD map of calcite in $3-74 XZ (130.58NE) section. Each pixel (total = 6847) represents an individual orientation analysis. Colonrs reflect variations in orientation defined by the three Euler angles. Red and black lines represent boundaries with >2 and > 10~ misorientation respectively between adjacent analyses. Black pixels show analyses of quartz or white mica (data not shown) or zero solutions. Grains 1 and 2 correspond to grains shown in (b) and (c) respectively. (b) & (e) Detail of grains 1 (475 points) and 2 (304 points) derived by assigning a spectrum of colour (blue to red) to a range of 15~ and 12~ misorientations respectively from the centre of each grain. Systematic changes in colour indicate orientation variations due to low-angle boundaries. The spatial distribution of colours can be directly compared to crystallographic orientations indicated in Fig. 5. (d) Lower hemisphere, equal angle projections of calcite orientation data (c-, a- and r-poles) illustrated in (a). X and Z correspond to shear zone coordinates shown in Fig. 2. Colours in (a) correspond directly to position of orientation data in (d). Orientation data from grains 1 and 2 are indicated.

to the colours in the respective orientation maps (Fig. 3b, c). The crystallographic axes within two grains of different bulk crystallographic orientation (grains 1 & 2) are dominantly dispersed around small circles centred on the centre of the pole figures, that is, the macroscopic Y direction of the shear zone coordinate system (Fig. 5a, b). In both grains, one crystallographic axis at the centre of the pole figures shows little dispersion (grain 1 = r-pole; grain 2 -- a-pole). The relative lack of dispersion of this crystallographic orientation indicates that this represents a principal

rotation axis associated with low-angle boundary formation within the grains. In detail there are subtle complexities to this simple picture. In grain 1 there is slight dispersion of the central r-pole and other crystallographic axes that indicate minor dispersions away from a single small circle distribution centred around Y. The exact orientation of this second rotation axis is difficult to constrain due to the limited total angular dispersions associated with this axis ( < 5~ H o w e v e r it is interpreted to lie in the lower left quadrant of the pole figure. In grain 2, a second minor dispersion seemingly

CONSTRAINING KINEMATIC ROTATION AXES IN HIGH-STRAIN ZONES

Misorientation Distribution

5

coordinate system (Fig. 6c, d, g, h) indicate very little preferred orientation.

Discussion

,

Calcite deformation and the role of different slip systems

o 0o

10

20

h'~

Z

30

40 50 60 70 Misodentation Angle (~

80

90

100

Z

-

Fig. 4. (a) Minimum misorientation angle distribution for orientation data indicated in Fig. 3a. White histogram corresponds to neighbour-pair misorientations (correlated data), the black histogram representing random-pair misorientations (uncorrelated data). Uncorrelated data show a subsample of 2000 randomly chosen random-pair misorientations. Solid black line indicates the theoretical misorientation angle distribution of randomly oriented trigonal minerals. (b) The distribution of misorientation axes for minimum correlated misorientation angles of 2-5 ~ and 5-10 ~ shown in (a). These data indicate a concentration of low-angle misorientations axes around the sample Y direction. around small circles centred around the f-pole in the lower right quadrant accommodates a few degrees of rotation. In both cases these additional rotational components are minor and only slightly modify the dominant dispersion centred around the sample Y direction, that is, the r-pole and a-pole for grain 1 and 2 respectively. Misorientation axes for 2 - 5 ~ and 5 - 1 0 ~ minimum correlated misorientation angles plotted with respect to sample coordinates for grains 1 & 2 show a relatively widespread distribution (Fig. 6a, b, e, f). This is particularly evident for the 2 - 5 ~ misorientation data. Such a result is expected because of large errors involved with calculating misorientation axes associated with low-angle misorientations (Prior 1999). However, in general the average distribution of misorientation axes is again close to the Y direction of the shear zone coordinate system (Fig. 2). The inverse pole figures showing the position of misorientations axes with respect to the crystal

In two grains of differing orientation, the dispersion of crystallographic axes associated with low-angle boundaries appears to be dominated by rotation about an axis that lies parallel to the kinematic Y direction of the sample (Fig. 5). For grain 1, the rotation axis associated with the dispersion of crystallographic axes has an orientation that corresponds to an r-pole (Fig. 5). The dominance of a single dispersion axis for this grain suggests that most deformation was accommodated on low-angle boundaries developed by activity of a single crystallographic slip system. However, the obliquity of lowangle boundary traces on the orientation map (Fig. 3b), and the presence of a second minor dispersion axis (Fig. 5a), indicates the presence of at least one other slip system. Similarly, low-angle boundaries in grain 2 (Fig. 5) also developed by multiple slip systems but were dominated by dispersion around an axis parallel to one of the a-poles. The correlation of axes associated with dominant crystallographic pole dispersion with different crystallographic directions in grains 1 and 2 (r-pole and a-pole respectively) indicates that the dominant slip systems were different in the two grains. Crystallographic pole dispersion patterns can be used to identify intragrain slip systems using simple models of tilt boundary formation during the formation of edge dislocations (Lloyd & Freeman 1991, 1994). For tilt boundaries, the rotation axis for crystal slip is the normal to the plane that contains the pole to the slip plane and the slip direction, and therefore the three features are mutually orthogonal in calcite (De Bresser 1991; Bestmann & Prior 2003). Given that the dispersion axis lies in the centre of the pole figure, the slip plane and slip direction responsible for dispersion must lie on the primitive circle of the pole figure. Using this model, the data from grain 1 indicate the probability of a-slip ((-12-10) (-2021)) as the active slip system. Experimental evidence for a-slip in calcite is rare (Turner & Heard 1965; Paterson & Turner 1970) and consequently is largely ignored as an important slip system (e.g. De Bresser & Spiers 1997). However, no other published calcite slip systems readily account for the observed dispersion around the r-pole (Fig. 5a). In addition, Schmid Factor

6

S . M . REDDY & C. B U C H A N

a) Grain 1

Z

{-~olz}

b) Grain 2

y

Z

X

c Fig. 5. Equal area, lower hemisphere projections of orientation data (poles to c, r, f, e, a and direction (-2021)) for grains 1 and 2 of Fig. 3. In both grains most orientation data are dispersed in small circles around the centre of the pole figures. Dispersion axes are coincident with the r- and a-pole in grain 1 and 2, respectively. Minor secondary dispersion in both grains can also be identified and relates to the operation of secondary minor slip systems that cumulatively accommodate a few degrees of rotation,

CONSTRAINING KINEMATIC ROTATION AXES IN HIGH-STRAIN ZONES z

Grain 1

Grain 1

\

. •

7

",4

..

""" "

f

9

9

"" i;,.'. :"'_-" ,.-,. ".~)" ",

c) m z

9.

~t

Grain2

":'" "-

~''~'.,

=

x .......................

~

~

~

" -.

Grain 2

"': "" ~'" ...... = u ;#

g)

,

.

"=

=,

,~ |

II

....-',....

." ~""

h)

Fig. 6. Equal area, lower hemisphere projections of minimum misorientation axes orientation for low-angle boundaries in grains 1 and 2 plotted in both sample coordinate (pole figure) and crystal coordinate (inverse pole figure) space for misorientation angles of 2 - 5 ~ and 5-10 ~

analysis, assuming deformation was by simple shear, indicates higher resolved shear stress associated with a-slip (Schmid Factor = 0.5) than any other calcite slip system, further supporting the possibility of a-slip in grain 1. Since grain 1 is anomalous in its orientation (Fig. 3a, d) and makes up less than 5% of the analysed grains, its contribution to deformation of the sample as a whole is minor. The additional secondary dispersion axis identified in the pole figures (Fig. 5a) cannot be readily identified because of its small angular misorientations. Grain 2 has an orientation similar to the majority of analysed grains and records dispersion about the central a-pole. This dispersion axis is common to both of the common calcite slip systems r- and f-slip. However, only in the case of r-slip is the slip direction also lying on the primitive circle at 90 ~ to the pole to the slip plane (cf. Bestmann & Prior 2003). This geometry discriminates between the two slip systems and indicates the probability that the major dispersion in grain 2 was accommodated by r-slip ((10-14) (-2021)). Importantly, despite the different slip systems, the rotation axis associated with low-angle boundary formation in both grains lies parallel to the macroscopic kinematic rotation axis (Y), suggesting a kinematic control on the activity of the dominant slip systems.

The recognition of specific crystallographic dispersion axes for grains 1 and 2 should also be recorded by the geometry of low-angle misorientation axes both in sample and crystal coordinate systems (Fig. 6). However, the misorientation data are not easily reconciled with the crystallographic pole dispersion data. Despite having an average orientation close to the Y axis, misorientation axes plotted in the sample coordinate system have a distribution that demonstrates a large degree of variation, particularly in the lower (2-5 ~ misorientations angles (Fig. 6a, b, e, f). The apparent almost random distribution of misorientation axes in the crystal coordinate system (Fig. 6c, d, g, h) is also inconsistent with the specific crystallographic dispersion axes identified earlier (Fig. 5), which would be expected to cluster around r- and a-poles for grains 1 and 2 respectively. There are several possible reasons for the observed discrepancy between the crystal pole dispersion axes and the misorientation axes. Misorientation axes calculated from small angular misorientations have a significant error (tens of degrees) associated with them (Prior 1999). Such an error is significant in sample coordinate space (Prior 1999) and is amplified in crystal coordinate space where axes are plotted in a smaller angular volume. The error

8

S.M. REDDY & C. BUCHAN

associated with the calculation of misorientation axes decreases with increasing misorientation angle (Prior 1999) and it is noticeable that in grain 1, where we have sufficient 2 - 5 ~ and 5 - 1 0 ~ misorientations to make a comparison, that the distribution of, the misorientation axes for the larger misorientation angles is more tightly clustered around the kinematic Y direction (Fig. 6b). Furthermore, the 5 - 1 0 ~ misorientation axes in crystal coordinate inverse pole figures (Fig. 6d) are less variable than the 2 - 5 ~ axes (Fig. 6c). At least part of the variation in the distribution of misorientation axes is therefore attributed to errors associated with the calculation of misorientation axes for small angular variations. The misorientation axis is the common axis between two different crystallographic orientations around which a rotation through the misorientation angle will bring the two orientations into coincidence (Pospiech et al. 1986). As such it is a geometric construct that need not relate to a specific process. Changes in orientation associated with a number of different slip systems may therefore yield misorientation axes that are a combination of the different slip systems and may not have simple relationships to crystallographic orientations (Lloyd & Freeman 1991). Although in both grains 1 and 2 the dominant slip system is associated with a crystallographic pole dispersion axis centred on the sample Y direction, the presence of other operating slip systems is also apparent from secondary dispersions. In grains 1 and 2, the secondary slip systems are minor and accommodate small misorientations so a combination of different slip systems should be more apparent in the 2 - 5 ~ misorientations than in the larger misorientation angle data. Consequently, some of the variation in the orientation of 2 - 5 ~ misorientation axes may be a result of a combination of minor amounts of slip on different slip systems. As a result of the potential complications and the errors outlined above, individual misorientation axes are considered unreliable for identifying crystal slip systems associated with the development of low-angle boundaries. Dispersion axes obtained from crystallographic pole dispersion patterns (Fig. 5), particularly those illustrating a large cumulative angle of rotation, are considered to be a more robust method. However, the data presented here indicate that a large number of low-angle misorientation axes yield an average orientation that is identical to the rotation axis associated with crystallographic dispersion. Many studies investigating slip systems within deforming minerals have recognized the

importance of critically resolved shear stress (CRSS) for the initiation of slip on a particular slip system (e.g. Lister 1978; Takeshita et al. 1987 for calcite). The CRSS model assumes that slip on a particular system will occur when the shear stress on that slip system reaches a threshold value. Numerical investigations of slip systems in stress space indicate that the relative activity of these systems is controlled by the geometrical relationship of the yield surface to the imposed strain vector rather than the value of the CRSS (Lister & Hobbs 1980; Takeshita et al. 1987) suggesting that there is a geometrical control, in addition to a CRSS control, on slip system activity. The dispersion pattern dominated by a single rotation axis shown by our data suggests that activity is dominated by a single slip system in each grain, with a minor component of at least one other slip system. This is inconsistent with deformation as assumed in the Taylor models that consider single crystal yield surfaces requiring up to five active slip systems (Von Mises criteria) (Takeshita et al. 1987). Taylor modelling ignores the effect of dynamic recrystallization and recovery that takes place in natural deformations, as well as other deformation processes such as diffusive mass transfer, diffusion accommodated grain boundary sliding and grain boundary migration that may relax the necessity for five active slip systems. Experimental studies involving dynamic recrystallization at high temperatures indicate that geometric softening associated with the preferred orientation of specific calcite slip systems may take place and lead to the accommodation of large shear strains (Pieri et al. 2001). Studies of dynamic recrystallization during natural deformation also indicate that within individual grains, deformation may be accommodated by the operation of a single slip system (Bestmann & Prior 2003). The data presented here are consistent with these findings and indicate a similar geometric control on the activity of particular slip systems, whereby the dominant slip system that operates within a particular deforming grain is one in which the geometric rotation axis associated with the slip system lies parallel to the mesoscopic kinematic rotation axis. Microstructural characterization o f the kinematic rotation axis

The data presented in this study demonstrate a geometrical coincidence ofintragrain rotation axes responsible for the dispersion of crystallographic axes (Fig. 5), mean low-angle misorientation axes orientations from single crystals (Fig. 6),

CONSTRAINING KINEMATIC ROTATION AXES IN HIGH-STRAIN ZONES low-angle misorientation axes collected from polycrystals (Fig. 4), and the macroscopic kinematic rotation axis of the HSZ. These observations indicate that rotation and misorientation axes associated with low-angle boundaries may provide geometric constraints on the orientation of the macroscopic kinematic rotation axis. This similarity of misorientation and rotation axes for low-angle boundaries with mesoscopic kinematic rotation axes has been noted previously in EBSD studies (e.g. Bestmann & Prior 2003), although the kinematic significance of this relationship has not been explored. In this study, microscopic analysis has been linked to well-constrained kinematic data in which the shear direction of the studied HSZ has been previously established (Reddy et al. 1999, 2003). In this high-strain example the mineral lineation direction lies orthogonal to the kinematic rotation axis (Fig. 2). However, it cannot be assumed a priori that mineral stretching lineations lie orthogonally to the kinematic rotation axis. For example, in the case of general shear, the maximum finite stretch may be dominated by the coaxial component of the deformation that is independent of kinematic rotation. Despite this, mineral lineations are widely used to constrain shear directions and identify the ideal plane on which kinematic indicators should be viewed (XZ of Fig. 2). This may lead to incorrect interpretations of shear sense in HSZ deforming by general shear. The data presented here indicate that for simple shear deformation the orientation of the kinematic rotation axis may be constrained by identifying the rotation axes associated with crystallographic dispersion across low-angle boundaries and geometrically coincident low-angle misorientation axes. Since crystallographic orientations are independent of the orientation of the analysed section, crystallographic dispersion and misorientation axes are readily identifiable from EBSD data independent of lineation orientation. Although the data presented here refer to conditions approaching simple shear, similar geometrical relationships may exist under general shear conditions. If so, the analysis of misorientation and dispersion axes may yield important information regarding the orientation of kinematic rotation axes in HSZ undergoing general shear. Further detailed studies are required to test this hypothesis.

Microstructural preservation and the temporal evolution of HSZ Owing to the evolving nature of microstructures and their syn- and post-kinematic stability,

9

low-angle boundaries may reflect only the last stages of the deformation history in a particular part of an HSZ. Consequently, kinematic information obtained from the analysis of low-angle boundaries may not represent much of the earlier kinematic evolution of the HSZ and may provide only a snapshot of the lattermost stages of deformation. This presents a problem in constraining variations of kinematic history during the temporal evolution of the HSZ. Recent detailed geochronological studies of the Gressoney Shear Zone indicate that the spatial distribution of deformation with time was heterogeneous, with strain migration and localization leading to the preservation of older parts of the shear zone between zones of younger deformation (Reddy et al. 1999, 2003). Preliminary microstructural data from other Gressoney Shear Zone samples, which record different deformation ages and variable kinematic senses, indicate that the Gressoney Shear Zone was not affected by a homogenous or pervasive, late-stage microstructural overprint and that microstructures (including low-angle boundaries) developed early in the deformation history may be preserved in different temporal domains. Therefore, different parts of the HSZ appear to have a microstructural memory that can be used to constrain the earlier parts of the deformation history. Consequently, the application of the technique presented in this paper to HSZ that record a complex evolution may lead to a significant advance in our understanding of kinematic partitioning in naturally deformed strain systems with respect to both space and time. This work was funded through ARC Large Grant A00106036. The data presented in this paper were collected at the Australian Research Council (ARC) funded Microstructural Analysis Facility (part of the Western Australian Centre of Microscopy)at Curtin Universityof Technology,Perth, Australia.GeoffLloyd,Neil Mancktelow and Denis Gapais are thankedfor comprehensiveand constructivereviews. This paper is TSRC publicationNo. 288 and contributes to the TSRC's Program 4: Tectonic Processes.

References BESTMANN, M. & PRIOR, D. J. 2003. Intragranular dynamic recrystallisation in naturally deformed calcite marble: diffusion accommodated grain boundary sliding as a result of subgrain rotation recrystallization. Journal of Structural Geology, 25, 1597-1613. DE BRESSER,J. H. P. 1991. Intracrystallinedeformation of calcite. Geologica Ultrajectina, 79, 1- 191.

10

S.M. REDDY & C. BUCHAN

DE BRESSER, J. H. P. & SPIERS, C. J. 1997. Strength characteristics of the r, f, and c slip systems in calcite. Tectonophysics, 272, 1-23. FOSSEN, H. & TIKOFF, B. 1993. The deformation matrix for simultaneous simple shearing, pure shearing and volume change, and its application to transpression - transtension tectonics. Journal of Structural Geology, 15, 413-422. FOSSEN, H. & TIKOFF, B. 1998. Extended models of transpression and transtension, and application to tectonic settings. In: HOLDSWORTH, R. E., STRACHAN, R. A. & DEWEY, J. F. (eds) Continental

Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 15-33. HANMER, S. & PASSCHIER, C. W. 1991. Shear-sense indicators: a review. Geological Survey of Canada, Paper 90-17, 72. JIANG, D. & WILLIAMS, P. F. 1998. High-strain zones: a unified model. Journal of Structural Geology, 20, 1105-1120. JONES, R. R., HOLDSWORTH, R. E. & BAILEY, W. 1997. Lateral extrusion in transpression zones: the importance of boundary conditions. Journal of Structural Geology, 19, 1201-1217. 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, 41-57. LISTER, G. S. 1978. Texture transitions in plastically deformed calcite rocks. Proceedings of the 5th

International Conference on Textures of Materials, 199-210. LISTER, G. S. & HOBBS, B. E. 1980. The simulation of fabric development during plastic deformation and its application to quartzite: the influence of deformation history. Journal of Structural Geology, 2, 355 -370. LLOYD, G. E. & FREEMAN, B. 1991. SEM electron channelling analysis of dynamic recrystallization in a quartz grain. Journal of Structural Geology, 13, 945-953. LLOYD, G. E. & FREEMAN, B. 1994. Dynamic recrystallization of quartz under greenschist facies conditions. Journal of Structural Geology, 16, 867- 881. LLOYD, G. E., FARMER, A. B. & MAINPRICE, D. 1997. Misorientation analysis and the formation and orientation of subgrain and grain boundaries. Tectonophysics, 279, 55-78. MAINPRICE, D., LLOYD, G. E. & CASEY, M. 1993. Individual orientation measurements in quartz polycrystals - advantages and limitations for texture and petrophysical property determinations. Journal of Structural Geology, 15, 1169-1187. PASSCHIER, C. W. 1998. Monoclinic model shear zones. Journal of Structural Geology, 20, 1121-1137.

PATERSON, M. S. & TURNER, F. J. 1970. Experimental deformation of strained calcite crystals in extension. In: PAULITSCH, P. (ed) Experimental and Natural Rock Deformation. Springer, Berlin, 109-141. PIERI, M., KUNZE, K., BURLINI, L., STRETTON, I., OLGAARD, D. L., BURG, J.-P. & WENK, H.-R. 2001. Texture development of calcite by deformation and dynamic recrystallization at 1000 K during torsion experiments of marble to large strains. Tectonophysics, 330, 119-140. POSPIECH, J., SZTWIERTNIA,K. & HAESSNER,F. 1986. The misorientation distribution function. Textures and Microstructures, 6, 201. PR1OR, D. J. 1999. Problems in determining the misorientation axes, for small angular misorientations, using electron backscatter diffraction in the SEM. Journal of Microscopy, 195, 217-225. RAMSAY, J. G. 1980. Shear zone geometry: a review. Journal of Structural Geology, 2, 83-99. REDDY, S. M., WHEELER, J. & CLIFF, R. A. 1999. The geometry and timing of orogenic extension: an example from the Western Italian Alps. Journal of Metamorphic Geology, 17, 573-589. REDDY, S. M., WHEELER, J. et al. 2003. Kinematic reworking and exhumation within the convergent Alpine Orogen. Tectonophysics, 365, 77-102. ROBIN, P. Y. F. & CRUDEN, A. R. 1994. Strain and vorticity patterns in ideally ductile transpression zones. Journal of Structural Geology, 16, 447 -466. SANDERSON, D. J. & MARCHINI, W. R. D. 1984. Transpression. Journal of Structural Geology, 6, 449-458. TAKESHITA, T., TOMI~, C., WENK, H.-R. & KOCKS, U. F. 1987. Single-crystal yield surface for trigonal lattices: Application to texture transitions in calcite polycrystals. Journal of Geophysical Research, 92, 12917-12930. TIKOFF, B. & TEYSSIER, C. 1994. Strain modeling of displacement-field partitioning in transpressional orogens. Journal of Structural Geology, 16, 1575-1588. TIKOFF, B. & GREENE, D. 1997. Stretching lineations in transpressional shear zones: an example from the Sierra Nevada Batholith, California. Journal of Structural Geology, 19, 29-39. TURNER, F. J. & HEARD, H. C. 1965. Deformation in calcite crystals at different srain rates. University

of California Publications in Geological Sciences, 46, 103-126. WHEELER, J., REDDY, S. M. & CLIFF, R. A. 2001a. Kinematic linkage between internal zone extension and shortening in the more external units in the NW Alps. Journal of the Geological Society, London, 158, 439-443. WHEELER, J., PRIOR, D. J., JIANG, Z., SPE1SS, R. & TRIMBY, P. W. 200lb. The petrological significance of misorientations between grains. Contributions to Mineralogy and Petrology, 141, 109-124.

Fracture and vein patterns as indicators of deformation history: a numerical study D A N I E L K O E H N , J O C H E N A R N O L D & CEES W. P A S S C H I E R

Tectonophysics, Institute for Geosciences, Becherweg 21, University of Mainz, 55099 Mainz, Germany (e-mail: koehn @mail. uni-mainz, de) Abstract: Fracture and vein patterns in the brittle crust of the Earth contain information on

the stress and strain field during deformation. Natural examples of fracture and vein patterns can have complex geometries including combinations of extension and conjugate shear fractures. Examples are conjugate joint systems that are oriented with a small angle to the principal stress axis and veins that show an oblique opening direction. We developed a discrete numerical model within the modelling environment 'Elle' to study the progressive development of fractures in two dimensions. Results show that pure shear deformation alone can produce complex patterns with combinations of extension and shear fractures. These patterns change in geometry and spacing depending on the Young's modulus of the deforming aggregate and the initial noise in the system. A complex deformation history, including primary uniaxial loading of the aggregate that is followed by tectonic strain, leads to conjugate shear fractures. During progressive deformation these conjugate shear fractures may accommodate extensional strain or may be followed by a secondary set of extension fractures. The numerical patterns are consistent with joint, fault and vein geometries found in natural examples. The study suggests that fracture patterns can record complex deformation histories that include primary uniaxial loading due to an overlying rock sequence followed by tectonic strain.

Brittle deformation involving fracturing of rocks is a very important deformation mechanism in the upper crust of the Earth (Patterson 1978; Ranalli 1995). Many sedimentary rocks that are exposed today are filled with joints, faults and veins with characteristic patterns reflecting the importance of brittle deformation (Price & Cosgrove 1990). In addition, fracture and vein patterns are used by structural geologists as a map of the far-field stress orientation at the time of fracture formation (Ramsay & Huber 1983; Price & Cosgrove 1990; Oliver & Bons 2001). Fractures are generally classified into mode I, mode II and mode III fractures (Pollard & Segall 1987; Scholz 2002). Mode I fractures are extensional and open perpendicular to the maximum tensile stress. They accommodate strain by opening and may contain vein material that precipitates in the open space. Mode II fractures are shear fractures that develop at an angle to the maximum principal stress. They accommodate strain by slip along the fracture plane and are therefore called shear fractures or faults if significant slip has taken place along them (Bonnet et al. 2001). Mode III fractures are important in

three dimensions and will not be considered further in this paper since the presented models are two-dimensional. In order to understand what kind of fracture will form under a given stress condition and what orientation it will have with respect to one of the major stress axes it is useful to construct a Mohr circle diagram (Jaeger & Cook 1976) with the normal stress on the x-axis and the shear stress on the y-axis (Fig. 1). The Mohr envelope or failure curve in Figure 1 is constructed after Griffith (Griffith 1920; Jaeger & Cook 1976; Scholz 2002). If the Mohr circle crosses the failure curve in the regime of tensile stress, mode I fractures will develop, if it cuts the failure curve in the compressive stress regime, mode II fractures will form with an angle 0 to the direction of the compressive stress. We will term fractures that form by a coalescence of small mode I cracks, extension fractures (joints in a larger volume of rock or veins if they are filled with new material) and fractures that form by a coalescence of mainly mode II cracks, shear fractures (faults in a larger volume of rock if slip takes place). In addition, Hancock (1985) and Price & Cosgrove (1990) mention an intermediate

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. DeformationMechanisms, Rheology and

Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 11-24. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

12

D. KOEHN ET AL.

hybrid i ~ modeII (3"1 , Oli ~ ~

i

failurecurve Ze=4To(O'n+To)

,T , "

,,

, , ..,

", T

0-I (In

', '

~ AO

failure curve

Fig. 1. A Mohr circle diagram with the normal stress (o-.) on the x-axis and the shear stress (~-) on the y-axis. The failure curve is constructed using the Griffith failure criterion (Jaeger & Cook 1976). Two Mohr circles with different sizes are shown. The smaller circle cuts the failure envelope in the tensile regime and produces mode I fractures and hybrid extension/shear fractures. The larger circle cuts the failure envelope in the compressive regime and produces mode II shear fractures. The mean stress (o-m), differential stress (Act) and the two main principal stress components (o'l, o'2) are indicated next to the larger Mohr circle. To represents the tensile strength.

fracture type called hybrid extension/shear fractures that form at an angle to the compressive stress direction but still lie in the tensile stress regime (Fig. 1) and form probably by a coalescence of mode I and II cracks. One has to note, however, that the occurrence of hybrid extension/shear fractures as propagating cracks and their orientation with respect to the principal stress direction and thus their connection to the parabolic failure curve in Figure 1 is questionable (Engelder 1999; Ramsay & Chester 2004). Rocks in the brittle regime can only sustain small amounts of elastic strain before they start to form fractures (Means 1976). If fractures develop, the overall behaviour of the rock will be plastic (Paterson 1978). Fractures accommodate strain in a rock and relieve the overall stress. However, since fractures, are two-dimensional planar structures, in order to accommodate threedimensional strain, complex fracture patterns develop. Anderson (1951) introduced the idea of conjugate faults, which intersect in the intermediate stress direction. The plane perpendicular to the intermediate stress axis then includes the minimum and maximum stress directions and the slip vectors of the faults. This concept, however, only works if the intermediate stress is zero or

close to zero. Otherwise more complex fault or fracture patterns develop with at least four sets of coeval faults, arranged in an orthorhombic symmetry (Oertel 1965; Aydin & Reches 1982; Krantz 1988). Fractures that are filled with vein material pose a problem. Rocks are normally loaded by the overlying rock sequence before they are deformed tectonically, so both the vertical and horizontal stresses will be compressional. Therefore extension fractures are not expected to form since the Mohr circle will cut the failure envelope in the compressive regime (Jaeger & Cook 1976; Means 1976; Suppe 1985; Price & Cosgrove 1990). However, veins are formed quite often in the upper crust, so an alternative mechanism is needed to produce tensile stresses. Two possible sources are (1) a heterogeneous layered sequence where tensile stresses are concentrated in competent layers forming fractureboudinage or (2) the presence of a high fluid pressure that shifts the Mohr circle towards the tensile regime (Price & Cosgrove 1990; Ranalli 1995). In these cases one can find extension fractures that are filled with vein material if material can precipitate. An example of a hybrid shear fracture that is filled with vein material and laterally extends into extension fractures is shown in Figure 2a. An interesting question is if a vein set like the one shown in Figure 2b can also be associated with a single set of conjugate hybrid shear fractures. Examples of conjugate hybrid joint sets can be found in Hancock (1985). In this paper we model the dynamic development of fracture patterns in a two-dimensional approach. We simulate different deformation histories starting from a simple case where we investigate different internal model parameters. Then we extend to more complex deformation histories that include gravitational loading and tectonic strain. We will argue that the initial fracture geometries that later develop into veins contain a record of the gravitational loading that preceded tectonic deformation. It will be shown that even during a simple deformation history different stress states in the samples will lead to different kinds of failures.

The model In order to model two-dimensional fracture patterns we developed a discrete-element model based on the work of Malthe-Screnssen et al. (1998b) and combined it with the modelling environment 'Elle' (Jessell et al. 2001). With this kind of model, linear elastic behaviour of rock aggregates can be simulated with the full description of the strain and stress field. The

FRACTURE AND VEIN PATTERNS

(a)

13

(b)

Fig. 2. (a) Sketch of a vein that may have developed partly as a hybrid shear fracture (after Price & Cosgrove 1990, Fig. 1.63). (b) Bedding plane with conjugate sets of veins from a locality near Sestri Levante in the Liguride units, Italy (see hammer for scale).

model also allows the introduction of fractures by bond-breaking and can thus be used to investigate the behaviour of the aggregate beyond linear elasticity theory. Studies with this kind of model have shown that it reproduces the scaling behaviour and statistics of experimental fracture patterns (Walmann et al. 1996; MaltheS0renssen et al. 1998a, b, 1999) and more complex systems (Jamtveit et al. 2000; Flekk0y et al. 2002). In the discrete model circular particles are connected with their neighbours by linear elastic springs (Fig. 3a). We use a triangular lattice in two dimensions since it reproduces linear elasticity theory with only nearest neighbour interaction (Flekkcy et al. 2002). The force acting on a particle from a neighbour is proportional to the extension or compression of the connecting spring with respect to the equilibrium distance. By definition, compressive stresses in the model are negative and tensile stresses positive. The lattice is in equilibrium condition if all forces acting on a particle cancel out. If the equilibrium condition is perturbed, for example by applying a deformation to the lattice, particles

.(~)~~ ~

(b bound,ary (c)

fracture

Fig. 3. Sketches of the discrete particle model that is used for the numerical simulations. (a) Triangular lattice showing particles and springs that connect particles. (b) Boundary particles (in grey) are used to apply deformation on the modelled aggregate. (c) Fractures are induced by the possibility of breaking springs. Springs break once they have reached a critical tensile stress and are irreversibly removed from the model.

are moved according to resulting forces until all particles are in an equilibrium position defined by a given threshold. In order to find equilibrium for the whole lattice efficiently it is over-relaxed by moving the particles by an 'over-relaxation factor' beyond their equilibrium condition (Allen 1954). In the model, particles fill an initial area called the deformation box. Particles along the box boundaries are defined as wall-particles and are used to apply kinematic boundary conditions (Fig. 3b). These particles are fixed perpendicular to the boundary of the model but can move freely parallel to it, for example, free-slip boundary conditions are applied. In order to strain the model, walls are moved inwards to apply compression or outwards to apply extension. Once a wall is moved all particles in the deformation box are moved assuming homogeneous deformation. Then the relaxation algorithm starts. Once the lattice is in an equilibrium condition after a deformation step, all springs are checked for breaking. If the stress associated with a spring reaches a critical tensile value, fracturing is induced by breaking springs (Fig. 3c). The spring with the highest probability of breaking will do so and a new relaxation routine will start. This procedure is repeated until all springs are below the breaking threshold and a new deformation step is applied. The fracturing process is therefore highly nonlinear; the whole lattice can fail at once between two deformation steps. This microscopic failure criterion will result in shear and extension failure between two neighbouring particles because particles are connected with six springs to their neighbours. Particles with broken springs are still repulsive. The coupling of the discrete code with the 'Elle' environment is performed as follows. In 'Elle' the microstructure of a rock is defined by nodes, which are connected with boundary

14

D. KOEHN E T A L .

Fig. 4. (a) Combination of the 'Elle' modelling environment with the discrete model. In Elle, grains are defined by double and triple nodes that are connected by straight segments. Particles of the discrete model that lie within an Elle grain have the same properties. (b) Example of a microstructurethat was used for most of the simulations shown in this paper. segments that define grains (Fig. 4a). In the discrete code grains are filled with particles. All particles within a grain can have specific parameters like the elastic constant and the breaking strength of springs. Springs that connect particles of different grains define grain boundaries and have different properties to springs that lie within grains. In all the presented models grain boundary springs have about half the tensile strength of intragrain springs. This value is chosen in order to have an influence of grain shapes on fracture development. Therefore most fractures will be intergranular. Intragranular fractures mainly appear when shear fractures develop. We expect grain boundaries to have a lower cohesion due to the presence of fluids or impurities. In order to reduce lattice effects, noise is distributed in the model by three different methods. First an initial microstructure is drawn and used as an input file with a certain spatial distribution of grain boundaries (Fig. 4b). Then a statistical distribution is applied to the elastic constants of different grains. Each grain has an individual elastic constant chosen by a random function from a gauss distribution. The mean elastic constant of the whole deformation box is represented by the mean value of the distribution. The breaking strength of all springs is statistically dispersed using a linear distribution because a rock is also expected to be disordered on a scale smaller than the grain size (Malthe-SOrenssen et al. 1998b). This distribution is not dependent on grains and grain boundary springs still have about half the breaking strength of intragrain springs.

The presented simulations have a resolution of 184 800 particles within the deformation box. In all simulations we strain the model in small steps of 0.02%. In the model the unscaled mean value of the initial elastic constant is 1.0 and the breaking strength of intragranular springs is 0.006. Grain boundaries always have half the breaking strength of intragranular springs. We scale these values in the results section in order to discuss the models in a geological context. We assume that the breaking strength of grain boundaries is 30MPa. Therefore a value of 0.0001 in the model scales to 1 MPa. An initial spring constant of 1.0 thus represents an elastic constant of 10 GPa. This leads to a mean tensile strength of the aggregates used in the simulations of about 25 MPa and a mean compressive strength of about 90 MPa under pure shear deformation (starting from zero strain and zero stress). This compressive strength is relatively low because in pure shear deformation one principal stress is tensile if the experiment is not preloaded. If this stress is also compressive the breaking strength will be considerably higher. This relation is also seen in laboratory experiments and is predicted by theory (Jaeger & Cook 1976; Paterson 1978). If the aggregate in our simulations is preloaded before tectonic deformation is applied (the confining pressure is increased), the compressive strength rises to about 140-190 MPa.

Results We present a number of simulations of fracture development with different deformation histories

FRACTURE AND VEIN PATTERNS and show the dependence of the pattern on the initial noise in the system and differences in the mean elastic constant of the aggregate. First we discuss fracture patterns that develop under pure shear deformation. We investigate the influence of noise and mean elastic constant of the aggregate on pattern formation using examples that experience the same pure shear deformation history. Then we move on to discuss deformation patterns that develop during area increase (expanding the model area). Finally we include complex deformation histories where samples experience an initial uniaxial loading and then a tectonic deformation. This leads to a number of deformation histories, which produce the vein and fracture patterns geologists can find in some field areas.

mean stress and the differential stress. Note that compressive stress is negative. Figure 5b shows schematic Mohr circles that represent stress states for some of the deformation events shown in Figure 5a. The developing patterns are illustrated in Figure 5c. Starting from a completely relaxed state, during progressive deformation the stress field behaves according to linear elasticity theory. Principal stresses are oriented parallel to the directions of compression and extension and increase linearly with the same slope into the tensile and compressive field. Consequently the mean stress remains zero and the differential stress curve has a slope that is twice as steep as that of a single principal stress component (Fig. 5a). The corresponding Mohr circle grows but remains stationary at the intersection of the axis of the diagram. A peak in the curve of the tensile stress marks the intersection when the Mohr circle reaches the tensile part of the failure envelope. At this point tensile fractures grow in the sample perpendicular to the extension direction. They cross most of the sample and show a distinct spacing. These fractures accommodate extensional strain and relax tensile stresses. This effect is also seen in the stress-strain curve in Figure 5a. The tensile

Pure shear deformation First we examine fracture patterns that are generated in a sample that experiences pure shear deformation. The initial sample is not stressed, deformation starts from a completely relaxed state. Figure 5a shows the stress-strain relationship during deformation. Different curves correspond to the maximum and minimum stress, the

~(uPa) (a)

400

A(~(~~m

8o

-40

I

0

"

-

0.01

0.0Z

E

15

(b)~l;~n ~

(+)

~

/

(')

A~ T

,~

Fig. 5. Simulation of pure shear deformation of an aggregate. (a) Stress-strain relation showing the two principal stresses (oh, ~r2),the mean stress (Crm)and the differentialstress (Ao-).(b) SchematicMohr circle diagram of (a) showing Mohr circles during different stages of the simulation. They mark the initial starting configuration,mode I failure and mode II failure. Note that the x-axis is negative towards the right-hand side since compressivestresses are negative in this paper. (c) Three successive stages during the simulation. The picture on the left-hand side shows development of mode I fractures. The pictures in the middle and on the right-hand side show development of mode II shear fractures.

16

D. KOEHN ET AL.

stress drops, the mean stress starts to increase and deviates from zero and the differential stress has a slope that is less steep. The compressive stress is not much affected by extension fractures because the vertical load is still supported by the sample. The Mohr circle in Figure 5b starts to move from the tensile regime into the compressive region and continues to grow in radius. The next failure event is reached when the differential stress is high enough so that the Mohr circle cuts the upper part of the failure curve and shear fractures develop. Before this moment is reached the stress curves show increasingly flattening slopes due to the local development of small shear fractures prior to failure of the whole sample. The failure of the sample is accommodated by slip along shear fractures (shear fracture becomes a distinct fault), which relaxes the compressive stress in the sample in a number of steps as seen in the stress-strain curve. The differential stress is also relaxed until the sample reaches a quasi steady-state where stresses remain almost constant and the bulk behaviour is almost purely plastic. This deformation history will result in the primary growth of extension fractures and a secondary growth of conjugate shear fractures.

(a}

t

Extension fractures will progressively open to accommodate extensional strain so that veins can form whereas the shear fractures will mainly show slip along their surfaces and develop into faults. Thus, in a natural example one would expect only one set of veins that develops out of extension fractures.

Differences in initial noise and elastic constant Differences in initial statistical noise and in the mean elastic constant of the aggregate influence the developing pattern. By initial noise in the system we mean the width of the distribution of elastic constants of different grains and the breaking strength of individual grains. Figure 6 shows two simulations with identical initial microstructure and elastic constants as in Figure 5, but with varying width of the distribution of the breaking strengths. The simulation shown in Figure 6a has a distribution of breaking strengths from 24 to 36 MPa whereas the simulation shown in Figure 6b has a distribution of breaking strengths ranging from 12 to 48 MPa. For each simulation two successive stages at the same amount of strain are shown, deformation is

) i

,.

...,~ :

~,

/

/ ( ?

(b)

./

3,

9 4

'!

.

Fig. 6. Two simulations with the same boundary conditions but different initial distributions of breaking thresholds, where (a) has a narrower distribution of breaking thresholds than (b). (a) Less noise in the system produces stronger localization of structures. (b) More noise produces more dispersed fractures with different sizes.

FRACTURE AND VEIN PATTERNS pure shear where compression is vertical. Stage one shows the initial development of extension fractures and stage two the development of secondary shear fractures. Two important differences can be observed between the different distributions of breaking strengths. First a wider distribution of breaking strengths results in a more dispersed development of fractures whereas a narrow distribution results in localized fractures that cut the whole aggregate. Secondly, the absolute breaking strength is lowered using a wider distribution of breaking strengths; more fractures develop in Figure 6b than in Figure 6a. The first effect can be explained as follows. A narrow distribution of breaking strength, which means that the initial noise in the system is low, represents a very brittle material. A lot of springs in the aggregate will reach their tensile strength within a single deformation step. This will result in an almost instantaneous propagation of large fractures through the aggregate. In particular, an extension fracture will continue to propagate once it nucleates since tensile stresses at the fracture tip will increase while the fracture grows in length. Once a large fracture is present it will be able to relax stresses

(a)

in the surrounding aggregate so that additional fractures will only develop at a certain distance to the initial fracture. This stress shielding mechanism produces a distinct spacing of fractures in the aggregate. If the width of the distribution of breaking strengths in the model is larger (Fig. 6b), fractures can develop at lower strains. This will reduce the overall breaking strength of the aggregate. Since the strength of springs varies significantly, fractures that nucleate may stop propagating. Therefore a range of fractures develops initially with different lengths. More fractures nucleate in an aggregate with a wider distribution of breaking strengths and the developing spacing of fractures will not be as distinct as in the example with the narrow distribution. The material has a bulk response to deformation that is less brittle. The effect of the mean elastic constant (Young's modulus) of the aggregate on pattern formation and rheological behaviour is shown in Figure 7. The mean Young's modulus of the aggregate increases from Figure 7a to e. Figure 7a to e shows an aggregate with a mean Young's modulus of 1, 2.5, 5, 7.5, and 12.5GPa. All simulations were performed

(b)

,( '. 2

"

17

Cc)

" '

it

!

!

0

i

9 ,

~

_,oi ";'F""

/ I

I 0 . . . . . . . . . . . . . . 0.01 . . . . . .

-~oI 0

~0,Ol

0,0~ E

mode,,

I~

0.02

Fig. 7. (a) to (e) show simulations under the same pure shear conditions with increasing mean elastic constant. A lower elastic constant produces stronger localization and a wider spacing of mode I fractures. If the elastic constant is higher, mode II fractures dominate. (f) Stress-strain relation for (a). Mode I failure shows a sawtooth curve typical for fast propagation of large mode I fractures. (g) Stress-strain relation for (e). Mode I failure is continuous with no sudden drop of the tensile stress (02 curve). Mode II failure results in a drop of the differential stress.

D. K O E H N ET AL.

18

using pure shear deformation and the snapshots presented are taken at similar amounts of finite strain. The breaking strength of springs in all simulations has the same distribution and mean value. Consequently failure will occur at lower strains in an aggregate that has a higher elastic constant. Therefore the number of fractures increases from Figure 7a to e. In addition, the compressive stresses are not high enough to induce shear failure in Figure 7a and only to a minor extent in Figure 7b and c, whereas shear fractures dominate in Figure 7e. An aggregate with a lower elastic constant localizes extensional strain in tensile fractures, which open and form large veins in Figure 7a and to a minor extent in Figure 7b and c. Figure 7d shows a combination of extension fractures and shear fractures. The spacing of extension fractures depends on the mean elastic constant of the aggregate. In the simulation in Figure 7a one large vein develops within the model whereas Figure 7b shows two and Figure 7c three large veins. The spacing is reduced with an increase in Young's modulus. A softer material with a lower elastic constant can relax a larger region around an opening fracture and thus produces a larger spacing than a material with a higher Young's modulus. This relation becomes disturbed in Figure 7d and e, where shear fractures start to dominate the pattern. Spacing of shear fractures does not show a simple relation with increasing Young's modulus. Figure 7f shows the stress-strain relationship for the simulation presented in Figure 7a, and Figure 7g the stress-strain curve for the simulation presented in Figure 7e. These curves illustrate that the rheological behaviour of the aggregate also changes significantly with increasing elastic modulus. Failure due to the development of extension fractures shows significant differences in the stress-strain

(a)

't'

,.)

q

(b)

I n c r e a s e in a r e a

Figure 8 shows three simulations with an area increase of the deformation box. The simulations have horizontal and vertical tensile boundary conditions where the horizontal component is larger in Figure 8a and b. Figure 8a shows a simulation with a large differential stress where the horizontal stress component is three times as large as the vertical component. Figure 8b shows a simulation where the vertical stress component is about 70% of the horizontal stress component. Both extension components in the vertical and horizontal direction are identical in the simulation shown in Figure 8c. The fracture pattern in Figure 8a represents the dominant horizontal extension. A small number of horizontal fractures develop due to the vertical extension component. Figure 8b shows a different pattern where the dominance of the horizontal extension component can still be seen but the

(c)

.4

.r l

,f

relationship of the two examples and failure due to the development of shear fractures only occurs at the finite strain shown in the example of Figure 7g. Figure 7f shows a sudden drop of the tensile stress (0"2) at the beginning of extension fracture development. This indicates that the fractures develop relatively fast and grow large enough to invoke failure of the whole aggregate and reduce the tensile stress almost completely. The overall failure also reduces the differential stress. Figure 7g shows the aggregate with the highest Young's modulus in the presented sequence. It experiences a more continuous development of extension fractures that grow progressively while the aggregate reaches higher extensional strains. Therefore the tensile stress is released gradually and the differential stress continues to increase after extension fracture development.

i

9

,-r.

'~

"~

,~

~

..'

,'3-

(.

r

"1 (

>i

"],

r , N .tl

-is

. ,," ,J

"'t"

", J !_~

~

......... ]

Fig. 8. Three simulations with two extension directions and an increase in modelling area 9(a) Dominant horizontal extension produces well-oriented mode I fractures. (b) Both extension components become more similar so that fractures start to curve 9Orientation of fractures becomes more diffuse than in (a). (c) Horizontal and vertical extension components are the same, no relation of the orientation of fractures and the model boundaries can be observed. Fractures form polygons similar to mud cracks.

FRACTURE AND VEIN PATTERNS vertical extension component also influences fracture development. Fractures start to curve and lose a dominant orientation with respect to a principal stress direction because the stresses are similar. Fractures still produce a distinct spacing. In Figure 8c the stress field has no dominant extension direction and the developing fracture pattern shows no preferred orientation with respect to the simulation boundaries. Fracture development and orientation is controlled by the initial noise in the system. Once the first fractures develop they influence the stress field in the aggregate, which leads to a polygonization with a distinct spacing similar to the structures found in mud cracks in shrinking sediments (Suppe 1985).

Gravitational loading and pure shear deformation In a normal tectonic environment one would expect that sedimentary sequences are loaded during their deposition and then experience

,ool

19

tectonic strain afterwards (Price & Cosgrove 1990). In order to investigate the effects of initial gravitational loading on fracture and vein patterns several simulations were performed with complex deformation histories including different tectonic settings 9 In the first simulation with a two-stage deformation history we load the sample uniaxially and then apply a pure shear boundary condition. The sample is loaded vertically with fixed sidewalls. The vertical stress is proportional to the deformation steps and the horizontal stress is compressive due to Poisson effects. After a given amount of vertical strain we apply a pure shear deformation with vertical constrictional strain and horizontal extensional strain. The developing fracture patterns are shown in Figure 9a in three successive steps. Figure 9b shows the stress-strain diagram for the simulation and Figure 9c the corresponding Mohr diagram with four different Mohr circles a to 6 that are indicated in the stress-strain diagram of Figure 9b. The region on the left-hand side

(c)

(b)

't

T ~5

8

(+)

(+)

-t00 0

o0.005

0.01

Fig. 9. Complexdeformation history involving a compaction event followed by pure shear deformation. (a) Three successive fracture patterns during the simulation. Conjugate mode II fractures dominate the pattern. Only at the latest stage do small mode I fractures develop in the model in order to accommodate local tensile strain. (b) Stress-strain relation of the simulation. Both principal stresses are compressive during the compaction (regime L). Once tectonic deformation starts (regime T) the horizontal stress is relaxed but the differential stress increases. (c) Mohr circle diagrams for (b). The diagram on the left-hand side marks three successive stages during the compaction process. The diagram on the right-hand side shows the successive growth of the Mohr circle during tectonic deformation and the following failure that produces the dominant conjugate mode lI fractures. Note that the x-axis is negative towards the right-hand side since compressive stresses are negative in this paper.

20

D. KOEHN ET AL.

of the vertical line in Figure 9b indicates the gravitational loading regime (marked L for loading). The two principal stresses are compressive and the mean stress and differential stress increase successively. The development of the Mohr circle during initial loading is indicated in Figure 9c on the left-hand side where c~ marks the initial starting point and/3 and y the successive growth and movement of the Mohr circle into the compressive regime. After gravitational loading pure shear deformation shows an increase in compressive stress and decrease in tensile stress (Figure 9b in tectonic regime marked T). The mean stress stays constant and the differential stress increases with a steeper slope than in the loading regime. At point 6 in Figure 9b failure is initiated so that in Figure 9c the Mohr circle cuts the failure envelope. Since differential stress and mean stress are high the developing fractures are shear fractures with typical conjugate sets, which can be seen in Figure 9a. The Mohr circle will not reach the tensile part of the failure curve without crossing the failure curve in the compressive regime if pure shear deformation is applied after the uniaxial loading. The developing pattern will mainly consist of conjugate shear fractures. Small extension fractures will only grow at a larger amount of deformation in order to accommodate local tensile strains as can be seen in Figure 9a. However, if fluid pressure is involved, the Mohr circle may cross the failure curve in the tensile regime during deformation (see discussion). Gravitational loading and area increase

In this section we take a look at fracture patterns that develop during different kinds of initial loading followed by a tectonic deformation that is characterized by a large area increase so that both major components of the strain in the modelled plane are extensional. The developing fracture patterns and corresponding stress-strain curves are shown in Figure 10a to d. Figure 10a shows a simulation with a small amount of initial loading. Tectonic deformation is dominant with a large horizontal extension component. During extension the compressive stresses are relaxed and the horizontal stress component becomes tensile, which results in the development of vertical extension fractures that are opening. Failure of the whole aggregate is fast, which is illustrated by the sawtooth stress curves in Figure 10a. The material behaves in a brittle manner and shows a strong localization of strain in two to three tensile fractures that are opening and that show a distinct spacing.

Following failure the aggregate reaches a quasi steady-state plastic behaviour where stresses remain at relatively constant values while the strain increases. Small fluctuations in the tensile stress represent growth of secondary extension fractures to accommodate strain. Figure 10b shows a simulation where gravitational loading is dominating. Failure of the aggregate starts during initial loading when a peak differential stress is reached. During this stage conjugate sets of shear fractures develop. Successive tectonic deformation reduces all stress components and leads to a local growth of small extension fractures. The resulting fracture pattern is very similar to that of Figure 9, where a smaller amount of initial loading was followed by a pure shear deformation. Both simulations show a dominant primary growth of conjugate shear fractures and a successive development of secondary small-scale extension fractures. An intermediate pattern between the simulations shown in Figure 10a and b is shown in Figure 10c. Here the aggregate is initially loaded by an intermediate amount so that initial loading itself does not induce failure. The aggregate is then extended in the horizontal and vertical direction where the horizontal extension component is dominant. During failure of the aggregate, conjugate shear fractures develop with a smaller angle towards the principal compressive stress direction than in Figure 10b. The fractures in Figure 10c are of the hybrid extension/shear type of Price & Cosgrove (1990), an intermediate type of fracture. In the simulation shown in Figure 10c these conjugate fractures accommodate successive amounts of extensional strain and are opening. Figure 10d shows a different type of gravitational loading where the aggregate is initially loaded by the same horizontal and vertical amount. Therefore the differential stress stays at zero and the two principal stress components as well as the mean stress increase by the same amount. The following tectonic deformation has a horizontal extension component. During extension all stresses are relaxed. The differential stress starts to increase as a result of the large horizontal extension component. This resulted in the growth of conjugate hybrid extension/ shear fractures similar to those shown in the simulation of Figure 10c. In summary gravitational loading can lead to the development of primary conjugate sets of hybrid extension or shear fractures that will accommodate extensional strain during successive deformation. It can be seen therefore that fractures and veins in a rock have a record of

FRACTURE AND VEIN PATTERNS

(a)

'LJ;............. ,.i' ;

}

,

(b)

,. '...

,

0

(MPa L

21

r

100 o (

o

o

~o21

AO

zc

O

/

o~

o~

i

lOO

-20

0

0.02

0.04

-20

13

01

0

0.01

E

-40

-40 0

0.01

0.0Z s

0

O,OOS

s

Fig. 10. Four different simulations and the corresponding stress-strain curves for deformation histories involving different kinds of compaction that are followed by tectonic deformation that involves extension in the modelling plane. (a) Small amount of compaction results in a situation where mode I fractures are dominant. (b) Large amount of compaction results in the development of conjugate mode II fractures during compaction. The following tectonic extension produces local mode I fractures but strain is mainly accommodated by the already existing fracture network. (c) A simulation with intermediate compaction followed by tectonic extension results in the growth of hybrid extension/shear fractures that are opening. (d) Compaction that is hydrostatic can also lead to hybrid extension/shear fractures if the following tectonic deformation is strongly non-hydrostatic so that extension is dominating in the horizontal direction.

22

D. KOEHN E T A L .

the gravitational loading of the rock prior to tectonic deformation.

Discussion We compare the results of the numerical study with the vein sets shown in Figure 2. The vein in Figure 2a may have started as a hybrid extension fracture and may reflect a stress state that would produce that type of fracture. This stress state would then be represented by the whole history of the vein development. However, our simulations show that quite a number of different stress states may develop during a deformation history. Therefore this vein could also form as a shear fracture (or hybrid shear depending on its orientation relative to the principal stress axis) initially and then open afterwards due to an extension component during deformation. Such a setting could also produce vein patterns similar to the ones shown in Figure 2b where conjugate shear (or hybrid shear) fractures may form first and then later extension takes over and the initial fractures are opening as veins. However, further study on the internal geometries of such vein sets is needed in order to prove that the proposed scenario can develop in nature. We use hyperbolic Mohr circle diagrams to illustrate the stress states and failure in our simulations. There is still much debate on whether a hyperbolic Mohr circle can be used for intermediate failure types between mode I and mode II, namely hybrid extension/shear fractures. Engelder (1999) illustrates that these intermediate structures cannot be explained as propagating cracks by linear elastic fracture mechanics theory. He rather concludes that they may form by an out-of-plane propagation of mode I fractures that are subject to a shear stress. In our simulation a number of scenarios produce patterns that are dominated either by mode I or mode II fractures. In intermediate settings quite often a combination of mode I and mode II fractures develop where both failure modes still show distinct sets of fractures. Only in a narrow range do we find fractures that have an intermediate orientation (between mode I and II), which fall into the category of hybrid extension/shear fractures (Fig. 10c). These may well be combinations of mode I and mode II fractures as proposed by Engelder (1999). It is interesting to see that in a number of simulations quasi steady-state behaviours are reached (Figs 5a and 10a, c). The fracture systems that develop can accommodate additional strain with only minor increase in stress, which is

probably due to additional growth of small fractures or friction along shear fractures. Prior to this quasi steady-state behaviour, stresses drop significantly when the whole systems fails. The remaining tensile stress component is very small since most tensile stresses can be accommodated by opening fractures or veins. Compressive stresses are higher and can probably be related to friction along shear fractures as long as one component of the stress tensor in the modelling plane is compressive. As long as the system does not heal after failure stresses will probably remain low. If this steady state is also reached in natural rocks, it has a strong influence on the strength of brittle faults and earthquake behaviour (Scholz 2002). Once a rock fails by fracturing it may deform in a plastic way without significant increase in stress (Paterson 1978; Zhang et al. 1990; Karato 1995). Once a fault forms and develops a cataclasite it may behave as a plastic material with no major increase in stress if fractures are not healing. Therefore no successive earthquakes will be expected along such a fault. However, if veins grow and fill the fractures, the cataclasite may heal and retain a certain strength again. Then it may fracture by catastrophic failure at high stresses (Scholz 2002). What we have not yet taken into account is the effect of fluid pressure on the developing fracture patterns. This is certainly an important component that we will include in the future. A high fluid pressure can shift the Mohr circle towards the tensile regime so that mode I fracture may develop during the loading history instead of conjugate shear fractures (Price & Cosgrove 1990; FlekkCy et al. 2002). However, conjugate shear fractures may also develop in this case if the differential stress is already high due to a non-hydrostatic initial loading. Fluid pressure and developing fractures will influence each other depending on how effectively fractures can drain an existing fluid pressure. If the tectonic stress is non-hydrostatic or if fluid pressure gradients exist along anisotropies, fluid pressure will provoke failure (Bercovici et al. 2001). Once fractures develop, the fluid pressure may drop because the permeability of the rock increases. Whether or not the fluid pressure influenced fracture and vein development in the natural samples that were presented in this paper is not clear. Bedding parallel veins in the area may have formed during high fluid pressures because they are located along the bedding planes between porous sandstones and less porous mudstones. Fluid pressure gradients can be expected along these anisotropies, which may invoke failure. The veins that are presented

FRACTURE AND VEIN PATTERNS in this paper are however located within the more porous sandstones and are probably more associated with tensile strain than a higher fluid pressure. However, additional research is needed in order to fully understand the effect of fluid pressure on the development of these systems. In the following paragraphs we will discuss some additional boundary effects of the model that we used. The box boundaries have an influence on the development of shear fractures. Box boundaries reflect shear fractures because the boundaries behave like rigid pistons and boundary particles are fixed perpendicular to the walls. They cannot accommodate a shear displacement of the boundary so that shear fractures that hit the walls are repelled and a new shear fracture grows back into the model. This effect will influence the density of shear fractures in the model and also their spacing. In some cases shear fractures will dominate near to the compressive walls (upper and lower wall) whereas extension fractures dominate in the centre of the model (Fig. 7). However, this effect is not always observed (Fig. 10c). An additional problem arises because the underlying lattice is triangular. Therefore fracturing is partly anisotropic, which is especially important in the case of shear fractures where slip should take place. The anisotropies of the lattice can be overcome by adding distributions to the elastic constants and spring breaking strengths that are larger than the lattice anisotropies (Malthe-SCrenssen et al. 1998b). Mode I fractures show no lattice-preferred orientation in our model, which can be best observed in Figure 8c where stresses are hydrostatic and no lattice directions nor any effects of the model boundaries are observed. Shear fractures may however still be influenced by the lattice (Fig. 10b) but can also develop in non-lattice directions (Fig. 10c). One has to note that the direction of the shear fractures in Figure 10b is more or less a lattice direction, but also the direction where the fractures would develop anyhow. The threshold of relaxation in the model can also change the developing pattern. This is recorded in detail in Malthe-SCrenssen et al. (1998b). Since we are moving all particles assuming homogeneous strain in the model before relaxation starts and since our relaxation threshold is very small relative to the applied strain, it has no major influence on the developing pattern. However, if a significantly higher relaxation threshold is used, different spacing of mode I fractures will result since the system will be progressively slower and relaxation will be more localized.

23

Conclusions Modelling of progressive fracture development in aggregates shows that different type of fractures can grow during one tectonic deformation event. Pure shear deformation produces primary extension fractures that are followed by a secondary set of conjugate shear fractures. These patterns change depending on the properties of the aggregate, namely its mean elastic constant and the initial noise in the system. Pure extension of the modelling area leads to one dominant extension fracture set or to a polygonal set of fractures if the principal stresses are equal. A more complex deformation history that includes gravitational loading as well as tectonic strain can lead to the development of conjugate shear fracture sets that can accommodate extensional strain and thus form veins during later stages of deformation. These conjugate sets of veins may be followed by secondary extension fractures. These complex veins may record gravitational loading preceding the tectonic deformation. Their orientation can be used to find the orientation of the non-hydrostatic stress field that existed during gravitational loading or tectonic deformation. We thank T. Engelder and J. Cosgrove for constructive reviews. We also thank A. Malthe-SCrenssen for his help with the discrete element code. JA acknowledges funding by the DFG-Graduiertenkolleg 'Composition and Evolution of Crust and Mantle'.

References ALLEN, D. M. DE G. 1954. Relaxation Methods. McGraw Hill, New York. ANDERSON,E. M. 1951. The Dynamics of Faulting and Dike Formation. Oliver and Boyd, Edinburgh. AYDIN, A. & RECHES, Z. 1982. Number and orientation of fault sets in the field and in experiments. Geology, 10, 107-112. BERCOVICI, D., RICARD, Y. & SCHUBERT, G. 2001. A two phase model for compaction and damage, 3: Applications to shear localization and plate boundary formation. Journal of Geophysical Research, 106, 8925-8942. BONNET, E., BOUR, O., ODLING, N. E., DAVY, P., MAIN, I., COWIE, P. & BERKOWITZ, B. 2001. Scaling of fracture systems in geological media. Reviews in Geophysics, 39, 347-383. ENGELDER, T. 1999. The transitional-tensile fracture: A status report. Journal of Structural Geology, 21, 1049-1055. FLEKK~3Y, E. G., MALTHE-S~3RENSSEN, A. & JAMTVEIT, B. 2002. Modeling hydrofracture. Journal of Geophysical Research B8, ECV 1, 1- 11. GRI~ITH, A. A. 1920. The phenomena of rupture and flow in solids. Transactions of the Royal Socie~ Series A, 221, 163-198.

24

D. KOEHN ET AL.

HANCOCK, P. L. 1985. Brittle microtectonics: principles and practice. Journal of Structural Geology, 7, 437-457. HUDSON, J. A. 8z COSGROVE,J. 1997. Integrated structural geology and engineering rock mechanics approach to site characterization. International Journal of Rock Mechanics and Mining Science, 34, 577. JAEGER, J. C. 8z COOK, N. G. W. 1976. Fundamentals of Rock Mechanics. Chapman and Hall, London. JAMTVEIT, B., AUSTRHEIM, H. 8z MALTHESORENSSEN, A. 2000. Accelerated hydration of the Earth's deep crust induced by stress perturbations. Nature, 408, 75-78. JESSELL, M. W., BONS, P. D., EVANS, L., BARR, T. D. & STt3WE, K. 2001. Elle, The numerical simulation of metamorphic and deformation textures. Computers and Geosciences, 27, 17-30. KARATO, S. 1995. Rock deformation: Ductile and brittle. Reviews in Geophysics, American Geophysical Union, 33. KRANTZ, R. W. 1988. Multiple fault sets and threedimensional strain; theory and application. Journal of Structural Geology, 10, 225-237. MALTHE-SORENSSEN,m., WALMANN,T., JAMTVEIT,B., FEDER, J. • J~SSANG, T. 1998a. Modeling and characterization of fracture patterns in the Vatnajokull glacier. Geology, 26, 931-934. MALTHE-SORENSSEN, A., WALMANN, T., FEDER, J., JOSSANG, T. & MEAKIN, P. 1998b. Simulation of extensional clay fractures. Physical Review E, 58(5), 5548-5564. MALTHE-SORENSSEN,A., WALMANN,T., JAMTVEIT,B., FEDER, J. 8z JOSSANG, T. 1999. Simulation and characterization of fracture patterns in glaciers. Journal of Geophysical Research, B104, 23157-23174. MEANS, W. D. 1976. Stress and Strain: Basic Concepts of Continuum Mechanics for Geologists. SpringerVerlag, New York.

OERTEL, G. 1965. The mechanism of faulting in clay experiments. Tectonophysics, 2, 343-393. OLIVER, N. H. S. & BONS, P. D. 2001. Mechanisms of fluid flow and fluid-rock interaction in fossil metamorphic hydrothermal systems inferred from veinwallrock patterns, geometry and microstructure. Geofluids, 1(2), 137. PATERSON, M. S. 1978. Experimental Rock Deformation - Brittle Field. Springer Verlag, Berlin. POLLARD, D. D. & SEGALL, P. 1987. Theoretical displacements and stresses near fractures in rocks: with applications to faults, joints, dikes and solution surfaces. In: ATKINSON, B. K. (ed) Fracture Mechanics of Rock. Academic Press, London, 277-348. PRICE, N. J. & COSGROVE, J. W. 1990. Analysis of Geological Structures. Cambridge University Press, Cambridge. RAMSAY, M. R. & CHESTER, F. M. 2004. Hybrid fracture and the transition from extension fracture to shear fracture. Nature, 428, 63-66. RAMSAY, J. G. & HUBER, M. I. 1983. The Techniques of Modern Structural Geology, 1: Strain Analysis. Academic Press, London. RANALLI, G. 1995. Rheology of the Earth, 2nd edn. Chapman & Hall, London, UK. SCHOLZ, C. H. 2002. The Mechanics of Earthquakes and Faulting, 2nd edn. Cambridge University Press, Cambridge. SUPPE, J. 1985. Principles of Structural Geology. Prentice-Hall, New Jersey. WALMANN, T., MALTHE-SORENSSEN, A., FEDER, J., JOSSANG, T., MEAKIN, P. ~r HARDY, H. H. 1996. Scaling relations for the lengths and widths of fractures. Physical Review Letters, 77, 5292-5296. ZHANG, J., WONG, T.-F. & DAVIS, D. M. 1990. Micromechanics of pressure-induced grain crushing in porous rocks. Journal of Geophysical Research, 95, 341-352.

Segmentation and interaction of normal faults within the Colfiorito fault system (central Italy) F. M I R A B E L L A , V. B O C C A L I & M. R. B A R C H I Geologia strutturale e Geofisica, Dipartimento di Scienze della Terra, Universith di Perugia, piazza Universith 1, 06100 Perugia, Italy (e-mail: mirabell@ unipg, it) Abstract: Fault segments belonging to a fault population can link and interact, eventually

forming a single larger fault, and thus affecting the estimation of the maximum expected earthquake. We present throw distribution data along the Quaternary normal faults of the Colfiorito fault system (central Italy), which consists of four main fault segments and where a seismic sequence occurred in 1997-1998. Throw values along the two central overlapping, en-rchelon segments (8.5-9.5 km long) were measured on a good stratigraphic marker, by constructing a set of closely spaced geological cross-sections, perpendicular to the fault strike. As these faults are commonly retained active and border Quaternary basins, we compare morphological and geological throws in order to verify the faults neotectonic activity. Geological and morphological throw distributions show good correlation, testifying that recent faulting affected the topographic surface and suggesting that the observed offset completely accumulated during the Quaternary. The throw distribution along the fault segments is asymmetric and reaches maximum values (500-550m) within the zone of fault overlap, suggesting mechanical interaction between the studied faults. Maximum length-throw correlation suggests that the studied faults grew according to a linear scaling relationship.

Faults are usually organized in systems of segments, which may interact and link forming a single fault (e.g. Anders & Schlische 1994; Cartwright et al. 1995). Fault interaction and geometrical irregularities between individual faults play an important role in the initiation and termination of seismic ruptures and in the distribution of aftershocks (e.g. Aki 1979; King & Nabelek 1985; Sibson 1986; Barka & Kadinsky-Cade 1988; Sieh et al. 1993; Amato et al. 1998; Chiaraluce et al. 2003). Fault discontinuities and interaction are also important in constraining fault growth conceptual models that consider either single continuous fault planes that increase their displacement as they grow (Watterson 1986; Walsh & Watterson 1988; Marrett & Allmendinger 1991; Cowie & Scholz 1992) or fault growth by segment linkage (Segall & Pollard 1980; Burgmann et al. 1994; Willemse et al. 1996). The primary expression of interaction between fault segments is on the way they grow and is reflected by the shape of the displacement profiles that show asymmetric slip distribution along strike (Peacock 1991; Gupta & Scholz 2000). The irregular way in which faults grow has also been attributed to the presence of

heterogeneities in shear strength but this effect is supposed to be of less importance than that due to fault interaction (Peacock & Sanderson 1991). This has also been suggested by numerical modelling, which indicates that mechanical interaction between neighbouring faults, even if not connected, can cause asymmetric slip distribution (Willemse et al. 1996). As fault dimensions are important for seismic hazard purposes (Wells & Coppersmith 1994), the interaction between adjacent faults affects the estimation of the maximum possible earthquake. Hence the possibility that two or more fault segments can interact to form a larger fault becomes a matter of interest, especially when studying active fault systems in populated areas. This paper is dedicated to the Colfiorito Fault System (CoFS) in the Umbria-Marche region (central Italy) where a seismic sequence characterized by a series of moderate magnitude earthquakes occurred in 1997-1998. We obtained the fault throw distribution along the two central segments of the fault system, where recent surface mapping (Barchi et al. 2001) provided a detailed fault pattern. From surface data, a set of 500 m spaced geological cross-sections oriented perpendicularly to fault strike was

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 25-36. 0305-8719/05/$15.00 ,:~ The Geological Society of London 2005.

26

F. MIRABELLA E T A L

constructed in order to draw the along-strike distribution. As recent fault activity is commonly supposed to affect the topographic surface (e.g. Burbank & Anderson 2001), we compare the vertical component of displacement (geological throw) with the difference between the maximum and minimum topographic elevations (morphological throw), measured at the footwall and hangingwall respectively, in order to discuss the neotectonic fault activity. Finally, by plotting maximum throw versus length for most of the normal faults cropping out in the investigated area, we derive a scaling law for the fault growth.

Geological setting The Colfiorito Fault System belongs to the active alignment of normal faults (Umbria Fault System - UFS, Barchi 2002) of the UmbriaMarche Apennines (Fig. la). The UFS faults border the most important N W - S E trending Quaternary continental basins (Gubbio, Colfiorito, Norcia, Cascia and Castelluccio basins).

The faults of the UFS are commonly referred to as active faults, on the basis of geomorphological (e.g. Ficcarelli & Mazza 1990; Coltorti et al. 1998; Messina et al. 1999), geological (e.g. Lavecchia et al. 1994; Calamita et al. 1999; Barchi et al. 2000; Boncio & Lavecchia 2000; Mirabella & Pucci 2002) and seismological evidence (Deschamps et al. 1984; Haessler et al. 1988; Amato et al. 1998; Ekstroem et al. 1998; Mariucci et al. 1999; Stramondo et al. 1999; Barba & Basili 2000). Historical seismicity (maximum intensity ---- 10 of the modified Mercalli scale, MercalliCancani-Sieberg intensity scale, MCS, CPTI 1999) and recent earthquakes (Norcia, 1979 Ms = 5.5; Gubbio, 1984 Ms = 5.3; Colfiorito, 1997-1998 Mw ~ 6.0; Deschamps et al. 1984, 2000; Haessler et al. 1988) indicate that the presently active stress field (extensional with S W - N E trending o-3) is consistent with the geological long-term stress field obtained by structural data and active since the Quaternary (Lavecchia et al. 1994; Mariucci et al. 1999). The Colfiorito Fault System is presently active. The area was recently struck by a long

Fig. 1. (a) Schematic structural map of the Umbria-Marche region showing the alignment of the intramontane basins along the Umbria Fault System (UFS). Historical seismicity is reported for the period 461 BC to AD 1979 (Boschi et al. 1997). Focal mechanisms and magnitudes are for the 1997-1998 Colfiorito sequence (Ekstroem et al. 1998), for the 1979 Norcia earthquake (Deschamps et al. 1984) and for the Gubbio earthquake (Dziewonski et al. 1985). (b) Schematic pattern of the Colfiorito Fault System and location of the four main segments.

SEGMENTATION AND INTERACTION OF NORMAL FAULTS 1997-1998 seismic sequence characterized by six main earthquakes with 5 < Mw < 6 all nucleated on SW-dipping normal faults (Chiaraluce et al. 2003). On 26 September 1997, two moderate magnitude main shocks (Mw = 5.8 and Mw = 6.0) occurred in a short time period (9 hours), rupturing in opposite directions and with hypocentres located at a distance of about 2 - 3 km from each other (Pino & Mazza 2000). At the surface, the pattern of the CoFS faults is complex. Several fault segments mainly strike about N135 ~ and dip to the SW, dips ranging from 50 ~ to 80 ~ Subordinate NS, EW and N E - S W normal faults are also present in the area. The faults dissect pre-existing compressional features related to the Umbria-Marche fold and thrust belt and some of them border the eastern side of the intramontane Pleistocene basins (Colle Croce-Annifo, Colfiorito, Cesi, S. Martino). In particular, the Colfiorito basin consists of lower-middle Pleistocene fluvio-lacustrine deposits (Ficcarelli & Mazza 1990; Coltorti et al. 1998) that reach a thickness of about 120 m in the depocentre (Messina et al. 1999). Despite the complexities of the fault arrangement, with patterns of overstepping synthetic and antithetic splays at different scales of observation, four main fault segments with greatest continuity are recognized, from North to South: M. Pennino, M. LeScalette, Cesi-S. Martino and M. Cavallo-M. Fema (Fig. lb). These main segments have similar characteristics, that is, they strike N W - S E , have similar along-strike length (ranging from 8 to 10 kin) and similar maximum throws of up to 550-600 m (Barchi et al. 2000; Mirabella & Pucci 2002). A structural sketch (Fig. 2a) of the two central segments (M. LeScalette and Cesi-S. Martino) is drawn from the recent detailed geological map (1:10 000 scale) of Barchi et al. (2001). The M. LeScalette segment comprises a major fault and subparallel faults, strikes ranging from N130 ~ to N140 ~ (Mirabella & Pucci 2002). Several minor faults are also present, strikes varying between N120 ~ and about N160 ~ The most continuous fault plane borders the Colfiorito basin and is in the northwestern part of the segment. It puts in contact debris deposits in the hangingwall with Cretaceous limestone (Maiolica fro, tithonian-aptian) in the footwall. Along this fault, 2 to 20 cm high free faces were locally exposed and interpreted by some authors as due to coseismic reactivation during the 1997 earthquake (Cello et al. 1998; Calamita et al. 1999). The Cesi-S. Martino segment, about 8.5 km long, is characterized by a greater variability in

27

the fault attitude (Fig. 2a). From NW to SE, it comprises the following three portions: the first that borders the Cesi basin, strikes about N130 ~ with dips in the range 55-60~ the second that borders the S. Martino basin, strikes about NI70 ~ with dips of 62-74~ the third strikes N150 ~ and dips about 60 ~ The Cesi-S. Martino segment is shifted about 3 km to the SW with respect to the M. LeScalette segment (Fig. 2b). The overlap zone is characterized by a complex pattern of SW dipping and subordinate NE dipping normal faults with strikes ranging from NI20 to N170 ~ On the basis of striated fault planes, the normal faults are prevalently dip-slip, although some planes possess a strike-slip component (Fig. 2c). The inversion of fault slip data, although lacking a clear antithetic system, indicates a roughly vertical 0-1 and S W - N E trending 0-3 (Barba & Basili 200O).

Throw profiles along the Colfiorito Fault System The throw values of the normal faults in the study area were measured along a set of 27 geological cross-sections through the fault segments. The cross-sections, numbered from SI to $27 (Fig. 2a), are oriented N55 ~ with a mean spacing of about 500 m and are perpendicular to the average fault strike. The Cesi-S. Martino segment is characterized by the presence of a single fault ( N i l ) , about 8.2 km long, on which all the measured throw has accumulated, whereas for the M. LeScalette segment the total throw is distributed on six smaller faults (N2, N3, N6, N7, N8, N9 in Fig. 2a). Considering the dextral en-~chelon array of the normal faults (Fig. 2a), the group of sections S1 to S12 shows the geometry of the M. LeScalette segment, while the group of sections S12 to $24 displays the geometry of the Cesi-S. Martino segment. The overlapping area between the two major fault segments is imaged by section S12 (Fig. 2a). Sections $25, $26, and $27 have been drawn in order to measure the throw distribution of small faults in the southernmost region, which are thought to have been activated during the 1997-1998 earthquake (Basili & Meghraoui 2001). Three sections, representative of the structural style of the CoFS, are represented in Figure 3a. Section $9 (Fig. 3a) shows that the hangingwall of the M. LeScalette faults comprises gently folded Cretaceous limestones and marls involved in pre-existing compressional structures (thrusts, transpressive faults and associated folds). The

28

F. M I R A B E L L A ET AL.

Fig. 2. (a) Schematic structural map of the study area derived from 1 : 10 000 mapping scale with intramontane basins and mapped faults. The studied normal faults are numbered from 1 (N1) to 15 (NI5). Thin lines labelled from S1 to $27 show the location of the 27 geological cross-sections used for the measurment of geological and morphological throws. (b) Stereographic projections (Schmidt equal area projection, lower hemisphere) of fault slip data measured along the main faults' scarps and along minor faults in the study area (after Barba & Basili 2000). (c) Schematic pattern of the main fault segments studied in this work and parameters of length and separation.

SEGMENTATION AND INTERACTION OF NORMAL FAULTS

Stratigraphy SW Oligocene __qZ~------~ middle-Eocene LowerTuronian I I [ upperAlbian ~ ~/11 Aptian-Albian ~ ~r~ LowerAptian~---~----~ UpperTithonian 0 Middle Tithonia~ I Toarcian

29

M.LeScalettefaults N7 N3 | N8 NE

i

0

SW .

1

Cesi S.Martinofault ......... " a~ ' %

M.LeScalettefaults

Nil

-NI0

---.. 0

-.. . . . . . . . . . . . .

,

0

N7

0

.x ".

.'" .-..

o

S

~

.

NE ,

~ .... .

0

Cesi S.Martinofault~ll NEI ~ I

S17~

2km

I]

@

Fig. 3. (a) Geological cross-sections ($9, S12, S17) across the M. LeScalette and Cesi-S. Martino segments. Section $9 crosses M. LeScalette faults, section $12 crosses both M. LeScalette and Cesi-S. Martino faults, and section $17 shows the Cesi-S. Martino fault geometry. Location of the sections is given in Figure 2a. (b) Umbria-Marche stratigraphy with the position of the Marne a Fucoidi marker.

total throw achieved by the M. LeScalette faults reaches about 500 m. Section S12 crosses both the southeastern termination of fault N7, belonging to the M. LeScalette segment, and the northwestern termination of fault N l l , which belongs to the Cesi-S. Martino segment. Section S17 shows the geometry of the central part of fault N l l where the throw is about 550 m. The throw distribution along the M. LeScalette and Cesi-S. Martino faults was measured along the geological cross-sections, from the offset of a continuous stratigraphic marker (Marne a Fucoidi formationM F fm, Aptian-Albian) which is a marly formation (Fig. 3b) well recognisable within the well-known Umbria-Marche carbonatic succession (Cresta et al. 1989). Besides, the stratigraphic setting of the area allowed us to reconstruct the geometry of the MF marker by stratigraphic correlation where it does not crop out. The M. LeScalette segment has a maximum length of 9300 m and an asymmetric throw distribution with a maximum value of about 500 m about 3 km SE of the segment centre (Fig. 4). The Cesi-S. Martino segment has a maximum length of about 8200 m, and comprises a single major fault ( N l l ) with an asymmetric throw distribution. The maximum throw of about 550 m occurs at about 0.7 km to the

NW of the segment centre corresponding to section $18 (Fig. 4).

Comparing geological and morphological throws In tectonically active areas, faulting commonly offsets the topographic surface. Normal faults with nearly pure dip-slip kinematics, such as the CoFS faults, are visible in the landscape. The sum of the topographic offset and of the depth of basin infill (morphological throw) is often considered to be comparable with the long-term geological throw achieved by active normal faults (e.g. Roberts 1996; Coltorti & Pieruccini 2000; Burbank & Anderson 2001; Pizzi et al. 2002). This is especially valid in regions of low erosion rates (Doglioni et al. 1998). As the recent activity of these faults is recorded by variations in topographic elevations, it is expected that geological and morphological throws display similar values. This approach can be applied to the CoFS faults that are of Quaternary age and that differ from other normal faults of central Italy that reactivated pre-existing normal faults (Calamita et al. 1998; Pizzi & Scisciani 2000; Pizzi et al. 2002). Reactivation of a pre-existing normal fault was also suggested for the Gubbio

30

F. MIRABELLA 200

vw"

'

"

'

"

'

ET AL.

"

'

"

'

"

'

"

'

5 ~

S 1 ~ 1 5 LeScalettesegment

~ 1 5 o / ~ N l O f a u l t .

Cesi-SMartino

segment

511t) 1000 1500 20(x) 2500 3000 3500 4(K,Ud faultlengthtrn) 600 INW"

I

"

I

"

I

'

I

'

!

'

I

"

I

"

I

"

I

'SE

600 NW"

[.

"~

I

9

I

"

.

segment

r ~"F ~

St~

; 4001"-

... S~/

Itotal geological

~300~ S" ~

_

throw

3

~."

""

~ /~

~

-

c, ~ ~"F ~.

/

I

,~E400L

$17

. ~

,

~

sl

"~-- S~4"~N"N

~

~

I

9

S~ /

i

"

I

9

I

9

I

9

I

" SE

.

\

Cesi-S.Martino segment \

\ total geological

-

throw

.

~19

/

".~ .,~300I

io

I

s~7~s

I

~

.

$1r

-

N1~"~~23

9

.

~200

~l

-0

I1 - ~200

I

]

~

I

~

I

~ 1 ~

~ 3 ~

[ ~12

1000 2000 3000 4000 5000 6000 7000 8000 c ~ fault length (m)

24

~l

100013

0

I

I

I

[

~

am

l

[

l

[

I

I

I

I [

l

1000 2000 3000 4000 5000 6000 7000 8000 9000 fault length (m)

Fig. 4. Throw distribution along the studied fault segments and relative position of the fault segments in map view (top-left).

normal fault NW of Colfiorito (see location in Fig. la), from good quality seismic reflection profiles (Mirabella e t al. 2004). In such circumstances, erosion erased a considerable part of the previously achieved morphological throw, giving morphological throw values that are smaller than the geological throw. For the M. LeScalette and C e s i - S . Martino segments of the CoFS, we measured the morphological throw on 1 : 10 000 maps along the trace of the geological sections that we used to measure the geological throw values. In the Colfiorito area, the topographic throw is a good approximation to the morphological throw (topographic throw + basin thickness) as the intramontane basins are a few tens of metres deep. The deepest basin, Colfiorito, reaches a maximum depth of about 120 m only in its central part. Hence, the values of the morphological throw are given by the difference between the maximum elevation of the fault footwall and the minimum elevation of the hangingwall (Fig. 5). Morphological and geological throw curves (Fig. 6) are quite similar. The two are asymmetric with the highest values shifted toward the nearby fault segment. The similarity in both shape and values suggests the observed longterm geological throw accumulated during the

Quaternary. From the age of the intramontane basins (about 1.8 Ma), it is possible to calculate the average slip rate by dividing the faultparallel offset (displacement) by the duration of faulting (1.8 Ma). Assuming a mean fault dip of 65 ~ and nearly pure d i p - s l i p kinematics, a

SW

NE

-...

I

o ~

I GT.~] I I

Marne a Eucoidi hn

....

"

~

I

Fig. 5. Schematic representation of the geological and morphological throws measured on the cross-sections. In this area, the morphological throw can be considered as equivalent to the topographic offset (difference between the maximum and minimum elevations at faults' footwall and hangingwall respectively), as the basins infilment is of few metres, reaching 120 m only in the depocentre of the Colfiorito basin (Messina et aL 1999).

SEGMENTATION AND INTERACTION OF NORMAL FAULTS 600

NIC

I

'

I

"

I

"

I

'

I

'

I

"

I

"M.LeScalettesegment

5OO

~ 4OO

.

"

I

''

I

"SE

] 600

'VW'

~

$9

'

..... ~7~

...,.,\

-

300 -

~ 200 - ~ ' ~ ~ - ~

N

/

\

;,~ "~

~1l

t -I

I ' i 9I ' I ' I " I ' I

' SEI

mort~hOolwOgl~cSal

Cesi-S.Martinosegmen; t "~#~,f7 16- "~.~tot.algeologicalthrow 1

morpholoicalthrow...*PS8 ~

-

"

31

.." ~ ' ~

f

I- .." /

;sl t 2oo~7

......

1 "~is24 t

100 0

1000 2000 3000 4000 5000 6000 7000 8000 9000 1000000 fault length (m)

1000 2 0 0 0 3000 4000 5000 6000 fault length (m)

7000

8000'9000

Fig. 6. Distribution of geological and morphological throws of M. LeScalette and Cesi-S. Martino segments.

maximum throw of about 500 m gives an average slip rate of 0.3 mm/year, in agreement with average slip rates for this portion of the Northern Apennines (<0.3 mm/year; Barchi et al. 2000).

Fault segmentation and segment interaction The maximum throws of M. LeScalette and Cesi-S. Martino segments are asymmetric, being shifted towards the SE and the NW ends, respectively (Figs 4 and 6). Such a throw distribution in en-~chelon fault systems is considered as evidence of fault interaction (Peacock & Sanderson 1991), as confirmed by elasticplastic modeling (Gupta & Scholz 2000) and numerical investigations (Willemse et al. 1996). According to Gupta & Scholz (2000), interaction between overlapping faults develops for a ratio of horizontal separation to total length lower than 15%. Similarly, an upper limit of 10% is found to control the interaction of overlapping strike-slip fault systems (An 1997). The separation between the M. LeScalette and Cesi-S. Martino en-6chelon segments is about 3.4 km and their overlap is nearly 1 km (Figs 2c and 7). Hence, the ratio of separation to total length is about 20%, which is higher than the 10-15% quoted above. Nevertheless, within the overlapping zone, several minor normal faults are present, the most important of which is fault NIO with a length of about 3.3 km and maximum throw of about 180 m (Figs 4 and 7). Fault NIO is not clearly linked to both the major faults but could likely play a connecting role between M. LeScalette and Cesi-S. Martino segments. In this view the M. LeScalette segment and the NIO fault have a ratio of separation to total length

of 12%, while the Cesi-S. Martino segment and the NIO fault have a ratio of about 9.4% (Fig. 2c). Moreover, the total geological and morphological throw distribution, from the northwestern tip of M. LeScalette segment to the southeastern tip of the Cesi-S. Martino segments, shows a decrease from about 500-550 m to about 250300 m in the interaction zone (Fig. 7a). This suggests that the NIO fault interacts both the northern M. LeScalette segment and the southern Cesi-S. Martino segment, forming a connecting structure between the two segments (Fig. 7b). This interpretation compares well with the observations of Peacock & Sanderson (1991) for the Kilve normal fault zone (Somerset, UK) and matches the interaction model of Gupta & Scholz (2000) (see inset Fig. 7c). The above evidence of fault interaction within the CoFS is important for seismic hazard assessment since it can provide inferences on the maximum expected earthquake. The focal mechanisms and the distribution of the main shocks and the aftershocks of the 1997-1998 earthquakes suggest that the M. LeScalette and Cesi-S. Martino segments were activated during the seismic crisis (Chiarabba & Amato 2003; Chiaraluce et al. 2003). The lengths of 9.3 km and 8.2 km of the two major M. LeScalette and Cesi-S. Martino fault segments are in good agreement with the magnitudes of 5.8 and 6.0 for the maximum events registered during the 1997-1998 seismic sequence according to the relationship of Wells & Coppersmith (1994). Considering the discussed degree of interaction between the two central fault segments of the CoFS, the possibility of a seismic rupture jumping from one fault segment to a nearby one, with a consequent increase in the maximum possible earthquake magnitude, cannot be excluded. In fact, by

32

F. MIRABELLA ET AL.

700]AUWi

I

'

I

'

I

'

I

'

M.LeScalette segment -9,5 km

600 U

total g e o l o g i c a l t h r o w

00L

Fmorphologicalthrow

r

400t--

"~

,,,

%1

~,, .~,~,~i~'~

200ti

0

$7/~x1~

S~'I~

t"

~

4000

(" M al ', ""--. .LeSc ette segment

'

I

'

~

1

~..l.-..~ "-

....................................

'

I

' SE

:

s 1]ltl / - -

\~

#--# /

',\~S19

--

~

-

17

0

6000 8000 10000 12000 length along faults strike (m)

................

I

Cesi-S.Martino segment ] -8,5km -s 18

Sllt\ SI2J### k ll

N

2000

I

1 linteraction zoneD, ' ' SI~

14000

16000

8000

NlOfault _............ CesiSMariin-D-segih-ei#-

...........

Fig. 7. (a) Distribution of geological and morphological throw along the M. LeScalette segment, NIO fault, and Cesi-S. Martino segment, showing a decrease of throw values in correspondence with the interaction zone. (b) Map of the fault segments. (c) Expected throw distribution in interacting normal faults' segments according to the Gupta and Scholz (2000) model.

considering the M. LeScalette and Cesi-S. Martino faults as a single fault, the empirical relationships on fault dimensions versus earthquake magnitude (Wells & Coppersmith 1994) would provide an earthquake of about M c. 6.2 in the area. Fault interaction within the Colfiorito fault system was also recently suggested on the basis of seismological data (Scholz & Gupta 2000; Chiaraluce et al. 2003). During the 1997-1998 crisis, two earthquakes of similarly moderate magnitude ( M w = 5 . 8 at 00:33 UT and

Mw = 6.0 at 09:40 UT) were registered in a short period of time and at a distance of less than 2 km (e.g. Amato et al. 1998; Chiaraluce et al. 2003). Although very close to each other, the two shocks were characterized by opposite rupture directions, that is, towards the SE and NW (Pino et al. 1999; Pino & Mazza 2000). This suggests nucleation of the seismic rupture at the border of a fault segment and triggering of a second rupture on a similar adjacent fault segment. The available data do not clearly indicate which segments of the CoFS interacted on

SEGMENTATION

AND INTERACTION

26 September 1997. However they indicate that the architecture of the system is such that mechanical interaction is likely to occur within it.

Since Menard (1962), a number of papers have been devoted to the understanding of fault growth and the scaling properties of fault displacement and length (e.g. Menard 1962; Walsh & Watterson 1988; Cowie & Scholz 1992; Dawers et al. 1993; Walsh et al. 2002). Despite the large scatter in types of data sets, there is a general agreement on the power-law relationship between displacement (D) and length (W): Wn

where c is related to material properties. The value of the exponent n has been debated. Some authors found a linear relationship (Elliott 1976; Ranalli 1977; Muraoka & Kamata 1983; Opheim & Gudmundsson 1989; Cowie & Scholz 1992; Dawers et al. 1993; Scholz et al. 1993), while others provided n values of 1.5 (Marrett & Allmendinger 1991; Gillespie et al. 1992) or of n = 2 (Walsh & Watterson 1988). Following the recognition of a relationship between D and W, several models attempted to quantitatively describe the fault growth process. Most early models assumed that faults progressively increase in both length and displacement during their activity. A more recent model

6001

'

I

'

I

i

I

'

FAULTS

33

(Walsh et al. 2002) predicts that faults at an early stage possess nearly constant lengths and grow mainly by increasing their cumulative displacement. This model requires that a wide range of fault lengths be achieved at an early stage of deformation and has been tested for a reactivated normal fault system in the Timor sea (NW Australia). The maximum length (Wmax) and throw (D~x) of the faults studied here are shown in Figure 8. The database includes all faults of the mapped area, for which tips could be isolated and accurately mapped, that is, a total of 15 faults ranging in length from 500 m to about 8000 m and with throws from 20 m up to about 550 m (Table 1). Among the different correlation functions (i.e. linear and power type), the function giving the lowest root mean square error (R 2) is of linear type: D = 0.0597-W 1. This result is in agreement with the review by Clark & Cox (1996) who have re-examined 11 data sets from previous analyses and found that most of the data sets are consistent with a linear relationship (i.e. n = 1), whereas exponents of 1.5 and 2.0 are inconsistent with the data.

Fault scaling

D=c.

OF NORMAL

Conclusions Our study of the M. LeScalette and Cesi-S. Martino fault segments provides evidence of fault interaction both in map view and in the distribution of throw along the segments. In map

I

'

I

'

I

'

I

'

I

'

O Nll .-. 500

. ~ 400

3OO 200

N14

~2 N9 ~

O~N3

oN10 y=0 0597x y~0=;0.~

N44

100

0

~lY 0

5

I , I , I , I , I , I , I , I , 1000 2000 3000 4000 5000 6000 7000 8000 9000

fault length (m) Fig. 8. Plot of the maximum length (Wma• against the maximum throw (Dmax)of the normal faults (15 faults).

34

F. MIRABELLA ET AL.

Table 1. Maximum fault length (Wmax)and the corresponding maximum throw values (Dmax) of the normal faults for which length -throw values could be well established in the mapped area. Position of faults are shown in Figure 2a Maximum length and throw of the mapped normal faults Fault

Length, W (m)

Throw, D (m)

N1 N2 N3 N4 N5 N6 N7 N8 N9 NIO Nil N12 N13 N14 N15

1700 4430 6790 1800 800 2750 3250 2450 1600 3330 8210 1170 3500 2370 500

70 270 360 90 50 170 220 180 125 18 550 150 220 190 20

view, the two fault segments display an en&helon arrangement, the northernmost segment (M. LeScalette) being shifted about 3 . 4 k m eastward of the southern Cesi-S. Martino segment, for an overlap of about 1 km. Within the separation zone, other minor SW-dipping and subordinately NE-dipping normal faults are present, among which the fault named NIO is the major connecting structure between the two segments. The shape of the throw distribution along the two segments is asymmetric, maximum values being close to the interaction zone. The cumulative throw distribution along the studied segments, from NW to SE, shows that throw values persist within the interaction zone. The possibility of mechanical interaction between fault segments is relevant to the estimation of the maximum possible earthquake and for hazard assessment. According to the relationship of Wells and Coppersmith (1994) between rupture length and moment magnitude, a magnitude of about M c. 6.2 could be reached. Seismological data (Pino et al. 1999; Pino & Mazza 2000) indicate that on 26 September 1997, mechanical interaction between two adjacent segments occurred in the Colfiorito fault system, confirming that the geometry of the system is such that mechanical interaction between the segments is likely to occur. Maxinmm lengths and throw data collected for 15 faults give a best scaling relationship of the linear type: Ornax/Wma x = 0.0597.

We thank C. Faccenna and an anonymous reviewer for their detailed comments and suggestions, which greatly improved the manuscript. We thank the Free Software Foundation (http://www.fsf.org) for providing all software employed for data analysis, figure editing, and text typesetting. This work was funded by grants from the Italian Ministry of Research (MIUR) and Regione Umbria to M. R. Barchi.

References AKI, K. 1979. Characterization of barriers on an earthquake fault. Journal of Geophysical Research, 84, 6140-6148. AMATO, A., AZZARA, R. et al. 1998. The 1997 Umbria-Marche, Italy earthquake sequence: a first look at the main shocks and aftershocks. Geophysical Research Letters, 25, 2861-2864. AN, L. 1997. Maximum link distance between strikeslip faults: observations and constrains. Pure and Applied Geophysics, 150, 19-36. ANDERS, M. & SCHLISCHE, R. 1994. Overlapping faults, intrabasin highs and the growth of normal faults. Journal of Geology, 102, 165-180. BARBA, S. & BASILI, R. 2000. Analysis of seismological and geological observations for moderate-size earthquakes" the Colfiorito fault system (Central Apennines, Italy). Geophysical Journal International, 141, 241-252. BARCHI, M. 2002. Lithological and structural controls on the seismogenesis of the Umbria region: observations from seismic reflection profiles. Bollettino della Societd Geologica Italiana, l(Vol. Spec.), 1,855-864. BARCHI, M., GALADINI, F., et al. 2000. Sintesi delle conoscenze sulle faglie attire in Italia Centrale: parametrizzazione ai fini della caratterizzazione della pericolositd sismica, GNDT, Gruppo Nazionale per la Difesa dai Terremoti, Korea, Spec. Publ. BARCHI,M., PuccI, S. etal. 2001. A geological map of the Colfiorito area. Convegno FIST GeoItalia 01, Volume dei Riassunti, 333-334. BARKA, A. & KADINSKu K. 1988. Strike-slip fault geometry in Turkey and its influence on earthquake activity. Tectonics, 7, 663-684. BASILI, R. & MEGHRAOUI,M. (2001). Coseismic and postseismic displacements related with the 1997 earthquake sequence in Umbria-Marche (central Italy). Geophysical Research Letters, 28, 26952698. BONCIO, P. & LAVECCHIA, G. 2000. A structural model for active extension in Central Italy. Journal of Geodynamics, 29, 233-244. BOSCHI, E., GUIDOBONI, E., FERRARI, G., VALENSlSE, G. & GASPERIN1, P. 1997. CFTI, Catalogo dei Forti terremoti Italiani dal 461 a.c. al 1990. ING-SGA, Bologna. BURBANK, D. & ANDERSON, R. 2001. Tectonic Geomorphology. Blackwell Science, Oxford, UK. BURGMANN, R., POLLARD, D. & MARTEL, S. 1994. Slip distributions on faults: effects of stress gradients, inelastic deformation, hetereogeous

SEGMENTATION AND INTERACTION OF NORMAL FAULTS host-rock stiffness, and fault interaction. Journal of Structural Geology, 16, 1675-1690. CALAMITA, F., PIZZ1, A., RIDOLFI, A., RUSCIADELLI, M. & SCISCIAN1, V. 1998. I1 buttressing delle faglie sinsedimentarie pre-thristing sulla strutturazione neogenica della catena appenninica: l'esempio della m.gna dei fiori (appennino central esterno). Bollettino della Societd Geologica Italiana, 117, 725-745. CALAMITA, F., COLTORTI, M., PIERANTONI, P., PIZZI, A., SC1SC1ANI, V. & TURCO, E. 1999. Relazioni tra le faglie quaternarie e la sismicita' nella dorsale appenninica Umbro-Marchigiana: L'area di Colfiorito. Studi Geologici Camerti, 14, 177-191. CARTWRIGHT, J., TRUDGILL, B. & MANSFIELD, C. 1995. Fault growth by segment segment linkage: an explanation for scatter in maximum displacement and trace length data from the canyonlands grabens of SE Utah. Journal of Structural Geology, 17, 1319-i326. CELLO, G., DEIANA, G. et al. 1998. Evidence for surface faulting during the September 26, 1997 Colfiorito (Central Italy) earthquakes. Journal of Earthquake Engineering, 2, 1-22. CHIARABBA, C. 8r AMATO, A. 2003. Vp and vp/vs images in the mw 6.0 Colfiorito fault region (central Italy): A contribution to the understanding of seismotectonic and seismogenic processes. Journal of Geophysical Research, 108, 2248. CHIARALUCE, L., ELLSWORTH, W., CHIARABBA,C. & Cocco, M. 2003. Imaging the complexity of an active complex normal fault system: the 1997 Colfiorito (central Italy) case study. Journal of Geophysical Research, 108(B6), 2294, 10.1029/ 2002JB002166. CLARK, R. & Cox, S. 1996. A modem regression approach to determining fault displacementlength scaling relationships. Journal of Structural Geology, 18, 147-152. COLTORTI, M. & PIERUCCINI, P. 2000. The planation surface across the Italian peninsula: a key tool in neotectonic studies. Journal of Geodynamics, 29, 323-328. COLTORTI, M., ALBIANELLI, A., BERTINI, A., FICCARELLI, G., LAURENZI, M., NAPOLEONE, G. & TORRE, D. 1998. The Colle Curti mammal site in the Colfiorito area (Umbro-Marchean Apennine, Italy): Geomorphology, stratigraphy, paleomagnetism and palynology. Quaternary International, 47(8), 107-116. COWIE, P. & SCHOLZ, C. 1992. Displacement-length scaling relationship for faults: data synthesis and discussion. Journal of Structural Geology, 14, 1149-1156. CPTI 1999. Catalogo Parametrico dei Terremoti Italiani. ING, GNDT, SGA, SSN, Bologna. CRESTA, S., MONECHI, S. & PARISI, G. 1989. Stratigrafia del mesozoico e cenozoico hell'area Umbro-Marchigiana. Itinerari geologici sull'Appennino Umbro-Marchigiano (Italia). Memorie Descrittive della Carta Geologica d'Italia, 39, 1-182.

35

DAWERS, N., ANDERS, M. & SCHOLZ,C. 1993. Growth of normal faults: displacement-length scaling. Geology, 21, 1107-1110. DESCHAMPS, A., INNACCONE,G. & SCARPA, R. 1984. The Umbrian earthquake (Italy) of 19 September 1979. Annales Geophysicae, 2(1), 29-36. DESCHAMPS, A., COURBOULEX,F. et al. 2000. Spatiotemporal evolution of seismic activity during the Umbria-Marche crisis, 1997. Journal of Seismology, 4(4), 377-386. DOGLIONI, C., D'AGOSTINO, N. & MARIOTTI, G. (1998). Normal faulting vs regional subsidence and sedimentation rate. Marine and Petroleum Geology, 15, 737-750. DZIEWONSKI, A., FRANZEN, J. ~ WOODHOUSE, J. 1985. Centroid-moment tensor solutions for April-June, 1984. Physics of Earth and Planetary Interiors, 37, 87-96. EKSTROEM, G., MORELLI, A., BOSCHI, E. & DZIEWNONSKI, A. 1998. Moment tensor analysis of the central Italy earthquake sequence of September-October 1997. Earth and Planetary Sciences Letters, 25(11 ), 1971 - 1974. ELLIOTT, D. 1976. The energy balance and deformation mechanisms of thrust sheets. Philosophical Transactions of the Royal Geological Society of London, A283, 289-312. FICCARELLI, G. & MAZZA, P. 1990. New fossil foundings from the Colfiorito basin (Umbro-Marchean Apennine). Bollettino della Societd Paleontologica Italiana, 29, 245-247. GILLESPIE, P., WALSH, J. ~; WATTERSON, J. 1992. Limitations of dimensions and displacement data from single faults and the consequences for data analysis and interpretation. Journal of Structural Geology, 14, 1157-1172. GUPTA, A. & SCHOLZ, C. 2000. A model for normal fault interaction based on observations and theory. Journal of Structural Geology, 22, 865-879. HAESSLER, H., GAULON, R. et al. 1988. The Perugia (Italy) earthquake of 29 April 1984: a microearthquake survey. Bulletin of the Seismological Society of America, 78(6), 1948-1964. KING, G. & NABELEK, J. 1985. Role of fault bends in the initiation and termination of earthquake rupture. Science, 228, 984-987. LAVECCHIA,G., B ROZZETTI,F., BARCHI,M., KELLER,J. & MENICHETTI, M. 1994. Seismotectonic zoning in east-central Italy deduced from the analysis of the Neogene to present deformations and related stress felds. Geological Society of America Bulletin, 106, 1107-1120. MARIUCCI, M., AMATO, A. & MONTONE, P. 1999. Recent tectonic evolution and present stress in the Northern Apennines (Italy). Tectonics, 18(1), 108-118. MARRETT, R. & ALLMENDINGER, R. 1991. Estimates of strain due to brittle faulting: sampling of fault populations. Journal of Structural Geology, 13, 735-737. MENARD, H. 1962. Correlation between length and offset on very large wrench faults. Journal of Geophysical Research, 67, 4096-4098.

36

F. MIRABELLA ET AL.

MESSINA, P., GALADINI,F., GALLI, P. & SPOSATO, A. 1999. Evoluzione a lungo termine e caratteristiche della tettonica attiva nell' area Umbro-Marchigiana colpita dalla sequenza sismica del 1997-98 (Italia Centrale). In: Peruzza, L. (ed) Progetm MISHA - Metodi Innovativi per la Stima dell'Hazard: Applicazione all'Italia Centrale. CNR-GNDT, Roma, 32-42. MIRABELLA, F. & PuccI, S. 2002. Integration of geological and geophysical data along a section crossing the region of the 1997-98 Umbria-Marche earthquake (Italy). Bollettino della Societ6 Geologica ltaliana, l(Vol. Spec.), 891-900. MIRABELLA, F., CIACC10, M. G., BARCHI, M. R. & MERLINI, S. 2004. The Gubbio normal fault (Central Italy): Geometry, displacement distribution, and tectonic evolution. Journal of Structural Geology, 2233-2249. MURAOKA, H. & KAMATA, H. 1983. Displacement distribution along minor fault traces. Journal of Structural Geology, 5, 483-495. OPHEIM, J. & GUDMUNDSSON,A. 1989. Formation and geometry of fractures, and related volcanisms, on the krafla fissure swarm, northeast island. Bulletin of the Geological Society of America, 101, 16081622. PEACOCK, D. 1991. Displacement and segment linkage in strike slip fault zones. Journal of Structural Geology, 13, 1025-1035. PEACOCK, D. & SANDERSON, D. 1991. Displacement and segment linkage and relay ramps in normal fault zones. Journal of Structural Geology, 13, 721-733. PINt, N. & MAZZA, S. 2000. The Umbria-Marche (central Italy) earthquakes: relation between rupture directivity and sequence evolution for the m w > 5 shocks. Journal of Seismology, 4, 451-461. PINt, N., MAZZA, S. & BOSCHI, E. 1999. Rupture directivity of the major shocks in the 1997 Umbria-Marche (central Italy) sequence from regional broadband waveforms. Geophysical Research Letters, 26(14), 2101-2104. PIZZI, A. & SCISIANI, V. 2000. Methods for determining Pleistocene-Holocene displacement in active faults reactivating pre-Quaternary structures. Example from the central Apennines (Italy). Journal of Geodynamics, 29, 445-457. PIZZI, A., CALAMITA, F., COLTORTI, M. & I~ERUCC~NI, P. 2002. Quaternary normal faults, intramontane basins and seismicity in the Umbria-Marche-Abrnzzi Apennine ridge (Italy):

contribution of neotectonic analysis to seismic hazard assessment. Bollettino della Societd Geologica ltaliana, l(Vol. Spec.), 923-929. RANALLI, G. 1977. Correlation between length and offset on strike-slip faults. Tectonophysics, 37, 1- 7. ROBERTS, G. (1996). Variation in fault- slip direction s along active and segmented normal fault systems. Journal of Structural Geology, 18, 835- 845. SCHOLZ, C. & GUPTA, A. 2000. Fault interactions and seismic hazard. Journal of Geodynamics, 29, 459 -467. SCHOLZ, C., DAWERS, N., Yu, J. & ANDERS, M. 1993. Fault growth and fault scaling laws: preliminary results. Journal of Geophysical Research, 98, 21 951-21 961. SEGALL, P. & POLLARD, D. 1980. Mechanics of discontinuos faults. Journal of Geophysical Research, 85, 555-568. SI~SON, R. 1986. Rupture interaction with fault jogs. In: DAS, S., BOATWmGHT, J. & SCHOLZ, C. (eds) Earthquake Source Mechanics. American Geophysical Union Geophysical Monograph 6, Maurice Ewing series, 157-167. SIEH, K., JONES, L. et al. 1993. Near-field investigations of the Landers earthquake sequence, April to July 1992. Science, 260, 171-176. STRAMONDO, S., TESAURO, M. et al. 1999. The September 26, 1997 Colfiorito, Italy, earthquakes: modeled coseismic surface displacement from SAR interferometry and GPS. Geophysical Research Letters, 26(7), 883-886. WALSH, J. & WATTERSON, J. 1988. Analysis of the relationship between displacements and dimensions of faults. Journal of Structural Geology, 10, 239-247. WALSH, J., NICOL, A. & CHmDS, C. 2002. An alternative model for the growth of faults. Journal of Structural Geology, 24, 1669-1675. WATTERSON, J. 1986. Fault dimensions, displacements and growth. Pure and Applied Geophysics, 124, 365-373. WELLS, D. & COPPERSMITH, K. 1994. New empirical relationships among magnitude, rupture length, rupture width, rupture area and surface displacement. Bulletin of the Seismological Society of America, 84, 974-1002. WILLEMSE, E., POLLARD,D. & AYDIN, A. 1996. Three dimensional analyses of slip distributions on normal fault arrays with consequences for fault scaling. Journal of Structural Geology, 18, 295- 309.

Textural controls on the brittle deformation of dolomite: variations in peak strength N. J. A U S T I N 1'2 & L. A. K E N N E D Y 1

1Earth and Ocean Sciences, The University of British Columbia, 6339 Stores Road, Vancouver, BC, V6T 1Z4, Canada 2present address: Earth, Atmospheric, and Planetary Sciences, The Massachusetts Institute of Technology, 77 Massachusetts Avenue, 54-710, Cambridge, MA, 02139-4307, USA (e-mail: naustin @mit. edu)

Abstract: Textural controls on the peak strength of dolomite are investigated through 23 triaxial deformation experiments, performed at confining pressures of 25, 50, and 100 MPa on texturally diverse dolomites. The mechanical data from these experiments are fit to an empirical failure criterion to elucidate the most significant parameters in dictating the peak strength of dolomite. Neither grain size nor porosity is required to quantitatively predict the peak strength of dolomite. Instead, only the effective Young' s modulus, and the empirically predicted uniaxial compressive strength, along with experimentally controlled confining pressure, are required in order to explain the peak strength to and R 2 better than 0.89. A Hall-Petch relationship does not apply to this data set as a consequence of the variability in intragranular and grain boundary textures, which appear to overshadow the role of grain size. It is, therefore, essential that grain boundary textures and intragranular flaws be examined prior to making predictions regarding the relative peak strengths of chemically and mineralogically similar dolomites.

The textural properties of dolomite can vary substantially within mineralogically similar formations, and it is these variations in porosity, grain size, grain boundary texture, and intragrain texture that result in very different peak strengths. Variations in peak strength in the brittle field can play an important role in the development and evolution of shear zones, and thus it is important to understand the textural parameters that influence this property. For the purpose of this paper, we use the term texture to refer to the geometrical aspects of, and mutual relations among, the rock' s component particles or crystals (Bates & Jackson 1984). In general, the textural properties of rocks dictate the location and concentration of the initial flaws, the importance of which have been well documented since Griffith's classic papers (Griffith 1921, 1924). Griffith's theory satisfactorily explains the initiation of cracks; however, it cannot account for crack propagation and coalescence (Hoek & Bieniawski 1966), which dictate the peak strength of rocks (Horii & Nemat-Nasser 1985; Ashby & Sammis 1990; Sagong & Bobet 2002). One model commonly used to explain crack propagation and coalescence is the wing crack model, which proposes the nucleation and growth of a wing crack when the shear stress along an inclined crack exceeds the

frictional resistance, providing an explanation for crack nucleation, propagation, and coalescence (Brace & Bombolakis 1963; Brace et al. 1966; Horii & Nemat-Nasser 1985; Ashby & Sammis 1990). Using experimental rock deformation data from triaxial experiments performed on texturally diverse, nearly stoichiometric dolomite, current empirical relationships are examined with the aim of advancing our understanding of the role of the textural properties of dolomite on its peak strength. The theories of compressional fracture mechanics (for a review, see Horii & NematNasser 1985; Ashby & Sammis 1990; Fredrich et al. 1990; Evans et al. 1990) are compared with experimental observations to show that very few rock properties are required to predict empirically the peak strength of dolomite. Grain size does not correlate well with the peak strength, a discrepancy that can be explained by variations in the grain boundary and intragranular textures.

Experiments Methodology To examine the role of rock texture on the peak and yield strengths of dolomite (see Table 1 for

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. DeformationMechanisms, Rheology and

Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 37--49. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

38

N.J. AUSTIN & L. A. KENNEDY

Table 1. List of symbols Symbol o-~ 03 0"p 0"y Pc rate ~b dm dx C Co /x qJ Eeff E R R2 K~c KI a Na l y

Description Principal compressive stress Least compressive stress Peak differential stress Differential stress at yielding Confining pressure Strain rate Porosity Mean grain size Maximum grain size Cohesion Unconfinedcompressive strength Coefficient of friction Coulomb friction parameter Effective Young's modulus True Young's modulus Correlation coefficient Square of the correlation coefficient Critical stress intensity factor Stress intensity factor Initial half flaw length Flaw density Length of the wing cracks The angle between the initial flaw and 0"~

a list of symbols used in this paper), a series of 23 deformation experiments (Table 2) were performed on texturally diverse dolomite from the Imasco Minerals quarry in the Badshot Formation of southeastern British Columbia, the Mighty White Dolomite quarry in Rock Creek, British Columbia, and the Graymont Dolime quarry in the Niagara Formation of Ohio. Badshot dolomite has the largest mean grain size of the sample suites examined (Table 2), with grains exhibiting extensive intragranular cleavage, lobate grain boundaries, and subgrains (Fig. l a). Rock Creek dolomite has the finest grain size of the sample suites studied (Table 2). Grains are subhedral and equigranular with very rare cleavage or twinning (Fig. lb). All the Badshot and Rock Creek samples have porosities of less than 1.5%. Niagara dolomite is intermediate in grain size between the Badshot and Rock Creek samples (Table 2). Grains exhibit anhedral to subhedral morphologies. Intragranular cleavage is present and visible, although less so than in the Badshot samples, but twinning is rare (Fig. lc). Unlike the Badshot and Rock Creek samples, Niagara samples contain vuggy porosity on the size scale of grains, resulting in porosities between 6.6% and 13.5%. All of these sample suites consist of greater than 85% nearly stoichiometric dolomite (Austin et al. 2005). The Large Sample Rig (LSR) triaxial rock press at the University of British Columbia was

used for these experiments (Austin et al. 2005). Experiments were performed on oven-dried right cylinders (baked at 358 K for 24 hours) with diameters of 22.2 mm and length to diameter ratios of c. 2: 1, at confining pressures of 25, 50, and 100MPa, using an argon gas confining medium, at 296 K, at a strain rate of c. 1 x 10 -5 s -1 (Table 2).

Experimental results

Differential stress versus axial strain curves are presented in Figure 2, along with the associated mechanical data in Table 2a, and micrographs of the deformed samples are presented in Figure l(d, e, f). The micrographs indicate that all deformation was accommodated by brittle mechanisms, consistent with what has previously been observed in dolomite under these conditions (Turner et al. 1954). Deformed Badshot samples exhibit extensive intragranular fracturing at all confining pressures, along with transgranular faults (Fig. ld). Rock Creek samples contain sample scale faults, and transgranular deformation: individual grains remain predominantly undeformed except for within 1 - 2 grains of fault zones (Fig. le). Niagara samples exhibit variable amounts of intragranular deformation (Fig. lf). In the vicinity of pores, grains contain extensive intragranular deformation, with microcracks making use of intragranular cleavage planes at all confining pressures; however, away from pores intragranular deformation is suppressed. Transgranular microcracks are also common, the morphology of which is generally controlled by porosity. In all sample suites, the deformed microstructures are independent of confining pressure. Values for Eeff, Oy, Co, C, and /x were obtained for each sample from analysis of the mechanical data (Table 2a). Effective Young's modulus and yield stress values were obtained by isolating the linear portion of the differential stress-axial strain curve, obtaining the slope of the best fit line to this region, and picking the stress at which the differential stress-axial strain curve deviated from this linear trend, respectively. Co, C, and /z were calculated for each sample suite based on Coulomb's failure criterion in principal stress space. O'1 = Co -q- tan ~9 ~3

(1)

where Co and qJ are converted to C and/x using the relationships sin/x = (tan q J - 1)/(tan t ) + 1)

(2)

THE PEAK STRENGTH OF DOLOMITE

39

t"q

+1 "~ ,.d

.,.... t".-

t"q I

<

,,....,,

9

+1

I~ ~

r

~

,...~ ,~. ~

:~

~

~

~

..~.. ,..~ ~.. ~1 .,~i..~. ~

~

r

'

~

'

t.~ ..~.. r162 ,...~ ~r~ r

~

cr ,~, ,,~. ~

~

O

~

~5 ~ C5 ~5 C5c5 C5~5 C5~ C5C5~ ~5 ~5 ~ ~ ~5 ~ ~5 ~5 ~5 ~ 1--,,

~

t",l r

~q r

r

r

~"-1 ~"q r

r

r

,...,,

...,.,, O

z O r~'-~

O

=]. •

. ~

,.,~..,~. ~ _r ' ' r~-t"~ ~-'-,~ . -. ~,~~..,~,1.~~t~ r1 I"" ~

t"q c'~ . . t'q ~ . ,-.e"" . t,-.-~ ~ ~t-~ ~r~ e~ ~ e ~

r

,.--~ t"-I

~e~-~~

~'~ ~

~

~

~'~ ~'~ ~7~

t"q t'-,i

40

N. J. AUSTIN & L. A. KENNEDY

(a) (d)

(b) (e)

(c) (f)

Fig. 1. Optical micrographs of undeformed samples from: (a) Badshot dolomite, highlighting lobate grain boundaries (L) and intragranular cleavage, present throughout the samples; (b) Rock Creek block RC4 highlighting straight grain boundaries (S); and (e) Niagara dolomite, highlighting the prevalence of intragranular flaws, giving the grains a 'dirty' appearance, along with straight grain boundaries (S). Images of the deformed samples: (d) is an optical micrograph of Badshot dolomite (CBB 1-9), highlighting deformation along intragranular flaws predominantly related to cleavage planes, along with the presence of transgranular cracks and sample scale faults; (e) is a secondary electron SEM image of deformed Rock Creek dolomite (RC5-1) highlighting microcracks, dilation along grain boundaries, and comminution of grains in the vicinity of faults; (f) is a secondary electron SEM image of Niagara dolomite (OH5-1), highlighting transgranutar microcracks that are commonly present along grain boundaries and intragranular flaws, and the strong link of both types of deformation to the presence of pores.

THE PEAK STRENGTH OF DOLOMITE (a) 600-

41

(b) 600-

CBB1-4 (103MPa) "-" 400'

~ 400"

B1-9 (104 MPa) CBBI-1 (51 MPa)

r~

RC4-2 (103 MPa)

o.

CBB1-7 (51 MPa) ~ 200' -9

200' ,CBB1-2 (24 MPa) CBBI-6 (24 MPa)

0 0

2

Strain (%)

4

6

(e) 600"

C4-7 (52MPa) ~ RC4-5(51 MPa) 1 ~RC4-6 (25 MPa) "l 1~ ~,~ "~ RC4-3(24 MPa) RC4-4 (23 MPa) i 0

"

, 2

Strain (%)

i 4

(d) 600-

RRC5-1(103MPa)

OHI-I (52 MPa)

400'

400"

"4

It ~

200"

~

RC5-7(51MPa)

i

0 i 0

H

5

-

200'

1 (103 MPa) OH5-3 (52 MPa)

0 2

Strain (%)

4

6

0

2

Strain (%)

4

6

Fig. 2. Differential stress-axial strain curves for: (a) Badshot dolomite, (b) Rock Creek (RC4), (c) Rock Creek (RC5), and (d) Niagara dolomite. The confining pressure of each experiment is indicated in brackets.

and C -- Co(1 - sin/x)/2cos/.~

(3)

(Sofianos & Halakatevakis 2002). In principal stress space, Co is the unconfined compressive strength, while ~b is a geometrical term, that relates to the coefficient of internal friction by equation (2) (Goodman 1980; Engelder 1993).

Controls on the peak strength of dolomite The relationship between sample properties, controlled experimental conditions, and the measured peak strength are examined using correlation coefficients (R) (Table 3, Fig. 3).

The low values observed in all correlation coefficients highlight the complexity of textural controls on variations in peak strength. There is a moderate negative correlation ( - 0 . 3 7 9 0 ) between porosity and peak strength. The parameters with the strongest correlations with peak strength are effective Young's modulus (0.4961), cohesion (0.5477), unconfined compressive strength (0.6060), and confining pressure (0.7890). Porosity is more strongly correlated to the yield stress (R = -0.5022), as is the effective Young's modulus (R = 0.5872), while cohesion, unconfined compressive strength, and confining pressure have a decreased correlation (R = 0.4708, 0.5521, and 0.6544, respectively) with yield stress. Significantly,

42

N.J. AUSTIN & L. A. KENNEDY

Table 3. Correlation coefficients between sample properties, experimental conditions, and experimental results for (a) our experiments; (b) this study, including the data from Hatzor et al. (1997) 4'

drn

dx

Eeff

1

-0.113 1

-0.057 0.989 1

-0.655 -0.234 -0.225 1

1

-0.242 1

-0.2 0.976 1

-0.745 0.047 0.01 1

Iz

Pc

e rate

O-p

O-y

-0.18 -0.24 0.648 0.494 0.649 0.495 0.221 0.344 1 0.98 1

-0.359 -0.829 -0.847 0.372 -0.694 -0.551 1

-0.116 0.08 0.032 0.058 0.178 0.182 -0.095 1

0.707 0.111 0.092 -0.794 -0.143 -0.231 -0.35 0.117 1

-0.379 0.114 0.089 0.496 0.548 0.606 -0.13 0.789 -0.273 1

-0.502 0.002 -0.029 0.587 0.471 0.552 0.02 0.654 -0.439 0.941 1

-0.411 0.668 0.625 0.372 1

0 . 1 7 1 -0.373 -0.521 0.285 -0.445 0.215 - 0.086 0.32 -0.705 0.448 -0.542 0.417 1 -0.431 1

-0.341 0.359 0.273 0.45 0.443 0.378 -0.545 0.611 1

-0.574 0.295 0.244 0.569 0.687 0.715 -0.438 0.859 0.545 1

C

Co

(a) 4' dm d•

Eeff C Co /~ Pc rate O-p O-y

(b) 4, dm dx Eeff C Co /z Pc e rate ~rp

-0.454 0.549 0.516 0.468 0.971 1

there exists only a very weak correlation between either mean or maximum grain size and peak strength or yield stress, over the large range of grain sizes being studied (Table 3). Our experimental data are also compared with the data of Hatzor et al. (1997) (Table 2b) performed on dolomite from the Amindav Formation in central Israel. Experiments from their data set were excluded where: (1) the experimental data was incomplete with respect to

I.

0.8-

o This Study (1997) This study, including Hatzor et al. []

0.6.

o o

0

Eeff

C

~ 0.4.

~

0.2,

~

0.

[]

o

"~ -0.2. ~ -0.4

o

-0.6

o

n

o

o

dx

dm

Co

Pc

o

measurements of porosity, grain size, effective Young's modulus, or peak stress, or (2) the modal mineralogy consisted of less than 80% dolomite. Co, C, and ~ were calculated for the data of Hatzor et al. (1997) using the same method as for the experiments performed as part of this study. Combining our experimental data with that of Hatzor et al. (1997) results in an adjustment of the correlation coefficients, although their relative values remain unchanged (Fig. 3). The correlation between peak strength and confining pressure (0.8586) is intensified, as is that with cohesion (0.6871), unconfined compressive strength (0.7154), and effective Young's modulus (0.5693). Notably, the correlation with porosity becomes more strongly negative (-0.5743), yet the correlation between peak strength and grain size remains poor (dm: 0.2947; dx: 0.2439).

A modified empirical f a i l u r e criterion

-0.8 Property

Fig. 3. The trends in correlation coefficients with peak strength for our data set and our data set combined with the data of Hatzor et al. (1997).

In order to elucidate the parameters most important in explaining observed variations in peak strength, the empirical failure criterion proposed by Hatzor & Palchik (1998) is examined: O'p = a ( E eff O'3) b / d~Cd 1/ 2

(4)

THE PEAK STRENGTH OF DOLOMITE The premise behind this relationship is that a H a l l - P e t c h relationship

however, when the Badshot and Niagara samples are included, the proportionality in equation (5) is no longer valid (Fig. 4). This discrepancy is highlighted by fitting the empirical relationship proposed by Hatzor & Palchik (1998) (equation 4) to the data obtained from our experiments and those from Hatzor et al. (1997) (Fig. 5a), using a N e l d e r - M e a d routine to minimize the absolute value of the misfit error, and comparing this relationship to one

(5)

O'p oc dm 1/2

43

holds true across a wide range of grain sizes (Olsson 1974; H u g m a n & Friedman 1979; Fredrich et al. 1990). Within the Rock Creek sample suite, it appears that a H a l l - P e t c h relationship is valid (Fig. 4);

600-

(a)

0 2 5 MPa o 50 MPa t, 1O0 MPa

500A

400.

0

[]

0

300-

A

0 [] 0

.,..~ ra

200.

0 0

100.

0

100

0

200

300

400

500

600

d m (/.tm) 600-

(b)

o 25 MPa n 5 0 MPa A 100 MPa

.-'- 500.

400-

=

300-

~ []

200-

o o

100-

0

0.05

0.1

0.15 1/2 d-l/2 m (~tm)

0.2

0.25

Fig. 4. The relationship between (a) mean grain size and (b) the inverse square root of mean grain size and peak differential stress for our experiments. Errors in stress measurements are + 2.7 MPa, and thus error bars are smaller than the data points. The standard deviations in mean grain size are reported in Table 1.

44

N.J. AUSTIN & L. A. KENNEDY

a)

b)

7OO

700 Eq. 6 .

Eq. 4 60O R 2 =0.4168

600 Rz =0.7937

5O0

5oo

400

Unconfined

30O

. .

~

100

o

O 0 700 r

1O0

200

300

400

~

500

6;0

700

o

J

"

lOO

2(Io

300

400

500

600

700

Eq. 9 with unit'balance .

'-~

400

400

"~

300

300

200

200

10o

1oo 1O0

.

.

.

J

600 500

o

.

700

Eq. 8 R ~=0.9200

500

o

.

a)

e)

600

.

Badshot Dolomite

20O

. ,...,

.

oO2oo, oo/ ; il , 400

~ "" .

.

200

300

400

500

600

700

0

0

I oo

200

300

400

500

600

700

M e a s u r e d P e a k D i f f e r e n t i a l Stress ( M P a ) Fig. 5. (a) The ability of equation (4) to explain the peak strength of dolomite with a = 2.4587, b = 0.4845, and c -- -0.1372, resulting in an R z of 0.4168. Note that this relationship cannot explain the unconfined samples or the Badshot samples. (b) Equation (6) fit to the data set with a = 1.3687, b -- 0.3624, and c -- 0.0301, producing an R 2 of 0.7937. Equation (6) can explain the coarse-grained Badshot samples (highlighted); however, unconfined samples are still not explained. (e) Equation (8) fit to the data set with a = 0.0003 and b = 0.8796, resulting in an R 2 of 0.9200. (d) Equation (9), optimized for a unit balance, with a = 0.5617, b --- 0.2227, and c = 0.7773, resulting in an R2 of0.8918.

with the same form and the same number of fit parameters, without the grain size term:

O'p = a(EeffOr3 )b / q~c

(6)

(Fig. 5b). Notably, it is the Badshot samples, with a coarse grain size, lobate grain boundaries, and extensive intragranular cleavage that are not fit by equation (4), but are fit by equation (6). This indicates that while a constant proportionality between peak strength and the inverse square root of grain size may exist in samples with limited textural variation, when the degree of variation is increased, the constant proportionality may no longer be assured. Further, in equation (6), the value of c (0.0301) approaches zero, indicating that porosity is not required to fit this data set, and thus equation (4) may be further modified, resulting in: O'p = a(EeffO'3) b

(7)

Equation (7) dictates that, as confining pressure goes to zero, the predicted strength of the rock also goes to zero. In Figure 5(a and b), based on equations (4) and (6), the strength of samples deformed at atmospheric pressure is significantly underestimated. Based on C o u l o m b ' s failure criteflon (equation 1), under zero confining pressure, the strength of the rock is defined as the uniaxial compressive strength (Co). According to C o u l o m b ' s failure criterion, equation (7) is modified with a Co term, resulting in o.p

=

a(Ee~o3) b + Co

(8)

which results in an R 2 of 0.92 (Fig. 5c). There is a unit imbalance in this relationship that may be addressed in one o f two ways. First, it may be assumed that the information contained within the fit parameter a has units of pressure, and that this information is correlated to the fit parameter b. Alternatively, Eeff and o'3 may be fit with different parameters, which are

THE PEAK STRENGTH OF DOLOMITE

45

l

0.9 0.8 0.7 _ Unit balance

"

~

_

~

~

0.6

b 0.5

._-o.3

0.4 0.3 0.2 0.1

_

a=N

0. l

0.2

0.3

0.4

0.5 C

0.6

0.7

0.8

0.9

Fig. 6. The range of b and c values that result in an R 2 value of >0.89, using equation (9), for varying a values, highlighting the range of a, b, and c values that can be used without reducing the ability of equation (9) to explain the data set. correlated so that they sum to 1 (Figs 5d and 6). This results in an equation with the form: b

o-p = aEeffO'~3 -}- Co

(9)

There is a range of fit parameters (a, b, c) that produce a unit balance without diminishing the ability of equation (9) to describe the experimental data (Fig. 6). When c = 1 and b = 0, equation (9) is linear with respect to confining pressure; however, b may be up to 0.65 without reducing the R 2 below 0.89 (which contains fewer significant figures than the optimized R 2 of 0.8918, thus accounting for sample variability). Confining pressure and effective Young's modulus are poorly correlated (0.058 and 0.320 for this data set and this data set combined with that of Hatzor et al. 1997), and thus the peak strength of dolomite must be related to the effective Young's modulus. The relationship between effective Young's modulus and the concentration and length of the initial flaws, as well as with the porosity, was theoretically analysed by Walsh (1965a, b, c). It was shown that the effective Young's modulus should be inversely proportional to increases in flaw length, density, or porosity (as observed in Table 3), and thus based on the later work of Horii & Nemat-Nasser (1985), Sammis & Ashby (1986), Ashby & Sarnmis (1990), and Zhang et al. (1990), an increase in effective Young's modulus should correspond to an increase in strength for materials with the same true Young's modulus. This is in agreement

with the empirical observations quantified in equation (9). From a practical standpoint, the effective Young's modulus of a material is a powerful parameter for use in an empirical relationship as it is readily obtained from the differential stressaxial strain curve of compression experiments. Furthermore, no assumptions need to be made regarding the location or nature of the initial flaws, or regarding the mechanism of deformation.

The relationship between flaws and grain size In order to understand why the Hall-Petch relationship breaks down for texturally variable dolomites, we examined measured peak strength variations using the wing crack model (Horri & Nemat-Nasser 1985; Ashby & Sammis 1990). The principle behind this model is that crack growth under a compressive stress is a stable phenomenon. When a solid is loaded in compression, cracks will nucleate at any of the numerous imperfections within the solid. The wing crack model assumes that cracks nucleate by sliding on a pre-existing, inclined flaw, resulting in the growth of wing cracks at the tips of the flaw. The stress required for the initiation of wing cracks is defined as: (7"1 = {[(1 -'~ 2 ) 1 / 2

~_ jL~]/[(1 -11-/,/2)1/2 __/_L] }O.3

+ {31/2/[(1 +/~2)1/2 _ i~]Kic/(Tra)l/2}

(10)

46

N.J. AUSTIN & L. A. KENNEDY

(Horri & Nemat-Nasser 1985; Ashby & Sammis 1990). The only parameters that influence the onset of wing crack propagation are KIr and/x, and the half length of the initial flaw, a. Therefore, at constant confining pressure, and for constant material properties, Kxc and /a., the stress required for the initiation of wing crack growth is governed solely by the magnitude of a. The peak strength, which is associated with the onset of shear failure, is governed by the coalescence of wing cracks, which, in turn, is sensitive to both the initial flaw length and the concentration of these flaws. For this work, we have chosen to use the two-dimensional, plane strain equation of Ashby & Sammis (1990): 01 -- {Cl + ( C 4 ( ( D / D o ) 1/2 - 1)/

(1 + (TrDo)UZ((D/Do) 1/2 - 1)/ (1 -D1/2)))}o- 3 -F { ( ( D / O o ) 1/2 - 1 -+- 0.1/COS T)l/2/

(1 q-(7rDo)l/Z((D/Do)

1/2

-

1)/

(1 - D1/2))

• (C4/(cos y ) l / 2 ) ( K i c / ( ~ a ) l / 2 ) }

(11)

where

C1 = ((1 + ]/2)1/2 _~_/d~)/((l Jr- ]3,2)1/2 --/d.)

(12)

C4 = 301/2 cos y/((1 + ]./,2)1/2 _ /L/,)

(13)

D = ~r(1 + acos T)2Na Do = -tr(a cos T)2Na

(14) (15)

The onset of strain softening occurs when l / a ~ 0.75 for an optimal orientation of the

initial flaws with respect to o-j (y), of 0.24rr (Horii & Nemat-Nasser 1985). Assuming that the material properties Kic and /z remain constant, and for the optimal l / a ratio and initial orientation of wing cracks, the peak strength of the material is sensitive to both the initial flaw length and initial flaw density (Horii & NematNasser 1985; Sammis & Ashby 1990). Using equations (11)-(15), the variation in peak o-1 for a spectrum of Na and a values was calculated, and the measured peak strength for each experiment was plotted (Fig. 7) using the /z values obtained from Coulomb's relationship (equation 1, Table 2). Plots are generated for a high KIc value for dolomite of 2.47 MPa m 1/2 (Gunsallus & Kulhawy 1984) and for an exceptionally low value of 0.1 MPa m 1/2 to highlight

the consequence of varying Klc, which results for different flaw types (Fredrich et al. 1990). In Figure 7, if either Na or a are prescribed, the other is fixed for a known peak strength and a constant Kic value. Starting with the common assumption that grain size is a proxy for the initial flaw length (Fredrich et al. 1990; Baud et al. 2000), a range of flaw densities per unit area are predicted using the curves for Kic = 2.47 M P a m 1/2 (Table 4). Flaw density in reference to unit area is difficult to analyse microstructurally, so these values are converted to the number of flaws per grain area based on the measured mean grain sizes. This relationship predicts that Niagara samples oh4 and oh5 contain the highest flaw densities per grain (0.60 and 0.51), followed by the Badshot samples (0.36), and Niagara sample ohl (0.28). The Rock Creek samples rc4 and rc5 have the lowest number of flaws per grain (0.07 and 0.10). Microstructural observations from both the undeformed and deformed samples are not consistent with these predictions (Fig. 1). The Badshot samples contain numerous cleavage planes within the grains, which are interpreted to act as initial flaws on a scale much smaller than the grain size, but with a very high density. In the deformed Badshot samples (Fig. 1d) there are numerous intragranular microcracks that are present along cleavage planes. The presence of smaller, more concentrated flaws are consistent with Figure 7; a decrease in flaw size necessitates an increase in flaw density. In both Rock Creek suites, it appears that flaw size scales with grain size, as deformation is predominantly accommodated along grain boundaries (Fig. le), and there are few flaws in the undeformed grains (Fig. lb). The low flaw density is possibly a consequence of grain boundary orientation: there may not be optimally oriented grain boundaries for wing crack initiation and propagation on every grain. In the Niagara samples, initial flaws are inferred to be a combination of pores, grain boundaries, and intragranular cleavage planes, which vary in size from the scale of grains to substantially smaller (Austin et al. 2005) (Fig. lc). Microcracks and fractures make use of all of these features (Fig. l f). Niagara dolomite has a more heterogeneous distribution of initial flaws than either the Badshot or Rock Creek samples, and thus it is difficult to predict which flaws act as the principal initial flaws for brittle deformation. This analysis has significant implications for the broad scale validity of the Hall-Petch relationship to brittle deformation. A HallPetch proportionality may be valid within

THE PEAK STRENGTH OF DOLOMITE (a)

47

(b) (C) 0 6

,x 10 6

10 x 106

\

KIC=2.47 KIc---0.1

~

~

~

0

0.2

0.4

0.6

0.8

(d) 1(

0.4

I

(t3 1( i 10 7

xlO

8

8

7

7

6

6

5

4

5

~

4

3

3

2

-"

~

1

0 0.1

2:

KIc=0.1

2

1

0.2

0.3

0,4

0.5

0.6

, 0.7

, 0.8

0',9

Ol

02

03

04

05

06

07

08

x 1()4

-g

0.8 x 10 3

(e) 7

t07

0.6

x [0 -3

• 10-3

09

(h)

(g) 7

1Gx 10

9

.

6

7

8

x IO 4

9

10

x 10 -5

(i) 7

10 7

1(

10

8 7 6 5 4 3 2 1

0

0".1

' 0.2

' 0.3

0 '. 4

' 0.5

0'. 6

' 0.7

0'. 8

' 0.9

I

O. 1 0 . 2

I

I

I

I

I

I

I

0.3

0.4

0.5

0.6

0.7

0.8

0.9

• 10 -4

011 012 013 014 015 016 017 018 019

x l o -4

x 16 4

!'),o.

3

=.

o

0.1

x 10-4

x 16 4

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.9 X 10-4

a(m) Fig. 7. Plots showing the relationship between the half flaw length (a), the flaw density (Na), and the measured peak strength for each experiment. The curves are plotted at constant peak stress for each experiment, and the vertical lines represent the measured mean grain size for each sample. (a), (b), and (c) are for Badshot dolomite at confining pressures of 25, 50, and 100 MPa, respectively; (d), (e), and (f) are for Niagara dolomite at confining pressures of 25, 50, and 100 MPa, respectively; (g), (h), and (i) are for Rock Creek (RC4) dolomite at confining pressures of 25, 50, and 100 MPa, respectively; and (j), (k), and (I) are for Rock Creek (RC5) dolomite at confining pressures of 25, 50, and 100 MPa, respectively. In all cases the peak strength is plotted using Kic values of 2.47 and 0.1 MPa m 1/2. individual sample suites, as appears to be the case for the Rock Creek samples, and for the A m i n d a v samples studied by Hatzor et al. (1997). If, however, the proportionality constant

is not the same for different formations, as appears to be the case for the current experimental data, then the relationship will not be valid.

48

N.J. AUSTIN & L. A. KENNEDY

Table 4. Calculated flaw densities Sample

N~ (flaws per mm2) Klc = 2.47

Flaws per grain Kic = 2.47

52 44 48 2 94 370

0.28 0.60 0.51 0.36 0.07 0.10

ohl oh4 oh5 cb rc4 rc5

Predictedflawdensitiesand the correspondingnumberof flawsper grain for each of the sample suites based on equation(1), using a Klc value of 2.47 MPa m 1/2 and an initial flaw length (2a) equivalentto the meangrain size.

The variability in strength as a function of texture is enhanced by variations in Kic. As Klc is decreased, the required flaw density for a given flaw size is also decreased, and correspondingly, for a given flaw density, the required flaw size is decreased (Fig. 7). Since Kic varies with flaw type in rocks (Fredrich et al. 1990), the nature of the flaws is also an important parameter in dictating the peak strength. Owing to the plethora of flaws in rocks, it will be a combination of flaw length, flaw density, and the K~c values of the given flaw types that will dictate which flaws are significant for failure, and thus which flaws dictate the peak strength. The theory of wing crack coalescence and shear failure dictates that the initial concentration of flaws plays as important a role in the peak strength of rocks as the initial flaw length. Both of these variables are strongly dependant on the texture of dolomite, and assumptions cannot be made about one without considering the effect on the other. Microstructural observations support the notion that including intragrain and grain boundary textures can dictate both the flaw density and the flaw length, thus varying the peak strength of the rock. Grain size cannot be assumed to be directly related to peak strength without taking textural variations into account.

Conclusion Using the wing crack model, this data set demonstrates that in texturally diverse suites of rocks, a H a l l - P e t c h relationship cannot be assumed a priori. Our experimental data confirm that factors other than grain size and porosity must be considered in modeling the peak strength. Careful observation of intragranular and grain boundary properties must be made in order to understand potential discrepancies. Unfortunately, intragranular and grain boundary textures are difficult to quantify, especially from two-

dimensional images. The effective Young's modulus of materials is, however, a powerful empirical tool for understanding how variations in texture may lead to variations in strength, due to its sensitivity to pre-existing flaws, including cracks and porosity. We have shown that it is possible to fit our data and those of Hatzor et al. (1997) using a relationship that is non-linear with respect to confining pressure, and that takes into account only the effective Young's modulus and the empirically defined uniaxial compressive strength. These results highlight the importance of understanding textural variations, especially when examining the mechanical response of mineralogically similar samples with very different textural properties. Funding for this research was provided by an NSERC operating grant to L. A. Kennedy,a CollaborativeResearch and Developmentgrant to L. A. Kennedy, and a B.P. research grant to L. A. Kennedy, as well as by an NSERC PGS-A scholarship to N. J. Austin and a U. B. C. University Graduate Fellowshipto N. J. Austin. The sample materials were kindly provided by Imasco Minerals, Mighty White Dolomite, and Graymont Dolime. We must thank Hans de Bresser and Ernest Rutter for their constructivereviews, as well J. K. Russell for many helpful discussions, and R. Rodwayfor his assistance in the design and maintenance of the triaxial rock press used in this study.

References ASHBY, M. F. & SAMMIS, C. G. 1990. The damage mechanics of brittle solids in compression. Pure and Applied Geophysics, 133(3), 489-521. AUSTIN, N. J., KENNEDY, L. A., LOGAN, J. M. & RODWAY, R. 2005. Textural controls on the brittle deformation of dolomite: the transition from brittle faulting to cataclastic flow. In: GAPAIS,D., BRUN, J. P. & COBBOLD, P. R. (eds) Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 51-66. BATES, R. L., & JACKSON,J. A. (eds) 1984. Dictionary of Geological Terms. American Geological Institute, New York. BAUD, P., SCHUBNEL, A. 8~ WONG, T. F. 2000. Dilatancy, compaction, and failure mode in Solnhofen limestone. Journal of Geophysical Research, 105(B8), 19 289-19 303. BRACE, W. F. & BOMBOLAKIS,E. G. 1963. A note on brittle crack growth in compression. Journal of Geophysical Research, 68(12), 3709-3713. BRACE, W. F., PAULDING,B. W. & SCHOLTZ,C. 1966. Dilatancy in the fracture of crystalline rocks. Journal of Geophysical Research, 71(16), 3939-3953. ENGELDER, T. 1993. Stress Regimes in the Lithosphere. Princeton University Press. EVANS, B., FREDRICH,J. T. & WONG, T. F. 1990. The brittle-ductile transition in rocks: recent experimental and theoretical progress. Geophysical Monograph, 56, 1-20.

THE PEAK STRENGTH OF DOLOMITE FREDRICH, J. T., EVANS, B. & WONG, T. F. 1990. Effect of grain size on brittle and semibrittle strength: implications for micromechanical modeling of failure in compression. Journal of Geophysical Research, 95(B7), l0 9 0 7 - l 0 920. GOODMAN, R. E. 1980. Introduction to Rock Mechanics. John Wiley & Sons, New York. GRIFFITH, A. A. 1921. The phenomena of rupture and flow in solids. Philosophical Transactions of the Royal Socie~ of London, 221, 163-198. GRIFFITH, A. A. 1924. Theory of rupture. Proceedings of the First International Congress on Applied Mechanics, 1, 55. GUNSALLUS, K. L. t~z KULHAWY,F. H. 1984. A comparative evaluation of rock strength measures. International Journal of Rock Mechanics and Mineral Science, 21(5), 233-248. HATZOR, Y. H., ZUR, A. & MIMRAN, Y. 1997. Microstructure effects on microcracking and brittle failure of dolomites. Tectonophysics, 281, 141 - 161. HATZOR, Y. H. & PALCHIK,V. 1998. A microstructurebased failure criterion for Amindav dolomites. International Journal of Rock Mechanics and Mineral Science, 35(6), 797-805. HOEK, E. & BIENIAWSKI,Z. T. 1966. Brittle fracture propagation in rock under compression. International Journal of Fracture Mechanics, 1, 137-155. HORII, H. & NEMAT-NASSER, S. 1985. Compressioninduced microcrack growth in brittle solids: axial splitting and shear failure. Journal of Geophysical Research, 90(B4), 3105-3125. HUGMAN, R. H. H. & FRIEDMAN, M. 1979. Effects of texture and composition of mechanical behavior of

49

experimentally deformed carbonate rocks. AAPG Bulletin, 63(9), 1478-1489. OLSSON, W. A. 1974. Grain size dependence of yield stress in marble. Journal of Geophysical Research, 79(32), 4859-4862. SAGONG, M. & BOBET, A. 2002. Coalescence of multiple flaws in a rock-model material in uniaxial compression. International Journal of Rock Mechanics and Mineral Science, 39, 229-241. SAMMIS C. G. & ASHBY, M. F. 1986. The failure of brittle porous solids under compressive states. Acta Metallica, 34(3), 511-526. SOEIANOS, A. I. t~z HALAKATEVAKIS, N. 2002. Equivalent tunneling Mohr-Coulomb strength parameters for given Hoek-Brown ones. International Journal of Rock Mechanics and Mineral Science, 39, 131-137. TURNER, F. J., GRIGGS, D. T., HEARD, H. & WEISS, L. W. 1954. Plastic deformation of dolomite rock at 380 degrees C. American Journal of Science, 252, 477-488. WALSH, J. B. 1965a. The effect of cracks on the compressibility of rock. Journal of Geophysical Research, 70(2), 381-389. WALSH, J. B. 1965b. Some new measurements of linear compressibility of rocks. Journal of Geophysical Research, 70(2), 391-398. WALSH, J. B. 1965c. The effect of cracks on the uniaxial elastic compression of rocks. Journal of Geophysical Research, 70(2), 399-411. ZHANG, J., WON6, T. F. & DAVIS, D. M. 1990. Micromechanics of pressure-induced grain crushing in porous rocks. Journal of Geophysical Research, 95(B 1), 341-352.

Textural controls on the brittle deformation of dolomite: the transition from brittle faulting to cataclastic flow N. J. A U S T I N 1'3, L. A. K E N N E D Y l, J. M. L O G A N 2 & R. R O D W A Y 1

1Earth and Ocean Sciences, The University of British Columbia, 6339 Stores Road, Vancouver, BC, V6T 1Z4, Canada 2Department of Geological Sciences, The University of Oregon, Eugene, OR, 97411 3present address: Earth, Atmospheric, and Planetary Sciences, The Massachusetts Institute of Technology, 77 Massachusetts Avenue, 54-710, Cambridge, MA, 02139-4307, USA (e-mail: naustin @mit. edu) Abstract: To investigate the role of texture on the brittle deformation of dolomite, 23 triaxial

deformation experiments were performed at confining pressures of 25, 50, and 100 MPa, dry, at room temperature, on dolomite from three texturally distinct sample suites. The variations

in the mechanical response of these mineralogically and chemically similar dolomites, and the ensuing microstructures, indicate that grain boundary textures promote or inhibit the ability of grains to shear and rotate with respect to one another, whereas the presence of intragranular flaws, such as cleavage, that act as weaknesses, promote intragranular deformation. In samples with porosities greater than c. 7%, inelastic pore collapse controls the transition from brittle faulting to extensive intragranular deformation and cataclastic flow. This porosity is much higher than has been observed at the onset of pore collapse in calcite, as a consequence of the inability of dolomite to deform by crystal plastic processes at room temperature. Combined, these textural features may dictate the transition from brittle faulting to cataclastic flow in brittle rocks in the upper crust. Rocks are inherently heterogeneous, and thus assumptions are required in order to understand which properties are significant in controlling their mechanical and microstructural response to differential stress in the brittle field. Of particular interest is the influence of rock texture on the microstructural evolution of dolomite. Here, texture is used to mean the geometric aspects of, and the mutual relations among, the rocks' component particles or crystals (Bates & Jackson 1984). Generally, the two textural properties that are considered in analysing the brittle deformation of rocks are porosity (Brace 1978; Hatzor et al. 1997; Baud et al. 2000) and grain size (Olsson 1974; Hugman & Friedman 1979; Fredrich et al. 1990). The roles of intragranular flaws and grain boundary textures are less often examined, despite the highly variable nature of both of these properties, and their effect on the distribution, concentration, and length of initial flaws. The concentration and length of the initial flaws are the principal parameters in fracture theory (Griffith 1921; Nemat-Nasser & Horii 1982; Horii & Nemat-Nasser 1985; Ashby &

Sammis 1990), and as such, must be accounted for in order to understand the brittle deformation of rocks. The transition from deformation by brittle faulting to cataclastic flow is of particular interest, as this transition is known to affect the permeability and transport properties of rocks (Zhu & Wong 1997). While this transition has been extensively addressed with relation to sandstones (Rutter & Hadizadeh 1991; Menendez et al. 1996; Zhu & Wong 1996, 1997; Wong et al. 1997) and limestones (Fredrich et al. 1990; Baud et al. 2000) there has been very little work done on dolomite. Dolomite is commonly associated with limestones, and thus to understand how these materials behave as multiphase rocks, the mechanical responses of the homogeneous end-members must be understood. Dolomite is also a common hydrocarbon reservoir rock because it commonly contains vuggy porosity, and is frequently extensively fractured (Antonellini & Mollena 2000). We have performed a series of triaxial rock deformation experiments at room temperature on texturally diverse dolomites covering a wide

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 51-66. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

N.J. AUSTIN ET AL.

52

range of porosities, grain sizes, and grain and grain boundary textures, with the aim of elucidating the relationship between texture and the brittle deformation of dolomite. Our analysis focuses predominantly on the microstructural response of these rocks to a differential stress. The results are compared primarily to the extensive experimental data set for quartz and calcite in order to highlight the importance of mineral and textural properties on the mechanical and microstructural response of rocks to deformation.

Methodology Apparatus description The Large Sample Rig (LSR) triaxial rock press used for this research (Fig. 1) is a significantly modified version of one previously used at Texas A&M University (Handin et al. 1972, 1986; Shimamoto 1977). It is a vertically mounted apparatus capable of confining pressures up to 300 MPa, with an argon confining medium. Confining pressure is monitored with an analogue gauge, with precision of + 2 MPa. Maximum displacements of 101.6mm may be obtained at displacement rates up to 1.8 • 1 0 - 2 c m s -1, with a maximum differential force of 980 665 N. Position is measured using a Trans Tek 0246 0000 displacement transducer, with precision of 0.01 mm. Force is measured with an external load cell with an observed precision of better than 1 kN, which translates to a stress of 2.65 MPa on a sample with a diameter

of 22.2 mm. Mechanical data is monitored digitally using Labview Software.

Sample assembly. The LSR can accommodate right cylinder samples of two sizes: (1) 2 2 . 2 m m x 50.8ram, with ends parallel to better than 0.0127 mm (used in this study), or (2) 47.6 mm x 95.3 mm, with ends parallel to better than 0.254 ram. One hardened H13 steel spacer is placed between the sample and each of the upper and lower pistons. A 1.588 mm diameter hole down their centre and a groove pattern on each end (Fig. 1) ensures that experiments are run under drained conditions, and may also be used to facilitate the even application of fluid pressure (not used in these experiments) across the ends of the sample. The sample assembly is jacketed in two wraps of polyolefin heat shrink tubing, sealed onto the ends of the sample assembly with twisted steel wire. While argon has been observed to diffuse through polyolefin, this has not been an issue at the timescale of the experiments. Calibration. The stiffness of the LSR was measured based on the spring analogy of Shimamoto (1977), which assumes that any extension of the tie rods is insignificant. Measurements were made by deforming fight cylinders of H13 steel ( E = 206.84GPa) and 6061-T6 aluminum (E = 9.73 GPa) (Table 1) in each of the sample assemblies (Table 2). These metals were chosen due to their significantly

Large Sample ~

Small Sample

External Load Cell /.x-----a

H13 Steel .__.... _ _ _ _ H 13 Steel

F

Sample

D I ~ I

~[i

~

H13L1 ~ ~

H13Steel

e

Fig. 1. A schematic of the triaxial rock press used in this study, highlighting the locations of the external load cell and displacement transducer (DCDT). In succession to the right are close-up views of the pressure vessel, and the samples assemblies for both the large and small samples sizes.

BRITTLE FAULTING TO CATACLASTIC FLOW Table 1. List of symbols Symbol o-~ o-3 o-p o-y o-m

Pc P* 4) dm dx E Eel Pm Pb e rate K~ K~c

Description Principal compressive stress Least compressive stress Peak differential stress Yield stress Mean stress Confining pressure Critical pressure for grain crushing Porosity Mean grain size Maximum grain size True Young's modulus Effective Young's modulus Matrix density Bulk density Strain rate Stress intensity factor Critical stress intensity factor

different elastic properties, allowing the stiffness of the LSR to be explored at both low and high loads. The stiffness is consistently lower at low loads, whereas the variance in stiffness is greater at low loads (Table 2), explaining the observed non-linearity at the onset of experimentally derived stress-strain curves. All reported strain data accounts for the average stiffness of the appropriate sample assembly as reported in Table 2.

Starting material Natural dolomite samples were obtained from the Imasco Minerals quarry in the Badshot Formation of southeastern British Columbia, the Graymont Dolime quarry in the Niagara Formation in Ohio, and the Mighty White Dolomite quarry near Rock Creek, British Columbia.

53

Chemistry and mineralogy. Bulk chemical composition was analysed for each of the three sample suites using X-ray fluorescence. Chemical analyses were obtained from four independent blocks of Badshot dolomite, from three independent blocks of Rock Creek dolomite, and from three independent blocks of Niagara dolomite (Table 3). CaO and MgO contents are consistent with nearly stoichiometric dolomite for all samples (Table 3). The CO2 content was estimated from loss on ignition (LOI) and is also nearly stoichiometric. The Badshot and Niagara samples have no significant concentrations of oxides other than CaO and MgO; the Rock Creek samples contain 7.28 moles per kg of Si02 as opposed to 56.1 and 48.2 moles per kg of CaO and MgO respectively. The mineralogy of each of the three sample suites was determined from optical thin section observation (Fig. 2), XRD, and Rietveld refinements (Raudsepp et al. 1999). Badshot samples are composed of > 9 7 vol% dolomite with accessory quartz, calcite, tremolite, and muscovite; Niagara samples are composed of > 9 5 vol% dolomite; and Rock Creek samples are composed of > 8 5 vol% dolomite, with accessory calcite and quartz (located in localized millimetre scale veins) and talc. Textural analysis: Grain size. Grain size is presented as the arithmetic mean of the equivalent circular diameter based on measured grain areas (Table 4, Fig. 3). Grain boundary maps, obtained by manually tracing thin section images, were imported into Scion Image and the grain areas were measured Hand sample scale variations were accounted for by calculating mean grain size values for each block. Maximum grain sizes were obtained by

Table 2. The stiffness of the LSR Sample assembly

Pc (MPa)

Material

Average stiffness (kg cm- 1)

Standard deviation

Standard deviation (%)

Maximum load (kg)

Small Large Small Large Small Small Small Large Large Large

0.1 0.1 0.1 0.1 24 52 107 24 52 107

Al Al H13 H 13 H 13 H 13 H13 HI3 HI3 H13

202170 337945 246027 393048 261561 283565 324325 386464 393504 396767

8806 20342 3876 2427 8321 6999 9754 3295 989 3125

4.36 6.02 1.58 0.62 3.18 2.47 3.01 0.85 0.25 0.79

5987 26444 22680 72575 22680 22680 22680 72575 72575 72575

The stiffnessof the LSR for both the large and small sampleassembliesat varyingconfiningpressures and for varyingloads, alongwith the standard deviation,obtainedfrom 15 duplicate measurements. A1refers to 6061-T6 aluminum,and HI3 refers to H13 tool steel.

54

N. J. AUSTIN ET AL.

o

~

~

~

~

~

~

~ ~ d d d d d d

eq ~.--,

O~

O~.

ddddcSddddcSdddddodddddddd

O~ ~, . .-;

dddd

q q q o o o q q o q q q q q o q q o q o q q q q q q q q q q ,.o.~

o~

z~

z~

ooooo~ooooooo

8 ~~

,.-.,

O~-

d~

zs t-,,

Q~

~g ea~

o

,~

~

.~'N [.- o

d d d d d d

. . . . . .

o d d d d d d d d d d d d d

~

,=

=

~ o~.~ ~

~

~

~

~

~

~

~

0 0 ~

BRITTLE FAULTING TO CATACLASTIC FLOW

(a)

(b)

(c)

(d)

55

Fig. 2. Optical micrographs of dolomite samples from each of the three formations examined in this study. (a) Badshot dolomite consisting of coarse grains with lobate grain boundaries (L) and finer polygonal grains (P). Note the presence of cleavage and twinning in almost all grains. (b) Rock Creek sample RC4 with rhombohedral grains and straight grain boundaries (S), and flaws within the grains (F). (c) Rock Creek sample RC5 illustrating the irregular grain boundaries that may result from pressure solution (PS), and the low concentration of flaws within the grains. (d) Niagara dolomite, with inclusion filled grains, giving a 'dirty' appearance, straight grain boundaries (S), and moderately developed cleavage within the grains.

averaging the 10 largest grains measured in each block.

Textural analysis: Porosity. Total porosity was determined by independently measuring the bulk and matrix density for each core (Table 4). The matrix density was measured with a Micromeritics multivolume pycnometer 1305 with a precision of 6.9 x 10 -5 MPa (0.01 PSI) and a Mettler H 2 0 scale with a precision of 0.00001 g. Measurements were performed on three independent blocks of Badshot dolomite, and on each of the blocks of Niagara and Rock Creek dolomite that were used for experiments. The bulk density was measured for each core individually. Volume was obtained by measuring

the diameter of the core at three points along its length, and the length at two different orientations using calipers with a precision of 0.01 mm, and the mass was measured using a Mettler P1200 scale with a precision of 0.01 g. The relationship (Pin -- Pb)lPm -- ~

(1)

was used to obtain porosity (Table 4). All of the Badshot and Rock Creek samples have porosities less than 1.5%, whereas the Niagara samples have varying porosities between 6.6 and 13.5%.

Textural analysis: Hand sample and grain scale textures. Badshot dolomite has been

56

N.J. AUSTIN ET AL. ,"4 ~

I

1

|

1

~

1

1

1

1

1

1

1

|

1

1

1

1

1

1

1

1

1

1

9

I

~.g

~o ~.~

~.o

~9 +~

~o

o~

m

BRITTLE FAULTING 0.16

TO CATACLASTIC

(a) CBB1

-

0.12 -

0.08

0.08

[

-.

0

0.16

-

1

~1'1'1'1' I'i'l'l'l' 0.4 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 4.8

0(

0.16

0.12

0.12

0.08

0.08

0.04

0.04

0.16 -

0.12

-

004.

(c) RC5

0

57

0.16- (b) RC4

0.12

0.04

FLOW

0

0.16 -

(e) O H 4

"L ' I'1'1' I' I' I'1' 2.4 2.8 3.2 3.6 4 4.4 4.8

(d) OH1

-

'l' ' l ' l ' l ' l ' l ' l ' l ' l ' l ' l ' 0.4 0.8 1.2 1.6 2 2.4 2.8 3.2 3.6 4 4.4 4.8

'1 '1'1 0.4 0.8 1.2 1.6

'J' 'J' I' I'l' I'1'1' 0.4 0.8 1.2 1.6 2 2.4 2.8 3.2 3.6 4 4.4 4.8

(f) OH5

0.12

-

0.08 -

0.08

0.04-

0.04-

0

' I' I' I'1'1' 0.4 0.8 1.2 1.6 2

' I' I'1'1' 2.4 2.8 3.2 3.6 4

I' I' 4.4 4.8

'1'1' I'1'1'1'1' 'l'l'l'l' 0.4 0.8 1.2 1.6 2 2.4 2.8 3.2 3.6 4 4.4 4.8

Log Grain Size (gm) F i g . 3. G r ain size h i s t o g r a m s plotted as log grain size, against r e l a t i v e f r e q u e n c y for (a) B a d s h o t d o l o m i t e b l o c k C B B 1 , (b) R o c k C r e e k b l o c k R C 4 , (e) R o c k C r e e k b l o c k R C 5 , (d) N i a g a r a b l o c k O H I , (e) N i a g a r a b l o c k O H 4 , and (f) N i a g a r a b l o c k O H 5 .

58

N.J. AUSTIN ET AL.

metamorphosed to amphibolite facies and subsequently annealed (Colpron et al. 1996). At the hand sample scale, the rock is isotropic, except for minor, randomly oriented tremolite porphyroblasts, and weakly aligned muscovite. At the thin section scale, coarser grains have lobate grain boundaries, exhibit minor to no undulose extinction and few deformation twins, although some well-developed subgrains are present (Fig. 2a). There is no noticeable shape preferred orientation. Growth twins are common, as is the presence of well-developed {10i 1} cleavage (Fig. 2a). Finer grains are predominantly polygonal, with no undulose extinction, very rare twinning, and poorly developed cleavage. These are inferred to be the product of dynamic recrystallization (Fig. 2a). Dolomite from the Rock Creek quarry is massive, with volumetrically minor amounts of variably oriented quartz and calcite veining along with sealed cracks visible in hand sample. In thin section, grains are predominantly subhedral and equigranular (Fig. 2b, c), with no noticeable shape preferred orientation, although sample RC4 generally has more rhombohedral grains and straighter grain boundaries (Fig. 2b) than RC5 (Fig. 2c). Twinning and {10il} cleavage are very rare within grains, as are intra- and transgranular microcracks (Fig. 2b, c). There are intragrain flaws, although these are more prevalent in the coarser RC4 block than in the RC5 block (Fig. 2b, c). Niagara dolomite is also massive, but has extensive porosity, visible as variably sized vugs in hand sample. In thin section, the anhedral to subhedral grains, which often have straight grain boundaries, contain opaque inclusions, giving them a dirty appearance (Fig. 2d), and making individual grains difficult to distinguish. As with Badshot and Rock Creek dolomite, Niagara dolomite has no shape preferred orientation. {10il} cleavage is common (although less so than in the Badshot samples), but twinning is rare (Fig. 2d). Pores are generally the size of the grains, and are surrounded by cleavage surfaces, resulting in sharp tips around pores. Experiments

A total of 23 experiments were performed on dolomite from the three locations described above, at confining pressures of 25, 50, and 100MPa, at a temperature of 296K, and a strain rate of c. 1 x 10 -5 s -1 (Table 4). Badshot samples were obtained from one block (CBB1), and were treated as one sample suite. Rock Creek samples were obtained from two

distinct blocks (RC4 and RC5), with each block being treated as an individual sample suite. The Niagara samples were obtained from three blocks, spanning porosities from 6.6% to 13.5%; however, all Niagara samples were treated as a single sample suite as at most two cores could be obtained from any individual block. Experiments were designed to ensure that, to the extent possible, Niagara samples with similar textural properties (specifically porosity) were examined at each confining pressure (Table 4).

Results Mechanical

The peak and yield stresses of all 23 experiments are presented in Table 4, along with the experimental conditions and sample properties. Effective Young's modulus was obtained by isolating the linear region of the differential stress-axial strain plot, fitting this with the best-fit line, and measuring the slope. The yield stress was obtained by picking the point where the measured differential stress-axial strain curve deviated from the linear fit used to obtain the effective Young's modulus, while the peak stress was obtained from the maximum stress on the differential stress-axial strain curve. All sample suites have a progressive increase in peak strength with increasing confining pressure (Fig. 4). Badshot dolomite exhibits ductility at all confining pressures (Fig. 4a); at confining pressures of 50 and 100 MPa, it has a stickslip response, with the period between events decreasing with increasing confining pressure. One sample deformed at 100 MPa maintains an apparent steady stress state between ~4 and 5.5% strain. Within the Rock Creek samples, RC4 samples (Fig. 4b) are consistently weaker than RC5 (Fig. 4c); however, this difference decreases with increasing confining pressure. All Rock Creek samples have an elastic-brittle response and exhibit very little ductility. Mechanically, the Niagara samples are the most variable (Fig. 4d). At 25 MPa, both samples (with porosities of 12.8 and 13.5%) exhibit significant ductility followed by strain softening. At 50 MPa, the two samples respond very differently; sample OHI-1 ( r exhibits an elastic-brittle stress-strain response very similar to the Rock Creek samples, whereas sample OH5-3 (~b = 12.8%) exhibits rapid yielding followed by gradual strain softening. At 100 MPa, OH5-1 (~b = 9.4%) exhibits significant ductility and strain softening, with a minor stickslip response.

600]

BRITTLE FAULTING TO CATACLASTIC FLOW (a)~ 600]

59

CBB1-4 (103 MPa)

400-

~4001

RC4-2 (103 MPa)

4-7 (52 MPa)

200.

~ 200 / ,CBB1-2 (24 MPa) CBB1-6 (24 MPa)

~

RC4-5(51 MPa)

j ~RC4-6 (25 MPa) "l H , ~ --" R_C4-3(24 MPa)

0[ 0

(c)

2

Strain (%)

4

0

6

600"

(d)

RC5-1 (103 MPa) C5-4 (104 MPa)

,,..,,

~, 400'

2

Strain (%)

4

6

600 =

OHI-1 (52 MPa)

400"

r~ m o~

~

~k " ~ R C 5 - 5 (52 MPa) !k RC5-7(51 MPa)

200'

OH4-1 (24 MPa) OH5-1(103 MPa) ( ~

200"

~RC5-3 (24 MPa) ~RC5-2 (24 MPa)

OH5-3 (52 MPa)

OH4-2 (24 MPa) ,

i

0

2

-

Strain (%)

i

i

4

6

0

~

0

2

Strain (%)

4

Fig. 4. Differentialstress-axial strain curves for all 23 experiments: (a) Badshot dolomite (CB samples), (b) Rock Creek, block RC4, (c) Rock Creek, block RC5, and (d) Niagara Dolomite (OH1, OH4, and OH5 samples). The confiningpressure of each experiment is indicated in brackets.

Macro and micro structures

The relationship between confining pressure, sample suite, and fracture orientation is illustrated in Figure 5. Fracture orientations were measured by manually tracing sample scale faults, and measuring the angle between the fault and the long axis of the core, which was parallel to the principal compressive stress during deformation. For both the Badshot and Rock Creek samples the mean angle between major fractures and o'1 decreasd as confining pressure was increased from 25 to 100 MPa, whereas in the Niagara samples, major fractures are rare, and have varied orientation.

All Badshot samples contain one major fault at the hand sample scale, accompanied by crumbling at the sample ends, possibly as a result of friction between the sample and the spacers. There is generally a small amount of comminuted material within the fault zones. The exception to this is sample CBB 1-7, deformed at 50 MPa, which has damaged ends, but contains no major faults or cracks. In thin section, deformed Badshot samples contain faults and comminuted material within the fault zone (Fig. 6a, b); transgranular microcracks, which cut across grains and make use of grain boundaries, generally in a subvertical

60

N.J. AUSTIN ETAL. 70-

-

60.

O Badshot o Niagara

50.

tx Rock Creek (RC4) • Rock Creek (RC5)

B o

&

40.

<

30.

.~ 2o. ~

•

•

10.

0

20

40

60

80

100

120

Confining Pressure (MPa)

Fig. 5. A graphical depiction of the variation in fracture orientation with varying confiningpressure for each of the sample suites.

orientation or subparallel to the faults (Fig. 6b); and intragranular microcracking (Fig. 6a, b). At P c = 25MPa intragranular microcracking is more prevalent in the vicinity of faults, along cleavage planes within the grains, whereas at Pc = 50 MPa and 100 MPa, the extent of intragranular microcracking is greater throughout the sample. At Pc = 100 MPa the extent of intragranular microcracking is such that individual grains can often not be identified. While faults and transgranular microcracks tend to cut across grains, intragranular microcracks are present along crystallographic weaknesses, and thus terminate, or change orientation at grain boundaries due to variations in grain orientation (Fig. 6a, b). All Badshot samples, regardless of the confining pressure, are in the transitional zone between brittle faulting and cataclastic flow. Rock Creek samples all have similar characteristics at the hand sample scale, regardless of the conditions of deformation. All samples contain faults, the orientations of which are shown in Figure 5. In most cases, faults are accompanied by extensive damage at the ends of the samples, again interpreted to be a result of friction between the spacers and the samples. In thin section, Rock Creek samples deformed at 25, 50, and 100 MPa all exhibit similar microstructures consisting of transgranular microcracks along grain boundaries, resulting in dilation, and faults (Fig. 6c, d). Intragranular deformation is observed within one to two grains of the faults, and extensive comminution

of grains has occurred within the faults (Fig. 6c, d). Distal to the faults, dolomite grains are undeformed; however, at intermediate distances to the faults, there is dilation along grain boundaries as a result of grain rotation and shear (Fig. 6c, d). Coarse-grained calcite in veins exhibits intragranular deformation. All Rock Creek samples have deformed in the brittle faulting regime. Niagara samples exhibit only very minor deformation at the hand sample scale. In both samples deformed at 25 MPa this deformation is localized at the ends, again likely due to friction between the spacer and the sample. At P c - - 50 MPa (OH5-3), a single, diagonal fracture is observed, but no other significant deformation can be seen in hand sample. At 100 MPa (OH5-1) there is extensive damage in the middle of the sample, and one diagonal fault. Intragranular deformation is observed to various extents throughout all Niagara samples (Fig. 6e, f). Grains that surround pores form damage zones, consisting of intense intragranular deformation, with microcracks making use of cleavage planes, resulting in comminution of grains. Away from pores, intragranular deformation is suppressed, resulting in textures reminiscent of the undeformed samples. Transgranular microcracks are common, generally controlled by porosity, and often linking pores. Between pores, microcracks are present both along grain boundaries and across grains (Fig. 6e, f). Where pores are favorably oriented,

BRITTLE FAULTING TO C A T A C L A S T I C FLOW

(a)

(b)

(c)

(d)

(e)

(f~

Fig. 6. Images of the deformed samples. (a) and (b) are optical micrographs of Badshot dolomite sample CBB1-9 deformed at P c = 100 MPa. Intragranular deformation along cleavage planes (I), coupled with the absence of major transgranular fractures, is visible in (a), whereas in (b) a sample scale fault is present (F), as are transgranular microcracks (T) which run subparallel to the fault and cut across grains. (e) and (d) are secondary electron SEM images from Rock Creek sample RC5-1, also deformed at P c = 100 MPa. The rotation of grains relative to one another has led to dilation along grain boundaries (D). Comminution of grains (C) is visible along the fault boundary. Away from faults, transgranular microcracks are present along grain boundaries (T), and the grains themselves are generally unreformed. (e) and (f) are SEM images from Niagara samples OH4-2 and OH5-1, deformed at P c = 25 and 100 MPa respectively. In (e) intragranular deformation (I) and comminution of grains are observed in the vicinity of the pores (P), while transgranular cracks are present along grain boundaries (T). In (f), transgranular cracks (T) running subparallel to crj are observed in the vicinity of faults, which link with a large pore just out of the image. There is also extensive intragranular deformation (I). The orientation of the principal compressive stress is indicated on the figure.

m i c r o c r a c k s parallel to cr~ step across pores, r e s u l t i n g in an e n - e c h e l o n g e o m e t r y (Fig. 6e, f). T h e r e is n o variation in m i c r o s t r u c t u r e s w i t h v a r y i n g c o n f i n i n g p r e s s u r e in the N i a g a r a

s a m p l e s . S a m p l e O H I - 1 d e f o r m e d in the brittle faulting r e g i m e , b a s e d o n its m e c h a n i c a l r e s p o n s e , w h e r e a s the rest o f the N i a g a r a s a m p l e s are in the c a t a c l a s t i c flow r e g i m e .

N . J . AUSTIN ET AL.

62

The transition from trans- to intragranular deformation The onset o f grain crushing in low-porosity d o l o m i t e

The transition from trans- to intragranular deformation between chemically and mineralogically similar, low-porosity Rock Creek and Badshot dolomites must be a function of some combination of the textural properties of the respective formations. In the Rock Creek and Badshot samples (all with porosities below 1.5%) deformed at confining pressures at or below 100 MPa, the transition from brittle faulting to cataclastic flow is relatively insensitive to variations in mean stress (Fig. 7). Both the Rock Creek and Badshot samples show a continuous and linear increase in peak differential stress with increasing mean stress, indicating that deformation is accommodated by brittle failure and/or dilatant flow (Wong et al. 1997; Baud et al. 2000). Furthermore, with the exception of a slight increase in the degree of intragranular deformation in Badshot samples deformed at 100 MPa, there is very little variation in microstructure as confining pressure, and thus mean stress, is varied. The role of grain size on the onset of hydrostatic grain crushing, which corresponds to the onset of intragranular deformation under a hydrostatic stress, was analysed by Zhang et al. (1990) based on Hertzian fracture mechanics. The principle behind this analysis is that particles

are in point contact, with a force applied normal to the contact. Owing to elastic deformation of the grains, the point of contact will expand into a circular area of contact, with the maximum hoop stress located at the perimeter of the contact. Grain crushing will occur when the externally applied pressure becomes sufficiently high such that the hoop stresses at the point contacts are large enough for Kt to exceed Kic. For randomly packed, variably sized spherical grains, the proportionality p* oc (t~dm)n

(2)

with n = - 3 / 2 , was theoretically obtained by Zhang et al. (1990) and was found to be in good agreement with data for consolidated sandstone, unconsolidated sand, and glass spheres, as well as for both naturally and experimentally deformed sandstones and quartzites (Wong et al. 1997). In using this relationship to analyse our data, mean stress was used as was done by Zhang et al. (1990) for the data of Hirth & Tullis (1989). Consequently, the critical pressure for grain crushing is lower due to the tangential contact forces at grain boundaries generated by the differential stress (Zhang et al. 1990; Papamichos et al. 1993; Menendez et al. 1996; Wong et al. 1997). The Rock Creek and Badshot samples plot in distinct regions of ~bdm-mean stress space (Fig. 8). Rock Creek samples, which deformed almost exclusively by brittle faulting and transgranular

600 •

o Badshot o Niagara A Rock Creek (RC4) x Rock Creek (RC5)

500

•

"~ 400.

IF •

S

"d 300.

~

200.

dl

.'L

~

A

100.

0 0

50

100

150 200 Mean Stress (MPa)

250

300

350

Fig. 7. The relationship between the peak differential stress and the mean stress for each experiment. Open symbols and x symbols indicate samples that deformed by brittle faulting, grey shading indicates samples that deformed in the transitional regime between brittle faulting and cataclastic flow, and black shading indicates samples that deformed by cataclastic flow.

BRITTLE FAULTING TO CATACLASTIC FLOW

63

1000000100000.

•

~K

O Badshot 13 Niagara

Cataclastic Flow

a RockCreek o Oughtibridge Ganister (Hirth and Tullis, 1989)

10000 1000-

nil

I

e

o..) 100. 10-

0.1 0.1

Brittle Faulting

]

l0 x dm (lam)

100

1000

Fig. 8. An examination of the grain crushing theory of Zhang et al. (1990) based on a plot of ~bdin v. mean stress for all of the experiments. The upper and lower envelopes have theoretically obtained slopes of - 3/2 (Zhang et al. 1990) and require constants of 5000 and 100 respectively 9Open symbols indicate samples that deformed by brittle faulting, grey shading indicates samples that deformed in the transitional regime between brittle faulting and cataclastic flow, and black shading indicates samples that deformed by cataclastic flow.

microcracking, inhabit a region of much lower ~b dm values than the Badshot samples, which deformed by a combination of brittle faulting and intragranular microcracking, consistent with Hertzian fracture theory (Zhang et aL 1990). Based on the transition in microstructures associated with varying grain size, at approximately constant porosity, equation (2) was used to obtain an estimate of the critical mean stress required for the transition from brittle faulting and transgranular microcracking, to cataclastic flow accommodated by intragranular deformation in dolomite (Fig. 8). Proportionality constants of 100 and 5000 were required to empirically fit the theoretical relationship in equation (2) to our observed data for the upper and lower stress limits of the transitional zone between brittle faulting and cataclastic flow (Fig. 8). The predicted pressures required for grain crushing in dolomite-rich rocks are much lower than predicted by Wong et al. (1997) for experimental data compiled for quartz-rich rocks, where the proportionality constants are 4743 and 9486 for the same limits to the transitional zone. The onset of grain crushing occurs at a lower axial stress during non-hydrostatic loading than required in hydrostatic loading (Papamichos et aL 1993; Wong et al. 1997); however, the deformation experiments of Hirth & Tullis (1989) on quartzite require a

proportionality constant of 8400, which while being lower than for hydrostatically loaded quartz and quartzite, is much higher than we observe for dolomite. The occurrence of grain crushing at lower mean stress than in quartzite is observed to be enhanced by the presence of {10[1} cleavage in dolomite, which microstructures indicate promotes the onset of intragranular deformation and cataclastic flow (Fig. 6a, b), consistent with the observations of Tullis & Yund (1987) on feldspar, where the ease of cracking along cleavage planes promoted cataclastic flow. The importance of the size and textural properties of grains on the transition from intragranular deformation to transgranular fracturing has been previously recognized in dolomite. Hugman & Friedman (1979), for example, observed that experimentally deformed anhedral crystalline Hasmark dolomite contained extensive intragranular microffactures, whereas micritic, euhedral Blair dolomite contained very few microfractures.

The influence o f grain and grain-boundary textures

Under a triaxial stress state, the transition from trans- to intragranular deformation is a function of the ability of grains to rotate and shear past

64

N.J. AUSTIN ET AL.

each other, which is commonly assumed to be controlled by the mean stress (Menendez et al. 1996). In our experiments, there is no clear change in the mode of deformation with increasing mean stress. Qualitatively, however, there is a relationship between the degree to which grain boundaries are interlocked (Fig. 2) and the degree of intragranular deformation. It, therefore, appears that this transition is promoted as a function of texture rather than mean stress. Badshot dolomite, with lobate grain boundaries, is incapable of deforming readily by grain rearrangement, but due to the presence of {10il} cleavage, it can readily accommodate intragranular microcracking. In contrast, Rock Creek samples have rhombohedral grains with relatively straight grain boundaries, which can readily deform by shearing along grain boundaries, leading to grain rotation, transgranular cracking, and deformation by brittle faulting, yet have very few intragranular flaws, thereby inhibiting intragranular deformation. The importance of texture on the transition from brittle faulting to cataclastic flow in dolomite, unlike what is commonly observed in quartz and calcite, is a consequence of the different properties of the minerals. Quartz does not contain well-developed cleavage planes and, therefore, grain boundaries act as the principal flaws over a broad range of textures, which enhances the relationship between pressure and the onset of intragranular deformation (Menendez et al. 1996). In calcite, which like dolomite contains perfect {10il} cleavage, fractures are strongly cleavage controlled at hydrostatic stresses above that required for grain crushing (Zhang et al. 1990). Calcite, however, is more susceptible to crystal plastic deformation than dolomite (Turner et al. 1954; Fredrich et al. 1990; Baud et al. 2000). The onset of extensive intragranular deformation in calcite is strongly linked to the pressure sensitivity of the onset of crystal plastic deformation mechanisms. Intragranular cracking has been observed to occur as a result of interaction between brittle and plastic deformation mechanisms, particularly twins (Fredrich et al. 1989). Twinning has not been observed to occur in dolomite at 296 K regardless of pressure (Higgs & Handin 1959; Barber et al. 1981). Basal C slip may be active in dolomite at 296 K (Handin & Fairbairn 1956); however, significant ductility associated with crystal plastic mechanism has not been observed below temperatures and pressures conducive for twinning (Higgs & Handin 1959; Barber et al. 1981). Consequently, a relationship between pressure and the onset of intragranular deformation related to crystal plastic deformation is not expected in dolomite.

The transition from deformation by transgranular microcracking and faulting to intragranular deformation and cataclastic flow is strongly influenced by grain and grain boundary textures in dolomite. Rock Creek samples, which deform by transgranular microcracking and faulting, have little to no pre-existing cleavage or twinning within the grains and contain predominantly straight grain boundaries. The principal weaknesses within the rock are grain boundaries, which allow grains to shift relative to one another, leading to the formation of transgranular grain boundary microcracks and faults. Conversely, grain boundaries in the Badshot samples are lobate, and undeformed grains contain numerous cleavage and twin planes. The principal weaknesses are the cleavage planes within the grains, and thus intragranular deformation predominates, resulting in cataclastic flow. The influence o f p o r o s i t y

Niagara samples with porosities greater than 7% are substantially more ductile than that with a porosity of 6.6% (OHI-1) (Fig. 4d). Samples with porosities greater than 7% plot higher in 05 dm space than any of the Badshot samples (Fig. 8), indicating that they are more likely to exhibit intragranular deformation, whereas that with a porosity of 6.6% plots amongst the Badshot samples. The intragranular deformation predicted from Figure 8 for samples with 05 > 7% occurs preferentially in damage zones, in the vicinity of pores. In vuggy rocks deformed under a differential stress, pores act as stress risers, promoting deformation in neighbouring grains. This is enhanced by the presence of cleavage, which causes irregularities in the surfaces of pores from where cracks can initiate and propagate. The lower porosity Niagara sample (OHI-1) plots amongst the Badshot samples in Figure 8, despite its significantly more elasticbrittle mechanical behaviour, a difference attributed to its finer grain size than the Badshot samples, coupled with straight grain boundaries that act as weaknesses, promoting transgranular deformation. The varying mechanical response of the Niagara samples is interpreted to be the result of pore collapse (Brace 1978; Gowd & Rummel 1980; Logan 1987; Rutter & Hadizadeh 1991; Baud et al. 2000). In Solnhofen limestone, however, Baud et al. (2000) observed the onset of pore collapse in samples with porosities as low as 3% whereas this mechanism is not observed in Niagara samples until porosities greater than c. 7%. The difference between our experiments on dolomite and those of Baud et al. (2000) on

BRITTLE FAULTING TO CATACLASTIC FLOW Solnhofen limestone may, as discussed in relation to grain size, be related to the ability of calcite to deform by crystal plastic processes at room temperature, a distinction they noted in comparing their experimental results to previous work on silicate rocks (Tullis & Yund 1992; Hirth & Tullis 1994). The stress intensification around pores can readily reach the critical resolved shear stress for twinning or slip in calcite, whereas this is not possible for dolomite at room temperature (Turner et al. 1954). The onset of crystal plasticity leads to crack tip blunting (Hertzberg 1996), which inhibits crack propagation and coalescence, and prevents brittle faulting. Further, twinning produces intragrain weaknesses that promote intragrain deformation and cataclastic flow (Baud et al. 2000). In calcite, this results in a ductile response (Baud et aL 2000), whereas in dolomite crystal plastic process are not operative at low temperatures, and the rock continues to deform by brittle faulting to higher porosities than in calcite, promoting shear localization and inhibiting the onset of cataclastic flow. The role of porosity on the onset of intragranular deformation is highly texturally dependent. Dolomite deforms in the brittle faulting regime at much higher porosities than calcite, where crystal plastic processes inhibit crack growth and propagation. Further, pore size and geometry dictate the degree to which stresses are intensified in the vicinity of the pore, and the concentration of pores dictates whether these localized regions of deformation can interact. The ability of the damaged zones around pores to interact will likely dictate the overall mechanical and microstructural response of the rock to deformation.

Conclusion The transition from brittle faulting to cataclastic flow in dolomite is a complex process that is intricately related to the textural properties of the rock. In dolomite with porosities below c. 7%, this transition is strongly linked to the ability of grains to rotate in relation to one another, an ability that can be influenced both by grain boundary texture and confining pressure. Intragrain flaws, of which cleavage is the most notable in dolomite, and lobate grain boundaries are key parameters that may promote the transition from brittle faulting and transgranular microcracking to extensive intragranular deformation and cataclastic flow. When grains contain well-developed intragranu1at cleavage and lobate grain boundaries, cleavage planes become more conducive to failure than grain boundaries.

65

In dolomite with porosities greater than c. 7%, inelastic pore collapse controls the transition from brittle faulting to cataclastic flow. At relatively low strains, this deformation is manifested as local intragrain deformation and comminution of grains in the vicinity of pores coupled with microcracks that connect and step across pores. This transition occurs at porosities of c. 7% in dolomite, unlike calcite where the transition has been observed in rocks with porosities as low as 3%, a difference that is strongly related to the textural properties of the rocks, as well as to the mechanical properties of the mineral. While porosity and grain size are important textural properties in controlling the transition from brittle faulting to cataclastic flow, so are grain boundary textures and the presence of flaws or inclusions within the grains. All of these must be considered in an examination of the controls of the mode of brittle deformation in dolomite if the spectrum of naturally occurring textures is to be investigated. Financial support for this work was provided by an NSERC operating grant to L. A. Kennedy, a Collaborative Research and Development grant to L. A. Kennedy, and a B. P. research grant to L. A. Kennedy, as well as by an NSERC PGS-A scholarship to N. J. Austin, and a U. B. C. University Graduate Fellowship to N. J. Austin. The sample materials were kindly provided by Imasco Minerals, Mighty White Dolomite, and Graymont Dolime. We thank Hans de Bresser and Ernest Rutter for their constructive reviews, as well as Steve Quane for his input and assistance with the development of the triaxial rock press used in this study.

References ANTONELLINI, M. & MOLLENA, P. N. 2000. A natural analog for a fractured and faulted reservoir in dolomite: Triassic Sella Group, northern Italy. AAPG Bulletin, 84(3), 3 1 4 - 3 4 4 . ASHBY, M. F. & SAMM~S, C. G. 1990. The damage

mechanics of brittle solids in compression. Pure Applied Geophysics, 133(3), 489-521. BARBER, D. J., HEAD, H. C. & WENK, H. R. 1981.

Deformation of dolomite single crystals from 20800 degrees C. Physics and Chemistry of Minerals, 7, 271-286. BATES, R. L. & JACKSON, J. A. (eds). 1984. Dictionary of Geological Terms. American Geological Institute, New York. BAUD, P., SCHUBNEL,A. & WONG, T. F. 2000. Dilatancy, compaction, and failure mode in Solnhofen limestone. Journal of Geophysical Research, 105(B8), 19 289-19 303. BRACE, W. F. 1978. Volume changes during fracture and frictional sliding. Pure and Applied Geophysics, 116, 603-614. COLPRON, M., PRICE, R. A,, ARCHIBALD, D. A. & CARMICHAEL, D. M. 1996. Middle Jurassic

66

N.J. AUSTIN ET AL.

exhumation along the western flank of the Selkirk fan structure: Thermobarometric and thermochronomertic constraints from the Illecillewaet synclinofium, southeastern British Columbia. GSA Bulletin, 108(11), 1372-1392. FREDRICH, J. T., EVANS, B. & WONG, T. F. 1989. Micromechanics of the brittle to plastic transition in Carrara marble. Journal of Geophysical Research, 94(B4), 4129-4145. FREDRICH, J. T., EVANS, B. & WONG, T. F. 1990. Effect of grain size on brittle and semibrittle strength: implications for micromechanical modeling of failure in compression. Journal of Geophysical Research, 95(B7), 10 907- l0 920. GOWD, T. N. & RUMMEL, F. 1980. Effect of confining pressure on fracture behavior of porous rock. Int. J. Rock Mech. Min. Sci., 17, 225-229. GRIFFITH, A. A. 1921. The phenomena of rupture and flow in solids. Philosophical Transactions of the Royal Society of London, 221, 163-198. HANDIN, J. & FAIRBAIRN, H. W. 1956. Experimental deformation of Hasmark dolomite. Geological Society of America Bulletin, 66, 1257-1274. HANDIN, J., FRIEDMAN,M., LOGAN, J. M., PATTISON, L. J. & SWOLFS, H. S. 1972. Experimental folding of rocks under confining pressure; buckling of single-layer rock beams. In: Flow and Fracture of Rocks. Geophysical Monograph, American Geophysical Union, 16, 1-28. HANDIN, J., RUSSELL, J. E. & CARTER, N. L. 1986. Experimental deformation of rock salt. In: HOBBS, B. E. & HEARD, H. C. (eds) Mineral

and Rock Deformation; Laboratory Studies; The Paterson Volume. Geophysical Monograph, The American Geophysical Union, 36, 117-160. HATZOR, Y. H., ZUR, A. & MIMRAN, Y. 1997. Microstructure effects on microcracking and brittle failure of dolomites. Tectonophysics, 281, 141-161. HERTZBERG, R. W. 1996. Deformation and Fracture Mechanics of Engineering Materials. John Wiley & Sons Inc., New York. HIGGS, D. V. & HANDIN, J. 1959. Experimental deformation of dolomite single crystals. Geological Society of America Bulletin, 70, 245-278. HIRTH, G. & TULLIS, J. 1989. The effects of pressure and porosity on the micromechanics of the brittle ductile transition. Journal of Geophysical Research, 94, 17 825-17 838. HIRTH, G. & TULLIS, J. 1994. The brittle-plastic transition in experimentally deformed quartz aggregates. Journal of Geophysical Research, 99(B6), 11 731-11 747. HORII, H. & NEMAT-NASSER, S. 1985. Compressioninduced microcrack growth in brittle solids: axial splitting and shear failure. Journal of Geophysical Research, 90(B4), 3105-3125. HUGMAN, R. H. H. & FRIEDMAN, M. 1979. Effects of texture and composition of mechanical behavior of experimentally deformed carbonate rocks. AAPG Bulletin, 63(9), 1478-1489.

LOGAN, J. M. 1987. Porosity and the brittle-ductile transition in sedimentary rocks. In: Physics and

Chemistry of Porous Media 11, AlP Conf. Proc., 154, 229-242. MENENDEZ, B., ZHU, W. & WONG, T. F. 1996. Micromechanics of brittle faulting and cataclastic flow in Berea sandstone. Journal of Structural Geology, 18(1), 1-16. NEMAT-NASSER, S. & HORII, H. 1982. Compressioninduced nonplanar crack extension with application to splitting, exfoliation, and rockburst. Journal of Geophysical Research, 87(B8), 6805-6821. OLSSON, W. A. 1974. Grain size dependence of yield stress in marble. Journal of Geophysical Research, 79(32), 4859-4862. PAPAMICHOS, E., VARDOULAKIS, I. & OUADFEL, H. 1993. Permeability reduction due to grain crushing around a perforation. International Journal of Rock Mechanics and Mineral Science, 30(7), 12231229. RAUDSEPP, M., PANI, E. & DIPPLE, G. M. 1999. Measuring mineral abundance in skarn. I. The Rietveld method using X-ray powder diffraction data. The Canadian Mineralogist, 37, 1-15. RUTTER, E.H. & HADIZADEH, J. 1991. On the influence of porosity on the low-temperature brittleductile transition in siliclastic rocks. Journal of Structural Geology, 13, 609-614. SHIMAMOTO, T. 1977. Effects of Fault Gauge on the

Frictional Properties of Rocks: An Experimental Study. PhD thesis, Texas A&M. TULLIS, J. & YUND, R. A. 1987. Transition from cataclastic flow to dislocation creep of feldspar: Mechanisms and microstructures. Geology, 15, 606-609. TULLlS, J. & YUND, R. A. 1992. The brittle ductile transition in feldspar aggregates: An experimental study. In: EVANS, B. & WONG, T. F. (eds) Fault

Mechanics and Transport Properties of Rocks. Academic, San Diego, CA, 89-117. TURNER, F. J., GRIGGS, D. T., HEARD, H. & WEISS, L. W. 1954. Plastic deformation of dolomite rock at 380 degrees C. American Journal of Science, 252, 477-488. WONG, T. F., DAVID, C. & ZHU, W. 1997. The transition from brittle faulting to cataclastic flow in porous sandstone: Mechanical deformation. Journal of Geophysical Research, 102(B2), 3009-3025. ZHANG, J., WONG, T. F. & DAVIS, D. M. 1990. Micromechanics of pressure-induced grain crushing in porous rocks. Journal of Geophysical Research, 95(B 1), 341-352. ZHU, W. & WONG, T. F. 1996. Permeability reduction in a dilating rock: Network modeling of damage and tortuosity. Geophysical Research Letters, 23(22), 3099- 3102. ZHU, W. & WONG, T. F. 1997. The transition from brittle faulting to cataclastic flow: Permeability evolution. Journal of Geophysical Research, 102(B2), 3027-3041.

Crack-seal patterns: records of uncorrelated stress release variations in crustal rocks FRAN(~OIS R E N A R D 1'2, M U R I E L ANDRI~ANI 1, A N N E - M A R I E B O U L L I E R 1 & PIERRE LABAUME 3

1LGIT, Universit~ Joseph Fourier, BP 53, 38041 Grenoble, France (e-mail: francois, renard@ Igit. obs. ujf-grenoble.fr) 2physics of Geological Processes, Institute of Physics, postboks 1048 Blindern, 0316 Oslo, Norway 3Dynamique de la lithosphkre (UMR5573), Universit~ MontpeIlier II, place E. Bataillon, 34095 Montpellier Cedex 5, France Abstract: Statistical properties of crack-seal veins are investigated with a view to assessing stress release fluctuations in crustal rocks. Crack-seal patterns correspond to sets of successive parallel fractures that are assumed to have propagated by a subcritical crack mechanism in the presence of a reactive fluid. They represent a time-sequence record of an aseismic and anelastic process of rock deformation. The statistical characteristics of several crack-seal patterns containing several hundreds of successive cracks have been studied. Samples were collected in three different areas, gold-bearing quartz veins from Abitibi in Canada, serpentine veins from the San Andreas system in California and calcite veins from the Apennine Mountains in Italy. Digitized pictures acquired from thin sections allow accurate measurement of crack-seal growth increments. All the samples show the same statistical behaviour regardless of their geological origin. The crack-seal statistical properties are described by an exponential distribution with a characteristic length scale and do not show any spatial correlation. They differ from other fracture patterns, such as earthquake data, which exhibit power-law correlations (Gutenberg-Richter relationship). Crack-seal series represent a natural fossil record of stress release variations (less than 50 bars) in the crust that show a characteristic length scale, associated with the resistance of rock to effective tension, and no correlation in time.

Introduction Syntectonic calcite or quartz veins with a c r a c k seal structure are common in rocks at low metamorphic grades and high fluid pressure (Beach 1977; Ramsay & Huber 1983; Passchier & Trouw 1995). The formation of these veins can be explained by a growth mechanism involving many repeated small increments, the c r a c k seal process (Ramsay 1980). The overall pattern is the result of a sequence of crack increments, followed by periods of precipitation in the open cracks (Fig. 1). Veins attributed to the c r a c k seal mechanism are commonly considered as evidence of episodic crack opening, driven by oscillations in fluid pressure or bulk stress (Ramsay 1980; Cox 1991; Fisher & Brantley 1992; Petit et al. 1999). It is common to observe several hundred successive c r a c k seals, which therefore represent a fossil record of local elastic stress releases in the crust.

However, to the authors' knowledge, no quantitative studies have been conducted on the level of stress variations and their correlations in time. Other fracturing processes in the crust exhibit well-defined correlations, the most famous being the power-law correlations widely found in large sets of fractures (Bonnet et al. 2001). Indicators of crack-seal processes are regularly spaced bands of small inclusions (typically small minerals, pieces of wall rock, or fluid inclusions), aligned parallel to the vein walls. Inclusion trails at high angles to the walls are better indicators of the opening direction than fibrous crystals. Opening per crack event is generally in the order of several micrometres to several tens of micrometres (Cox & Etheridge 1983; Cox 1987; Williams & Urai 1989; Xu 1997). The cracking event corresponds to the opening of a narrow fluid-filled crack along the vein

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. DeformationMechanisms, Rheology and Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 67-79. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

68

F. RENARD ET AL.

Fig. 1~ (a) Sketch of a polished section of fibrous gold-bearing quartz vein, Abitibi, Canada, showing decimetre-scale organization of fibrousribbons, separated by horizontal discontinuities, due to earthquake events (after Firdaous, 1995). Crack-seal patterns analysed in this study are located in the grey area. (b) Microscopic view of crack-seal structure. Quartz layers (light) alternate with tourmaline (dark). Pattern results from cracking events, followed by precipitation of minerals at the walls of open crack. (e) Diagram of crack-seal process. One tourmaline layer (dark) and one quartz layer (light) result in a cracking event, followed by precipitation.

margin, whereas, during the sealing period, precipitates fill the crack again (van der Pluijm 1984; Beutner & Diegel 1985; Ellis 1986; Gaviglio 1986; Cox 1987; Ramsay & Huber 1983; Labaume et al. 1991; Davison 1995; Fisher et al. 1995; Toriumi & Hara 1995). The precipitating crystals usually have a fibrous, elongate or blocky shape (Bons 2000). Urai et al. (1991) proposed a kinematic numerical model for the formation of fibrous

morphologies by a crack-seal mechanism, and predicted that, depending on the boundary conditions, fibres may or may not track the opening trajectory of the crack. In such a kinematic model, the displacement-tracking conditions in crack-seal veins are of two kinds. First, the vein wall should have a rough morphology and, secondly, the growth rate of fibres should be sufficiently fast to fill the open space before the next cracking event (Urai et al. 1991; Hilgers 2000).

STATISTICAL ANALYSES OF C R A C K - S E A L PATTERNS

Computer models of crack-seal fibrous veins have focused on studying the kinematics of vein-precipitation by varying different parameters, such as mineral morphology (e.g. prismatic growth), crack width, roughness, opening frequency and opening trajectory (Hilgers 2000). The nutrients filling the vein can be transported by diffusion from the wall rock or by advection into the flow through the fracture network (Taber 1916, 1918; Boullier & Robert 1992; Fisher & Brantley 1992; B o n s & Jessell 1997). Precipitation in the vein can be due to variations in fluid pressure (e.g. associated with fracturing) or to some nucleation process where new crystals can only grow on the fracture walls. Wiltschko & Morse (2001) proposed that cracking events could be due to the force of crystallization of the growing crystals, inducing subcritical crack propagation through stress concentration at the crack tip. The crack-seal mechanism has been invoked to sustain the concept of fluid valving, an idea derived from considerations of fluid behaviour within and surrounding faults and shear zones, particularly associated with the seismic cycle (Sibson et al. 1988). As crack-seal structures record a time sequence of geological events, they represent a unique natural record of episodic events of cracking, whereby time periods are directly imprinted in the rock. In this contribution, the statistical properties of crack-seal patterns in various rocks are studied and are related to stress release variations and their correlations in time. The aim is to

69

characterize the apparent regularity of crackseal increments and their spatial correlation. This provides new constraints on local states of stress in the crust in various geological settings.

Geological setting and description of the crack-seal samples Three kinds of crack-seal samples were collected and analysed (see Table 1): three quartz veins in diorite from Abitibi, Canada, two serpentine veins from California, and one calcite vein from the Apennines, Italy. It is shown below that, even if they come from various geological areas, these three kinds of samples have similar statistical distributions and spatial correlation properties.

Gold quartz-bearing veins, Val d'Or, AbitibL Canada The structural setting of the gold-quartz veins at Val d' Or (Abitibi) has been described by Robert (1990) and Robert et al. (1995). The Val d'Or district is located in the Archean Abitibi greenstone belt, in which several shear zones of different scales were active during the late stages of a N - S shortening event and contain the quartz-tourmaline-carbonate-pyrite (QTC) veins considered in this paper (Robert 1990; Robert et al. 1995). The QTC vein system at the Sigma deposit occurs in deformed andesites, porphyritic diorites, and feldspar porphyry dikes

Table 1. Top, origin and mineralogy of crack-seal veins. Middle, statistical properties of crack-seal samples:

number of crack-seal events in the sample, thickness of vein, mean and standard deviation of thickness of individual crack-seal events in each vein, minimal thickness of individual cracks. Bottom, stress release variations recorded by crack-seals, assuming elastic properties.

Reference sample

S42abc

S42cdeg

S421hijm

Location

Country Wall rock Vein filling minerals

Abitibi Abitibi Abitibi Canada Canada Canada Diorite Diorite Diorite Quartz + Tourmaline

Santa Ynez Santa Ynez CA, USA CA, USA Serpentinized Peridotite Chrysotile

Apennines Italy Limestone Calcite

No. of successive cracks Vein thickness (~m) Mean crack thickness (~m) Std. crack thickness (~,m) Min, crack thickness (~m)

229 10 100 44.4 48.7 12.0

436 16300 37.6 35.5 8.1

1106 31900 28.9 30.8 7.2

88 350 4.1 2.1 1.3

251 570 2.3 0.8 0.8

379 56372 148.7 86.9 21.8

E (kbar) a W (X 10 -6 m) b

500 44

500 38

500 29

300 4

300 2

150 149

c (m)C wE/4c (bar)

0.5-3 11-66

0.5-3 10-60

0.5-3 7-42

0.02 15

0.02 7.5

0.1 56

aWall-rock Young's modulus. bAverage crack thickness (see Fig. 5). CCrack length data represent maximum values of the actual crack lengths.

SY3P

SY32

X45FXC

70

F. RENARD ET AL.

(Robert & Brown 1986). The Sigma deposit combines subvertical shear zone-hosted fault veins and subhorizontal extensional veins, which display mutually cross-cutting relationships, cyclic growth and deformation textures and earthquake evidence (Boullier & Robert 1992). A seismic fault-valve model, involving fluctuations in fluid pressure, has been proposed to explain such cyclic sequences (Sibson et aI. 1988). The crack-seal vein studied here is a 60 cm thick subhorizontal tabular extensional vein, extending into intact rocks away from the associated fault veins in the Sigma deposit (Fig. la). It consists of several centimetre-scale wall-parallel ribbons of quartz and tourmaline representing growth layers and attesting to incremental development during successive earthquakes. These ribbons are themselves made of numerous crack-seal increments ranging between 7 and 200 ~m in thickness (Figs lb and 2). The opening direction was vertical, which implies a fluid pressure slightly higher than the lithostatic load during the formation of the vein (Boullier & Robert 1992). Quartz infilling has an elongate blocky structure in which crystals have a vertical long axis. The width of the crystals is 1 ~m on average. Delicate undeformed tourmaline needles grow in rosettes on the inclusion bands on both sides of a crack-seal increment. We infer that they have crystallized in fluid-filled open-space cavities (Fig. 2b). The crack-seal structure is comparable to that described by Henderson et al. (1990); that is, inclusion bands are locally interrupted. Such interruptions of the crack-seal sequence also occur at a smaller scale of a few bands within the quartz crystals (Figs 1 and 2a). However, crack-seal sequences remain identical on both sides of the interruption, regardless of its scale, and may be followed over the entire width of large thin sections (7 cm) and from one sample to another. The delicate geometry of crack- seal inclusion bands and the presence of undeformed tiny tourmaline fibres, that are not broken, indicate that the crack propagated slowly. Robert & Boullier (1994) have proposed that these crack-seal textures are induced by a steady-state mechanism of subcritical crack propagation in order to explain their constant shape mimicking the irregularities of the crack walls. The cracks followed the weak boundary between the host rock and the previous sealed crack. Like other authors (Robert & Boullier 1994; Cox 1995; Wiltshko & Morse 2001), we infer that crack propagation occurred in subcritical conditions. Based on experimental work, crack velocity may be estimated in the range 10 -6 to 10 -3 m s -1, depending on stress,

Fig. 2. (a) Crack-seal sequence (crossed polars) of

dark layers of tourmaline and light layers of quartz in gold-bearing quartz veins, Abitibi, Canada (sample S42abc). (b) Tourmalinefibres have nucleated from the crack wall-formingrosettes. This indicates that nucleation occurred far from equilibriumand that crystals grew in an open space. temperature, and fluid composition (Atkinson 1984; Gurguen & Palciauskas 1994).

Serpentine veins, San Andreas system, California Serpentine samples have been collected north of Santa Barbara (California), along the Santa Ynez fault, in the Blue Canyon. From late Jurassic to late Cenozoic times, this area has undergone subduction of the Pacific plate (Atwater 1989), resulted in an accretionary complex along the coast. In California, this so-called Franciscan subduction complex is a mrlange of different sedimentary and ultramafic rocks, showing various degrees of deformation and metamorphism under HP-LT (e.g. Ernst 1971; Page 1981). Bodies of serpentinite are distributed throughout the mrlange (Page 1972). They are derived from

STATISTICAL ANALYSES OF CRACK-SEAL PATTERNS mantle peridotites that are partially to totally hydrated. Outside the highly deformed zones, massive serpentinites are preserved and show a network of hydrothermal veins, comparable to those described in oceanic ridge serpentinites (Dilek et al. 1997; Stamoudi 2002). Extensional veins have a banded internal fabric (Fig. 3). They cross-cut at various orientations the serpentinized peridotite matrix that shows the typical mesh texture of replacement of olivine and pyroxene. They have been attributed to incremental stress release during progressive unroofing of serpentinites in the M A R K area (Dilek et al. 1997). Like other veins of the same type, the banded veins considered here have irregular edges that are strictly parallel to each other. These criteria are similar to those invoked for incremental opening and filling by a c r a c k seal mechanism (Ramsay 1980). Minimum incremental opening is around 1 ~m and the maximum observed is less than 5 ~m. Nanometer-size tubes of chrysotile (tubular serpentine mineral) have a preferred orientation perpendicular to the vein margins and fill each crack antitaxially. Thus, fibres do not directly track the displacement path, as observed in quartz veins by Cox (1987). However, the opening direction can be clearly followed thanks to a few inclusions of wall-rock and the undulose extinction linking edge irregularities across the vein. Bands are almost separated by a nanometric free space and tubes do not show any trace of deformation. In these veins, there is no evidence of cataclastic deformation; this

Fig. 3. Crack-seal sequence (crossed polars) in a serpentine vein, Santa Ynez fault, San Andreas system, California (sample SY3P). The individual cracks are filled with antiaxial serpentine fibres. The delicate geometry of the fibres has been conserved throughout time. The only disturbances are due to trapped impurities. This is indicative of a slow cracking mechanism, such as subcritical crack growth.

71

could indicate that the mechanism of vein opening was a subcritical crack propagation process. Calcite veins, Apennines, Italy Calcite crack-seal veins were collected in the Northern Apennines (NW Italy), a belt formed by stacked allochthonous units thrust toward the northeast from the Oligocene to the Recent (Elter 1973). In the northern part of the Bobbio window, where the sample comes from, several superficial units belonging to the Sub-Ligurian and Ligurian accretionary complexes have been thrust over lower Miocene turbidites of the Tuscan foredeep (Labaume & Rio 1994). The allochthon consists of mainly deep-sea marlstone, claystone and sandstone units, affected by low-angle faults. The tectonic pile was later folded antiformally, probably above a thrust. Calcite veins are abundant in the marlstones of the allochthon, where they are commonly associated with scaly deformation marking the lowangle shear zones (Labaume et al. 1991). The most common veins are tabular bodies 1 to 5 cm in thickness and several metres to tens of metres in length. They consist of several superposed ram-thick calcite sheets with striated surfaces. Internal microstructures show that each calcite sheet is a shear vein. Calcite has crystallized in a large number of rhomb-shaped cavities, which are releasing oversteps between the two bounding shear surfaces (Fig. 4). The acute angle between the shear surfaces and the initial rupture, responsible for each overstep, gives an unambiguous indication of shear sense. This angle suggests a mode I rupture, that is, subparallel to the local direction of greatest principal stress. The direction of subsequent opening was

Fig. 4. Crack-seal sequence (ordinary light) in a calcite vein, Apennines, Italy (sample X45FXC). Calcite crystals appear in light colours whereas dark marlstone particles detached from walls of vein line the limit of each crack-seal surface. White arrows indicate direction and sense of shear; black arrows indicate shear surfaces.

72

F. RENARD ET AL.

controlled by movement along the shear surfaces. This is a similar geological setting to what Davison (1995) has described in limestones from Kilve, UK. Releasing oversteps form sequences of many narrow calcite veinlets, separated by thin bands of encasing sediment (matrix bands) and bounded by closely spaced ( < 5 mm) shear surfaces. Individual veinlet width varies from 0.025 to 1 mm, the most common values being around 0.1 ram. For each veinlet, the shaperatio (length of initial rupture to opening width) is usually well in excess of 10. Matrix bands are mostly extremely thin, relative to adjacent veins, and locally discontinuous. The veinlets formed sequentially, each releasing overstep opening only after the crystallization of calcite in the previously formed neighbouring overstep. The presence of the matrix bands shows that successive ruptures occurred preferentially in the sediment rather than at the calcite-sediment interface. This may be due to incomplete cementation of the very fine-grained sediment, allowing the pores to be invaded by calcite close to the veinlet walls. During the following rupture this thin band of sediment impregnated by calcite remained glued to the vein-fill, thus forming the matrix band that separates neighbouring veins. The crack-seal vein sequences are commonly offset a few tenths of millimetres by microtransform faults subparallel with the shear surfaces (Fig. 4). Some of these transform faults appear to have initiated at heterogeneities (e.g. fossil debris or a large siliciclastic grain), which locally hardened the sediment, thus deviating the ruptures.

Statistical analysis of crack-seal pattern Polished thin sections observed under an optical microscope with crossed polars were used to extract digital pictures with a high-resolution CCD camera. The position of each crack on a line was measured parallel to the direction of vein opening. The vein wall is set as the origin, and successive cracks are characterized by their distance from the origin. The distance between crack n and crack n + 1 is the aperture of crack n. This was calculated directly on the digitized images. The resolution of the measurement is around 0.5 micron (pixel size of the camera) for the highest optical magnification. Using the crack position data and spatial derivative it was possible to construct cumulative displacement curves, distribution histograms, and crack thickness frequency plots. In addition, the signal was binarized by replacing the information on crack positions, pixel by pixel: the

pixel is replaced by a '0' if it does not contain a crack and by ' 1' if it contains a crack boundary. This binarization process enables the relative spatial location of each crack to be reproduced. The binarized data set, determined directly from the previous one, is used to calculate Fourier transforms of the signal and therefore to study the spatial correlation properties of the pattern.

Crack thickness histograms: a characteristic length scale For each crack-seal vein, the position of each crack is plotted as a function of vein opening (Fig. 5, left). This plot shows how successive crack thicknesses stack spatially and characterizes the series of opening events. The cumulative opening of the vein is found to be almost a linear function of crack position. On average, the number of cracks on a segment parallel to the opening vector depends only on the length of the segment. The derivative of this curve gives the thickness of each successive crack as a function of its position in the vein (Fig. 5, middle). The histograms of thicknesses (right column of Fig. 5) indicate that there is a characteristic length scale and a tail of thicker cracks. These thickness- frequency distributions are shown in the following to follow an exponential relationship.

Distribution function: exponential distribution of crack thicknesses The crack thickness data (Fig. 5, middle) are sorted according to decreasing thickness. We use the so-called rank-ordering technique (Zipf 1949), whereby variables are arranged in descending order wl > w2 > ... > w,, and crack thickness w is plotted as a function of the

Fig. 5. Analysis of crack-seal data for six different crack-seal patterns (see Table 1). For each sample, the data analyses indicate similar results. Left column: cumulative displacement (i.e. distribution function) of the cracks, perpendicularly to the vein. Crack labelling starts near the vein wall (no. 1) and stops towards the centre of the vein (last crack-seal increment number). Middle column: thickness of cracks as a function of position in the series. This is the derivative of the curve in the left. Right column: histograms of crack thicknesses for the various samples. All distributions have tails for large thicknesses and a characteristic peak indicating a preferential crack thickness. The mean crack thickness is also shown.

STATISTICAL ANALYSES OF CRACK-SEAL PATTERNS

i 1~176176176 1oo0J Abtibs42bj 4ooi ' 600 5000

1oo 200 300 00 crack-sealnumber x 104 2 [ __ Abitibi,S42cde9 I ~o 400

0oi / 0oi /

~200

>~

~

,

-~ u=

0

200 400 crack-sealnumber

100 200 300 crack-sealnumber

~ol[

~o 0 600

~,otl ~Olt

~SantaY.... ~ P l

El0

~ooI~ 0~

0

~

o oo

s

~

~4oo

.~ E

=

0

0

50 100 crack-sealnumber

~"

100 200 300 crack seal number

0 1 2 3 4 5 6 7 8 910 12 crack-sealthickness(micron) l ~_mooo,.~.~m, ....

~4

4o ~ol/

~

~-2o

2 0

0

oi_

~

~, 200

mean:4.1 micron~

o

50 100 crack-sealnumber

c

30.

-~ glo ~-

~

~

0 -' 0 100 200 300 crack-sealthickness(micron)

~20

~5

>9 loo I =

~ 10

~oo b- 15f j ~ ~

300/

2

~ ~Otlll

0 t ............... 0 500 1000 1500 crack-sealnumber

500 1000 1500 crack-sealnumber

. . . . .

0 100 200 300 crack-sealthickness(micron)

,o[i

~ 100

o 400/II J .~

o

200 400 600 crack-sealnumber

,oo

>~

mean:38 micron

g2Orll

~ ~oo

0

~ ~0 [ i i l l l ~ ~ 00 100 200 300 crack-sealthickness(micron) 40[ [ I I

~oo

4 x104

J

men44lmcron .........

_- 200

0

._s g

4~

10

~o 0

100 200 300 crack-sealnumber

0

0 5 10 crack-sealthickness(micron)

X 10 4

6 : f ~ X i

cI 7 6 0 0 r ~

In

E400 200

~=

0~

0

,

200 400 crack-sealnumber

0 ~

0

200 400 crack-sealnumber

~, 0 200 400 600 crack-sealthickness(micron)

?3

F. R E N A R D ET AL.

74 ..ICI 't-

.s

3

.o "t-

~ 2,5

-o 2.5

-~ 2.5

3

t:t

o

o

2

2

~- 1.5

-1.5

C

1.5

:6 6~ _o 0.5 0

100 200 300 S42abc crack seal No.

0

200 400 600 S42cdeg crack seal No.

1

1.4

..Q "C

'i=.

.-~ 1.2 "0

~

~o u'}

/'Y" J " ' -

1

f',

0,8

(-.n J

0.6

I

0.8 c.-

~0.5

0.4

0

500 1000 1500 S421hijm crack seal No.

3

~

2.5

~

2

g

1.5

0.6 0.2

= 0.4

r

m

r

'5 02 6~ o 0

I

0

._

50 100 SY3P crack seal No.

I

0 0.2 [

0

100 200 3OO SY32 crack seal No.

1 0

200 400 X45FXC crack seal No.

Fig. 6. Distribution of crack thicknesses. The data of the middle column in Figure 5 are sorted by decreasing thickness (rank-ordering technique). In a log-linear plot, the sorted data (black line) plot close to a straight line (red) over 80% of the signal (except for large and small cracks), indicating an exponential distribution. Residuals (blue line) have a mean value close to zero and no specific trend.

rank n. Represented in a log-linear diagram, the rank ordering plot is qualified by a straight line on more than 80% of each data set (Fig. 6). This plot represents the distribution function of the cracks in a direction parallel to the opening vector and this describes the cumulative distribution of the probability density function. The distribution is mostly linear in a log-linear diagram, indicating that the data can be described by an exponential function (Otnes & Enochson 1978) where the number n of cracks of thickness w obeys the relationship

equal to Wo. Table 1 shows that for all crackseal patterns, the mean crack thickness and the standard deviation are reasonably close. This measurement and the linear trends in Figure 6 indicate that the crack-seal data can be described by an exponential distribution, the deviation from an ideal distribution being small at small and large crack thicknesses. The large number of events (up to 1106 successive cracks for sample S421hijm) ensures that the exponential distribution is representative.

Fourier analysis: a b s e n c e of spatial correlation The characteristic scale Wo represents a physical length or arises through the dynamics of the fracturing process (Bonnet et al. 2001). Note that for the serpentine samples, the exponential distribution is not smooth, because the average crack aperture is close to the resolution of the optical measurement process. For an exponential distribution, the mean and the standard deviation of the data set should be

To study the spatial correlation of successive cracks, the Fourier transform of the binarized crack-seal patterns was calculated. In a l o g log plot, the Fourier spectrum shows a flat linear relationship (Fig. 7). This indicates that the crack-seal data behave like a random series, having no preferential correlation in the signal: the variations of crack thickness are not correlated spatially between each other.

STATISTICAL ANALYSES OF CRACK- SEAL PATTERNS 6

__u -t

gE 4 E O

75

6

I

......

- - initial data .. 9 randomly permuted dab

g

\1 ~.,.

4

initial data randomly permuted dab

E

E O

2

c-

- \

C

2

9

'-

0 0

O

u- 2 O

S42cdeg

S42abc

4

4

3

2

-

1

2

5

4

log of frequency (micron-1)

3

2

1

0

log of frequency (micron -1)

~ 6 ii[ initial data I .... randomly permuted dab

'5

~

E 4 E b

3

.

.

nitial data .

.

BI

E

~2 t-

"t-

"

0

~ o

-2

4

3

2

JSY3 1 3

1

log of frequency (micron-1) 5

\

I --:

2

1

0

log of frequency (micron-1) 6

initial data '

- -

daa

E E 3

initial data randomly permuted dab

E

2 t~

2

1 0 ii

ii . . . . . . . . . . . . . . . .

5Y32

or) O

--2

3

45FXC 2

1

0

log of frequency (micron -1)

2

5

4

3

2

1

0

log of frequency (micron -1)

Fig. 7. Fast Fourier Transform (FFT) analysis of crack-seal patterns. In log-log plot, the overall flat shape of the Fourier spectrum indicates that noise dominates the data and that no specific organization can be detected. The raw data (full lines) and the randomly permutated data (dotted lines) show a similar horizontal linear trend.

This interpretation was confirmed by performing Fourier transforms on randomly permuted sets of data. For each c r a c k - s e a l sequence, the positions of the cracks along the profile parallel to the vein opening direction were randomly permuted. This transformation gives series with the same m e a n and variance as the original ones and for which the initial order (spatial correlation), if any, is destroyed (dotted lines on Fig. 7). Performing Fourier transforms on these new sets

does not change the flatness of the Fourier spectrum. The question as to whether binarization of the original patterns taken with the camera could induce a bias on the high frequencies in the Fourier analysis needs to be addressed. To test this effect, Fourier transforms were also performed on the raw data, taken directly from the camera. Here again, the Fourier spectrum showed a flat linear relationship.

76

F. RENARD E T AL.

Another test, not shown here, is the autocorrelation function of the crack-seal sequence. For all samples, the autocorrelation function drops to zero immediately (as for randomly permutated data sets) thereby confirming that the crack-seal patterns are uncorrelated in space. Therefore, the successions of crack-seal thicknesses do not show any visible spatial organization. This is a memoryless process (the thickness of crack number n does not depend on the thickness of crack n - 1), which is also a property of exponential distributions.

Discussion To summarize the statistical analysis, cracks have an exponentially decaying thickness distribution (Fig. 5). The mean and standard deviation crack distribution values are close and the distribution is almost linear in a log-linear diagram (Fig. 6). Moreover the cracking events are not spatially correlated (Fig. 7). The crack-thickness distribution has all the properties of an exponential distribution. The exponential probability density function has been widely used to accurately describe the failure times in various systems (Fisz 1963; Otnes & Enochson 1978). It characterizes systems containing a large number of elements (transistors, elastic fibres, and so on) which break with a constant failure rate probability. As an analogy to these systems, the exponential distribution of cracks indicates that they are characterized by a single failure rate, which can be controlled, for example, by the external loading conditions. Geological systems such as mid-oceanic ridges (Cowie et al. 1993) also exhibit an exponential distribution. Based on numerical simulations on a spring network model, Spyropoulos et al. (2002) have concluded that exponential distributions of fractures in rocks are indicative of an increase in brittle strain in a regime dominated by the heterogeneity in the system. Moreover, Gupta & Scholz (2000) have observed a transition from a power-law to an exponential distribution of faults as brittle deformation increases in the Afar area. The three types of crack-seal veins show the same trend of statistical properties, regardless of the geological context, the type of rock and the vein kinematics (extensional or shear vein). The linear trend in the cumulative displacement plot (Fig. 5, left) is interpreted as being the fact that the far-field boundary loading conditions and the physico-chemical parameters during the crack-seal deformation did not vary notably; see also Davison (1995) for similar data. There

is neither acceleration nor slow-down in average vein growth rate. As a result, the variations observed in crack thickness should reflect disturbances around a steady-state fracturing process in a constant environment. This steady-state fracturing process is determined both by far-field stress conditions and by the local petrography and stress field in the wallrock. For example, the grain size in the wallrock could control the average crack thickness through nucleation processes or local variations in the wall-rock yield strength. Petrographic evidence also confirms that the P-T conditions did not change during vein formation. Intergrowths of quartz and tourmaline form the crack-seal pattern in the Abitibi quartz veins (Fig. 2). These two crystals have grown from hydrothermal conditions and mineral paragenesis and the fluid inclusions trapped in the successive layers indicate that the thermodynamic conditions (lithostatic pressure, temperature) have not changed much during the whole vein opening history (Firdaous 1995; Robert et al. 1995). Such observations are similar for the three types of veins. The crackseal sequence represents cycles of breakage events and periods of quiescence during which the filling minerals have grown from a fluid. During that time, stress and temperature were kept more or less constant. Consequently, crack-seal structures have recorded a sequence of events occurring in a rock domain with almost invariant geological conditions. In this respect, they have recorded small variations around a steady state. A crack-seal vein accumulates elastic strain due to tectonic loading until the stress on it reaches the yield strength. At this stage a new crack forms. The presence of a well-defined peak in the crackseal thickness distribution indicates that the yield stress varies around a mean value. The uncorrelated distribution of crack-seals and the exponential distribution of crack thicknesses can be interpreted by two effects. On one side, they can reflect spatial heterogeneities in the rock as a crack opens. On the other side, they can reflect time-effective stress and/or fluid pressure variations at the vein boundary. In the first assumption, the exponential distribution could be due to fluctuations in the crack length or in the local rock strength. The average crack length was measured in the field or directly on the samples (Table 1) and remains almost constant for all the cracks in each vein. In addition, each crack follows exactly the same parallel path as the previous one. Moreover there is no evidence that the crack thickness is related to the local mineralogy of the wall-rock.

STATISTICAL ANALYSES OF CRACK-SEAL PATTERNS For the Abitibi samples, the wall-rock contains centimetre-long quartz phenocrystals along which several tens of cracks have developed. Locally, these cracks also show an exponential distribution, even if there is no variation in the wall rock. Therefore there is no structural evidence in the sample that the heterogeneity content varies spatially and that there were local variations in the strength of the vein. In the second assumption, the crack thickness varies, whereas all the textural or geometric parameters of the veins remain unchanged. As the crack lengths are almost constant, the crack thickness variations reflect disturbances in the crack thickness/crack length ratio. Rock mechanics indicate that this ratio is directly related to the state of stress around a fracture and/or at a fracture tip. Given that, after the cracking event, the crack remains open, several authors have proposed that high fluid pressures open the cracks (Sibson et al. 1988; Boullier & Robert 1992). For a pressurized, elastically opened crack, crack width is proportional to the driving pressure (Pollard & Segall 1987). With this assumption, the overpressure or, necessary to open a crack of length c with an aperture w is

wE o- = - 4c

(2)

(Gu6guen & Palciauskas 1994). This corresponds to the elastic stress released when a new crack has opened. Crack surface areas are similar for all the individual cracks; only the crack thickness varies. In this case, the thickness can be used to calculate typical stress variations that allowed crack opening assuming that the crack length is known. What crack-seal structures record in fact are variations in elastic stress release during the vein formation. The values for c given in Table 1 represent a maximum length for the cracks as observed in the outcrops or directly on samples. Given these lengths and the typical elastic modulus of the wall-rocks, the variations in crack thickness indicate minimum effective driving pressure variations ranging from 10 to 50 bars. These variations are small compared to the main stresses. However, they are sufficient to explain the crack-seal patterns.

Conclusions Crack-seal veins, recording several hundred successive crack events, have been analysed. All data sets show similar statistical properties, irrespective of the various geological settings and vein kinematics: the exponential crack

77

distributions have many similarities with a random data set that has a well-defined characteristic length. In this way, the distributions are different from other fracture patterns, which obey power-laws (as in earthquakes) or other types of correlations. Crack-seal structures provide a fossil record of non-correlated crack sequences that have a well-defined average characteristic size: they have recorded noisy stress release variations in the crust. This project has been supported by the CNRS through an Action Thdmatique Innovante. We would like to thank C. Pequegnat and D. Tisserand for technical help. We thank S. Cox and P. Meakin for fruitful discussions and F. Robert for providing the Abitibi crack-seal samples. We acknowledge reviews by M. Jessell, X. Zhang and P. Cobbold.

References ATKINSON, B. K. 1984. Subcritical crack-propagation in geological materials. Journal of Geophysical Research, 89, 4077-4114. ATWATER, T. 1989. Plate tectonic history of the northeast Pacific and western North America. In: WINTERER, E. L., HUSSONG, D. M. & DECKER, R. W. (eds) The Eastern Paci[ic Ocean and Hawaii. Geology of North America, Geological Society of America, Boulder, CO, 21-71. BEACH, A. 1977. Vein arrays, hydraulic fractures and pressure solution structures in a deformed flysch sequence, S.W. England. Tectonophysics, 40, 201-225. BEUTNER, E. C. & DmGEL, F. A. 1985. Determination of fold kinematics fi'om syntectonic fibers in pressure shadows, Martinsburg slate, New Jersey. American Journal of Science, 285, 16-50. BONNET, E., BOUR, O., ODLING, N. E., DAVY, P., MAIN, l., COWIE, P. & BERKOWITZ, B. 2001. Scaling of fracture systems in geological media. Reviews of Geophysics, 39, 347-383. BONS, P. & JESSELL, M. 1997. Experimental simulation of the formation of fibrous veins by localised dissolution-precipitation creep. Mineralogical Magazine, 61, 53-63. BONS, P. 2000. The formation of veins and their microstructures. Journal of the Virtual Explorer, 2, 12. BOULHER, A. M. & ROBERT, F. 1992. Palaeoseismic events recorded in Archaean gold-quartz vein networks, Val d'Or, Abitibi, Quebec, Canada. Journal of Structural Geology, 14, 161 - 179. COWIE, P. C., SCHOLZ, C. H., EDWARDS, M. & MALINVERNO, A. 1993. Fault strain and seismic coupling on mid-oceanic ridges. Journal of Geophysical Research, 98, 17 911-17 920. Cox, S. F. & ETHERtDGE, M. A. 1983. Crack-seal fibre growth mechanisms and their significance in the development of oriented layer silicate microstructures. Tectonophysics, 92, 147-170.

78

F. RENARD ETAL.

Cox, S. F. 1987. Antitaxial crack-seal vein microstructures and their relationship to displacement paths. Journal of Structural Geology, 9, 779-787. Cox, S. F. 1991. Geometry and internal structures of mesothermal vein systems: implications for hydrodynamics and ore genesis during deformation, ln: Structural Geology in Mining and Exploration. University of Western Australia, Extended Abstracts, 25, 47-53. Cox, S. F. 1995. Faulting processes at high fluid pressures: an example of fault valve behavior from the Wattle Gully Fault, Victoria, Australia. Journal of Geophysical Research, 100, 12 841-12 859. DAVISON, I. 1995. Fault slip evolution determined from crack-seal veins in pull-apart and their implications for general slip models. Journal of Structural Geology, 7, 1025-1034. DILEK, Y., COULTON, A. & HURST, S. D. 1997. Serpentinization and hydrothermal veining in peridotites at site 920 in the Mark area. In: KARSON, J. A., CANNAT, M., MILLER, D. J. & ELTHON, D. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 153, 35-59. ELLIS, M. 1986. The determination of progressive deformation histories from antitaxial syntectonic crystal fibers. Journal of Structural Geology, 8, 701-709. ELTER, P. 1973. Lineamenti tettonici ed evolutivi dell Appennino. Acad. Naz. Lincei, Roma, 183, 97-109. ERNST, W. G. 1971. Do mineral parageneses reflect unusually high pressure conditions of Franciscan metamorphism? American Journal of Science, 271, 81-108. FIRDAOUS, K. 1995. Etude des fluides dans une zone sismog~nique fossile: les gisements aurif~res m~sothermaux archdens de Val d'or, Abitibi, Quibec. University Thesis, Institut National Polytechnique de Lorraine. FISHER, D. M. & BRANTLEY, S. L. 1992. Models of quartz overgrowth and vein formation: deformation and episodic fluid flow in an ancient subduction zone. Journal of Geophysical Research, 97, 20 043- 20 061. FISHER, D. M., BRANTLEY, S. L., EVERETT, M. & DzvoYIK, J. 1995. Cyclic fluid flow through a regionally extensive fracture network within the Kodiak accretionary prism. Journal of Geophysical Research, 100, 12 881-12 894. FISZ, M. 1963. Probability Theory and Mathematical Statistics. Wiley Publications in Statistics, Wiley, New York. GAVIGLIO, P. 1986. Crack-seal mechanism in a limestone: a factor of deformation in strike-slip faulting. Tectonophysics, 131, 247-255. GUEGUEN, Y. & PALCIAUSKAS,V. 1994. Introduction to the Physics of Rocks. Princeton University Press, Princeton, NJ. GUPTA, A. & SCHOLZ, C.H. 2000. Brittle strain regime transition in the Afar depression: implications for fault growth and sea-floor spreading. Geology, 28, 1087-1090. HENDERSON, J. R., HENDERSON, M. N. & WRIGHT, T. O. 1990. Water-sill hypothesis for the origin of

certain veins in the Meguma Group, Nova Scotia, Canada. Geology, 18, 654-657. HILGERS, C. 2000. Vein growth in fractures: experimental, numerical and real rock studies. Unpublished PhD dissertation thesis, University of Aachen, Germany. LABAUME, P., BERTY, C. & LAURENT, P. 1991. Syndiagenetic evolution of shear structures in superficial nappes: an example from the Northern Apennines, NW Italy. Journal of Structural Geology, 13, 385-398. LABAUME, P. & RIO, D, 1994. Relationships between the subLigurian allochthon and the Tuscan foredeep turbidites in the Bobbio window (NW Apennines). Memorie della Societd Geologica Italiana, 48, 309-315. OTNES, R. K. & ENOCHSON, L. 1978. Applied Time Series Analysis. Wiley, New York. PAGE, B. M. 1972. Oceanic crust and mantle fragment in subduction complex near San Luis Obispo, California. Geological Society of America Bulletin, 83, 957-972. PAGE, B. M. 1981. The Southern Coast Range. In: ERNST, G. W. (ed) The Geotectonic Development of California. Prentice-Hall, Englewood Cliffs, New Jersey, 330-417. PASSCHIER, C. & TROUW, R. 1995. Microtectonics. Springer, Berlin. PETIT, J.-P., WIBBERLEY, C. A. J. & RUIZ, G. 1999. Crack-seal, slip: a new fault valve mechanism? Journal of Structural Geology, 21, 1199-1207. POLLARD, D. D. & SEGALL, P. 1987. Theoretical displacement and stresses near fractures in rock with applications to faults, joints, veins, dykes and solution surfaces. In: ATKINSON B. K. (ed) Fracture Mechanics of Rock. Academic Press, London, 277-350. RAMSAY, J. G. 1980. The crack-seal mechanism of rock deformation. Nature, 284, 135-139. RAMSAY,J. G. & HUBER, M. I. 1983. The Techniques of Modern Structural Geology. Volume 2: Folds and Fractures. Academic Press, London. ROBERT, F. 1990. Structural setting and control of gold-quartz veins of the Val d'Or area, southeastern Abitibi Subprovince. In: Ho, S. E., ROBERT, F. & GROVES, D. I. (eds) Gold and Base Metal Mineralization in the Abitibi Subprovince, Canada, with Emphasis on the Quebec Segment. University of Western Australia Publication, 24, 164-209. ROBERT, F. & BOULLIER, A. M. 1994. Mesothermal gold-quartz veins and earthquakes. In: HICKMAN, S. H., SIBSON, R. H. & BRUHN, R. L. (eds) The Mechanical Involvement of Fluids in Faulting. U.S. Geological Survey, Open-file report 94-228, 18-30. ROBERT, F., BOULLIER,A. M. & FIRDAOUS,K. 1995. Gold-quartz veins in metamorphic terranes and their bearing on the role of fluids in faulting. Journal of Geophysical Research, 100, 12 86112 879. ROBERT, F. & BROWN, A. C. 1986. Archean goldbearing quartz veins at the Sigma Mine, Abitibi greenstone belt, Quebec. Part I: Geologic relations

STATISTICAL ANALYSES OF CRACK-SEAL PATTERNS and formation of the vein systems. Economic Geology, 81, 578-592. SIBSON, R. H., RO~Er~T, F. & POULSEN, K. H. 1988. High angle reverse faults, fluid-pressure cycling and mesothermal gold quartz deposits. Geology, 16, 551-555. SPYROPOLOUS, C., SCHOLZ, C. H. & SHAW, B. E. 2002. Transition regimes for growing crack populations. Physical Review E, 65, 056105-1056105-10. STAMOUD1, C. 2002. Processus de serpentinisation des

piridotites de Hess-Deep et de la zone de MARK." Approches chimiques et miniralogiques. PhD thesis, University Paris VI, France. TABER, S. 1916. The growth of crystals under external pressure. American Journal of Science, XLI (4th series, No. 246), 532-556. TABER, S. 1918. The origin of veinlets in the Silurian and Devonian strata of central New York. Journal of Geology, 6, 56-73. TORIUMI, M. & HARA, E. 1995. Crack geometries and deformation by the crack-seal mechanism in the

79

Sambagawa metamorphic belt. Tectonophysics, 245, 249- 261. URAI, J. L., WILLIAMS, P. F. & van ROERMUND, H. L. M. 1991. Kinematics of crystal growth in syntectonic fibrous veins. Journal of Structural Geology, 13, 823-836. VAN DER PLUIJM,B. A. 1984. An unusual 'crack-seal' vein geometry. Journal of Structural Geology, 6, 593 -597. WILLIAMS, P. F. • URAI, J. L. 1989. Curved vein fibres: an alternative explanation. Tectonophysics, 158, 311-333. WILTSCHKO, D. V. & MO~SE, J. W. 2001. Crystallization pressure versus 'crack seal' as the mechanism for banded veins. Geology, 29, 79-82. Xu, G. 1997. Fluid inclusions in crack-seal veins at Dugaid River, Mount Isa Inlier: implications for paleostress states and deformation conditions during orogenesis. Journal of Structural Geology, 19, 1359-1368. ZIPF, G. K. 1949. Human Behavior and the Principle of Least-Effort. Addison-Wesley, Cambridge.

Single-contact pressure solution creep on calcite monocrystals SERGEY ZUBTSOV x'2, FRAN(~OIS R E N A R D 1'3, J E A N - P I E R R E G R A T I E R l, D A G K. D Y S T H E 3 & V L A D I M I R T R A S K I N E 2

1LGIT-CNRS-Observatoire, Universitd Joseph Fourier, BP 53, 38041 Grenoble cedex, France (e-mail: Francois.Renard@ lgit. obs. ujf-grenoble.fr) 2Colloid Chemistry Group, Chemistry Department, Moscow State University, Moscow 119899, Russia 3physics of Geological Processes, University of Oslo, Postboks 1048 Blindern, 0315 Oslo, Norway

Abstract: Pressure solution creep rates and interface structures have been measured by two methods on calcite single crystals. In the first kind of experiments, calcite monocrystals were indented at 40 ~ for six weeks using ceramic indenters under stresses in the 50-200 MPa range in a saturated solution of calcite and in a calcite-saturated aqueous solution of NH4C1. The deformation (depth of the hole below the indenter) is measured ex situ at the end of the experimenc In the second type of experiment, calcite monocrystals were indented by spherical glass indenters for 200 hours under stresses in the 0-100 MPa range at room temperature in a saturated aqueous solution of calcite. The displacement of the indenter was continuously recorded using a specially constructed differential dilatometer. The experiments conducted in a calcite-saturated aqueous solution of NH4CI show an enhanced indentation rate owing to the fairly high solubility of calcite in this solution. In contrast, the experiments conducted in a calcite-saturated aqueous solution show moderate indentation rate and the dry control experiments did not show any measurable deformation. The rate of calcite indentation is found to be inversely proportional to the indenter diameter, thus indicating that the process is diffusion-controlled. The microcracks in the dissolution region under the indenter dramatically enhance the rate of calcite indentation by a significant reduction of the distance of solute transport in the trapped fluid phase. This result indicates that care should be taken in extrapolating the kinetic data of pressure solution creep from one mineral to another.

Pressure solution creep (PSC), coupled with cataclastic and crystal plastic deformation, plays a significant part in sedimentary and fault rock deformation in the upper crust (Ramsay 1967). This is a mechanism of a slow aseismic compaction and porosity reduction whereby stressenhanced chemical potential at the contact sites of mineral grains results in local supersaturation of the adjacent fluid, diffusion of the solutes out of the high concentration areas, and precipitation of the material on the grain faces with smaller chemical potential (Weyl 1959; Paterson 1973; Rutter 1976). According to this definition, a distinction can be made between reactionkinetics-controlled PSC, where the rate of deformation is limited by the dissolution or precipitation of the material, and diffusion-controlled PSC, where this rate is limited by solute diffusion along the grain contacts (Raj 1982).

The main difficulty in quantifying which parameters control PSC rates is that this very slow mechanism of deformation is difficult to reproduce experimentally. Moreover, this process largely depends on the shape of the solid that evolves with time because of deformation. This is the main reason why indenter experiments have been developed, where an inert indenter in contact with a crystal and its solution is shown to pressure dissolve the crystal (Gratier 1993; Dysthe et al. 2003). As the shape of the indenter does not change with time, the deformation is easier to quantify. Moreover, the deformation laws for this simple geometry are well known and allow separating the different parameters controlling the rate of deformation. By this means one expects to deliver creep laws for relevant geological conditions and to quantify both finite strain and rates of PSC deformation.

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and

Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 81-95. 0305-8719/05/$15.00 (~ The Geological Society of London 2005.

S. ZUBTSOV ET AL.

82

The displacement rate of a piston (~/) indenting a mineral by PSC is deduced from theoretical laws (Raj 1982; Rutter 1983) for the two types of PSC. For the diffusion-controlled case, the relationship is Ocw

~oc d 2

(1)

where c is the mineral solubility at stressed state, w the thickness of the water film under the indenter, D the diffusivity of rate limiting ion, and d the indenter diameter. Here, C:

CoeAIx/RT

(2)

where Co is the mineral solubility at zero stress, R the gas constant, T the temperature, and A/z the difference in chemical potential between the dissolution and precipitation zones (proportional to the applied normal stress O-n). At low A/~ (if o'n < 3 0 MPa; Rutter 1976) an approximation may be used to obtain the relation derived by Gratier (1993):

Dco w A l~ RTd 2

(3)

For the kinetic-controlled case, the relationship is

yoc k(C - C~ \ co /

(4)

or expressed more simply:

;y oc kc

(5)

Here ( c - Co)/Co is the saturation state of the solution and k the reaction rate constant of the interfacial reaction (dissolution or precipitation). For precise calculations it should be borne in mind that this constant is temperaturedependent. Using relation (2) at crn < 30MPa, the relation derived by Gratier (1993) for mineral indentation by kinetic-controlled PSC is obtained, namely:

kcoAl~ RT

(6)

However, relation (4) is based on the approximation that the rate of interfacial reaction is proportional to the saturation state of the adjacent solution. This is not always true. In the general case, this rate (and thus the rate of mineral indentation by kinetic-controlled

PSC) is in power relation with the saturation state (Morse & Arvidson 2002):

:y oc k(C - C------~~n \ c0 /

(7)

At low applied stress, a relationship similar to that of de Meer and Spiers (1999) can be derived from equation (7):

kco( A tz) n RT

~/o~ -

(8)

Here n is the empirical reaction order, which is different for different minerals. Moreover, it is different for different solution compositions and usually different for dissolution and precipitation processes (Morse & Arvidson 2002). For calcite in pure water, n was found to be equal to 1 for dissolution (Plummer et al. 1978) and 2 for precipitation (Reddy & Wang 1980). Generally, all of the preceding means that the relationship between the rate of mineral indentation by PSC and the value of the difference in chemical potential between the dissolution and precipitation zones (i.e. the value of applied normal stress On) is sophisticated and depends on many factors. Care should be taken when using the chemical potential to determine whether kinetics or diffusion-controlled PSC takes place in an experiment. On the other hand, equations (1) and (5) show that, if the process of crystal indentation is diffusion-controlled, the rate of indenter displacement is inversely proportional to the square of the indenter diameter, and in the case of kinetic control the indenter diameter has no influence on this rate. The relationship between indenter diameter and indentation rate seems to be a more reliable parameter on which to base conclusions concerning the controlling mechanism of PSC. Two types of laboratory experiment have been conducted previously to investigate PSC. In the single contact experiments, the crystal surface is indented by a piston or another crystal (Tada & Siever 1986; Gratier 1993; Dysthe et al. 2002). This type of experiment is conducted in order to observe the deformation process precisely controlling the applied normal stress, the contact surface area and the displacement of the indenter. These experiments represent a useful tool for validating existing theoretical pressure solution models (Dysthe et al. 2003). Unfortunately, for low-solubility minerals, such as calcite or quartz, the rate of single contact deformation is too slow and the laboratory experiment must last several months or the displacement recording system must be extremely sensitive. This is why a second experimental

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS approach is used for these minerals. It consists of powder compaction, that is, multiplying the number of grain contacts to increase the bulk deformation rate (de Meer & Spiers 1997; den Brok et al. 1999; Zang et al. 2002). However, the interpretation of the results obtained is ambiguous as other deformation processes (grain sliding, subcritical crack growth) can occur in addition to PSC (Zang et al. 2002). Contrary to the significant number of theoretical and experimental studies of cataclastic and plastic calcite deformation, the PSC of calcite has not previously been studied in great detail in spite of its importance in natural rock deformation processes. The main reason for this situation is the low solubility of calcite, which makes it difficult to investigate its PSC in laboratory conditions. However, a limited number of pressure solution compaction experiments using calcite powder have been performed. Baker et al. (1980) studied compaction of deep-sea carbonate Iceland spar and reagent-grade calcite powder at 3 5 - 1 0 0 M P a and 2 2 - 1 8 0 ~ The duration of experiments varied from 21 to 240 hours. After 240 hours of compaction, a porosity reduction of 9.9% was obtained. Unfortunately, there is no clear evidence in these experiments that cataclastic and crystal plastic deformation do not make a significant contribution to the porosity reduction value obtained. Recently, Zang et al. (2002) compacted finegrained calcite powders at ambient temperature. The compaction experiments were performed in drained conditions with a saturated calcite solution, applying a uniaxial effective stress of 1 - 4 MPa. All samples were initially precompacted at 8 MPa to minimize the contribution of grain sliding and reorientation to subsequent wet deformation. The absence of brittle deformation during the experiment was proven by the absence of recorded acoustic emissions. However, a mechanism of subcritical crack growth was considered to be a relevant deformation process. Dry experiments and experiments with inert fluid did not show any significant deformation. Deformation of the wet samples reached 1.1% maximum strain after 20 days of compaction. The compaction rate decreased with increasing strain, increasing grain size and decreasing applied stress. On the basis of these data, it was concluded that reaction kinetic-controlled pressure solution creep and possibly subcritical crack growth were the dominant deformation mechanisms. To our knowledge, no single contact PSC experiments have been performed to date because of the very slow rate of calcite single crystal deformation. There are at least three different approaches

83

described below to resolve this problem: modifying the fluid chemistry, inducing microcracks, and increasing the sensitivity of the deformation measurement device. Effect o f f l u i d chemistry

The nature of the fluid is an important parameter for the rate of PSC. Above all, it is obvious that the rate of deformation depends on the solvent capacity of the fluid; that is, it is proportional to the solubility of the mineral in the fluid. For example, Pharr & Ashby (1983) studied the rate of pressure solution deformation of precompressed potassium chloride and sucrose powders in methanol/water solutions. They found that the extent of deformation enhancement is a linear function of the solubility of the solid in the pore fluid. Gratier & Guiguet (1986) used a solution of NaOH to increase the deformation rate of quartz, starting from the idea that alkalinity enhances its solubility as well as its dissolution and growth kinetics. Moreover, the chemistry of the fluid can influence the rate of pressure solution deformation by phenomena such as adsorption of fluid components and complex formation. Skvortsova et al. (1994) investigated the effect of adding small quantities of dimethylaniline (DMA) on the rate of deformation of halite by PSC. They found that DMA forms complexes with salt and adsorbs on the salt crystal surfaces, resulting in a change from a diffusion-controlled to a reaction-controlled mechanism. This transition is accompanied by a sharp decrease in deformation rate. In addition, Zang et al. (2002) found that calcite powder PSC slows down significantly if a small quantity of MgC12.6H20 is added to the pore fluid. The reason for this effect is the adsorption of magnesium ions on the calcite crystal surfaces, which inhibits further calcite crystallization. The rate of calcite deformation by PSC can also be enhanced by changing the fluid chemistry. As calcite is a weak base, its solubility is increased in acid solutions. In particular, in weak acid solutions, a dynamic chemical equilibrium takes place so that calcite cannot only be dissolved in the solutions, but also precipitates if the solution is supersaturated. This is not the case for a total dissolution reaction, as for example when calcite dissolves in HC1. A wellknown natural example of this dynamic equilibrium is the dissolution of calcite in a solution saturated in carbon dioxide: CaCO3 q- C O 2 "q-H20 = Ca(HCO3) 2

(9)

The rate of mineral deformation by PSC is proportional to its solubility; consequently the rate

84

S. ZUBTSOV ET AL.

of calcite deformation by PSC should be enhanced in a weak acid solution saturated by calcite. Effect of microcracks Gratier et al. (1999) have shown in experiments of halite indentation by diffusion-controlled PSC that if the applied stress is high enough to form microcracks on the mineral surface under the indenter, the rate of deformation enhances dramatically. The reason for this effect is that the kinetics of diffusion along a trapped fluid phase under the indenter could be one order of magnitude less than in bulk fluid (Dysthe et al. 2002; Alcantar et al. 2003). If there are no microcracks under the indenter, the indentation rate is inversely proportional to the square of the indenter diameter. However, in the presence of microcracks, faster diffusion out of the contact may occur in the free fluid that fills the microcracks. Consequently, in this case the rate of mineral indentation is not limited by the diameter of the indenter, but by the distance between two adjacent microcracks. It should be noted that calcite is a brittle mineral and the formation of microcracks on its surface under the indenter seems plausible. Increase in recording sensitivity Another possibility for carrying out a singlecontact PSC experiment with calcite is to increase the sensitivity of the displacement recording system. Dysthe et al. (2002, 2003) developed a capacitance dilatometer especially designed for single-contact pressure solution experiments. The sensitivity of this device is about 10 - 9 m. Halite monocrystals were tested using the dilatometer and showed a creep rate of about 2 x 10 - 6 to 10 - 8 m/day under an applied stress of 4 MPa (Dysthe et al. 2003). In this contribution, two types of singlecontact PSC experiments with calcite monocrystals are described. The three abovementioned ways of activating PSC on calcite crystals were successfully used in this study. The interpretation of the results obtained provides new constraints for understanding the mechanism of PSC on calcite deformation.

Experimental methods P S C indenter experiments" with ex situ deformation measurement With the simple indenter technique described in detail by Gratier (1993), calcite monocrystals

were indented at 40 ~ using ceramic indenters under stresses in the 50-200 MPa range for six weeks in a saturated solution of calcite and in a calcite-saturated aqueous solution of NH4C1 (Fig. la). Natural calcite crystals (from Sweetwater Mine, Reynolds Country, Missouri, USA, delivered by Ward's Natural Science Establishment Inc., Rochester, NY, USA) were cleaved and then polished (1200 to 4000 mesh grinding paper and DP-mol cloth with 1 micron diamond suspension, all from Struers A/S, Rodovre, Denmark). The surfaces were cleaned with water. The aqueous solutions were checked to ensure that they wetted the crystals when the experiments were started. For some experiments without water, the calcite crystals were dried with nitrogen. For each experiment, a different crystal was used and the indenter was always applied on the same crystallographic face. The saturated solution of calcite was prepared by adding several grams of CaCO3 powder to 100 ml of distilled water boiled for 15 minutes before solution preparation in order to eliminate any traces of CO2, and heated to 60 ~ shaking once per day for several days. The aqueous calcite-saturated solution of NH4C1 was prepared by adding 10 g of CaCO3 to 100 ml of a 5% solution of NH4C1 in distilled water and heating to 60 ~ shaking once per day for several days. Before starting the experiments, the solutions were kept at 40 ~ (temperature of the experiments) for two days. Each deformation cell consists of a Plexiglas tube with a diameter of several centimetres, and a Plexiglas bottom. The ceramic indenters, 300 txm wide and of various lengths (see Table 1), are mounted under a free-moving Teflon | piston and put in contact with calcite monocrystals immersed in a saturated solution. Dead weights placed on the piston maintain a constant stress on the samples. The value of the dead weight is calculated for each indenter in order to impose the desired stress (Table 1). Liquid paraffin is added on the top of each cell to prevent evaporation of the fluid. The cells are placed in a furnace at constant temperature. After six weeks, the depths of the resulting holes are measured. Experimental conditions (stress, length of indenter, nature of fluid) for different deformation cells, together with the depth of the holes obtained, are presented in Table 1. The indenter lengths are different for each indenter (from 300 to 740 txm), but the diffusion distance of solutes under the indenter is always the same because of the same indenter widths (300 txm).

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS

1

2

aluminium block \

85

3

:lead a~au

Peltier elements

brass block

9

+/ iOmm

, I

"

isolation

J

_ . . . . . . s . . . . . . . . . . . . . . . . . . . . . . . . . . . I t .. . . . . . . . . . . . . . . . . . . . s !

,

"

oil ,

dry sample

'

upper ~ ~ ~ i n d e n ' e r i / / electr ...................... ............[ t / / ~ c e n t e r e l e c t r o d e i , . i .~i}

'\

soluiion

" wet sample " t ' - ' - ' ~ sapphire window

~ / f - - ~ \ \ ~\

--(

O

~-~

~'/ glass spheres glued on the indenter

Fig. 1. (a) PSC indenter experiments with ex situ deformation measurement. First stage: equilibrium between calcite crystal and its saturated solution. Second stage: the indenter is placed on the mineral surface and vertical stress is applied through a dead weight. Third stage: slow creep of the indenter downwards into the crystal. The deformation is analysed ex situ by measuring the depth of the imprint. (b) High-resolution PSC experiments with continuous displacement recording through a double capacitance device (Dysthe et al. 2002, 2003). The technique allows displacement to be measured accurately with a long-term resolution of 2 nm (see text).

86

S. ZUBTSOV ET AL.

Table 1. Experimental conditions and results for PSC indenter experiments with ex situ deformation measurement. All the experiments last for six weeks at constant temperature (40 ~ All the indenters have a constant width (300 lzm) and varying lengths

Table 2. Experimental conditions for high-resolution PSC experiments with continuous deformation recording

No.

1

1 3 4 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 26 27 28 29 30 31 32 33 34 35 37 38 40

Length Dead Stress of the weight (MPa) indenter (kg) (~m) 300 380 360 310 320 500 340 300 400 400 300 320 300 300 300 360 740 380 300 340 340 320 380 300 400 300 300 300 300 360 340 390 360 340 300

1.20 1.52 1.44 1.24 1.28 2.00 0.77 0.68 0.90 0.90 0.68 0.72 0.68 0.68 0.30 0.36 0.19 0.10 1.20 1.36 1.36 1.28 1.52 1.20 1.60 0.68 0.68 0.68 0.68 0.81 0.77 0.88 0.36 0.34 0.08

200 200 200 200 200 200 150 150 150 150 150 150 150 150 100 100 50 50 200 200 200 200 200 200 200 150 150 150 150 150 150 150 100 100 50

Solution Depth chemistry of hole (txm)

No.

2 3

Calcite Calcite Calcite Calcite Calcite Dry Dry Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite Calcite NH4C1 NH4C1 NH4C1 NH4C1 NH4CI NH4CI Dry Dry NH4C1 NH4C1 NH4CI NH4C1 NH4C1 NH4C1 NH4C1 NH4C1 NH4C1

0 0 0 5 5 0 0 19 0 5 5 18 5 0 0 0 5 5 52 65 29 48 108 46 0 0 15 4O 12 23 5 42 25 16 5

High-resolution P S C experiments with continuous deformation recording Experiments were conducted using the differential dilatometer described in Dysthe et al. (2002, 2003), which allows displacement to be measured accurately with a long-term resolution of 2 nm (Fig. l b). The device consists of two nominally identical capacitances, indenters and samples, one wet and one dry. This symmetrical design serves to eliminate all effects except those caused by differences on the two sides (presence/absence of water). The temperature is controlled by thermostatically regulated water

Fluid chemistry

Dead weight (g)

Saturated aqueous solution of calcite Saturated aqueous solution of calcite Dry (dried with nitrogen, immersed in hexadecane)

0; 80; 167 167 167

flowing through brass tubes spiraling down the external part of the instrument. The electronic circuit is a specialized lock-in amplifier built in-house and based on that of Jones (1988). About 2 0 - 3 0 glass spheres with diameters ranging from 250 to 300 txm were glued along the edge of a cylindrical glass support. This kind of indenter geometry provides good mechanical stability during the experiment. As the experiment goes on, the spheres penetrate into the 1.5 mm diameter calcite crystal. The contact surface area between the indenter and the crystal gradually increases, resulting in a progressive decrease in stress, The samples and solutions were prepared as described for the long-term calcite monocrystal indentation experiments. Table 2 shows the experimental conditions for these experiments. All experiments were run at stabilized room temperature. The temperature was controlled with circulating water from a thermally controlled bath, with fluctuations of about _+0.1 ~ that are larger than that described in Dysthe et al. (2003) for pressure solution on halite. Because the solubility of calcite is lower than halite and not very dependent on temperature (0.07% change with 0.1 ~ temperature change) the calcite experiments are not as sensitive to temperature fluctuations as the halite experiments. From time to time the electrical noise in the environment of the instrument causes short fluctuations in the measured indentation. These fluctuations make quantitative interpretation difficult for some parts of the experimental curves.

Results P S C indenter experiments with ex situ deformation measurement Figure 2 shows a typical hole obtained after calcite monocrystal indentation. The depth of holes was measured after the experiments using

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS

87

120

100

S.

80

60

,.'"

40

50

Fig. 2. Typical hole in a calcite monocrystal obtained after six weeks of indentation. The shape of the hole, close to a parallelepiped, represents that of the indenter. The depth of each hole is measured under microscope at the end of the experiment. a microscope with a resolution of three microns. The results are presented in Table 1. This table shows that none of the dry experiments (experiments 8, 9, 28, and 29) show any trace of deformation even under a maximum stress of 200 MPa. The experiments conducted with saturated calcite solution also usually do not show any deformation or show imprints of more than 5 ~m depth after six weeks (experiments 1, 3, 4, 6, 7, 11-13, 15-20). Depths greater than 5 ~m have been observed in two experiments only (10 and 14), which can be interpreted as brittle deformation. Therefore, it may be concluded that, under these conditions, there is no significant calcite deformation if saturated calcite solution is used as a fluid. On the other hand, the experiments conducted in a calcite-saturated aqueous solution of NHaC1 show significant deformation (Table 1, experiments 21-24, 26, 27, 30-35, 37, 38, 40). Moreover, these data indicate a correlation between the depth of the hole and the applied stress (Fig. 3).

High-resolution PSC experiments with continuous displacement recording Experimental data. Figure 4 shows the results of the first PSC experiment with the double capacitance dilatometer. During this experiment the load was varied as shown in Figure 4a. In this experiment the vertical displacement of the indenter is recorded continuously. This displacement corresponds to the indentation of some glass spheres in the calcite crystal. The

100 150 Stress, MPa

200

250

Fig. 3. Relationships between depth of hole and applied stress in the long-term indentation experiments conducted in a calcite-saturated aqueous solution of NHaC1. The straight line represents a guide for the eye.

temperature during the experiment is shown in Figure 4b. Figure 4a and b show that differential measurement makes the double capacitance insensitive to temperature fluctuations of the order of 0.1 ~ The disturbances in the indentation curve during the first five hours of the experiment arise from the failure of the temperature control system (see Fig. 4b). Figure 4a shows that there are large short-term fluctuations in the measured indentation. These are caused by electrical noise in the environment of the instrument. However, it is possible to separate the rapid jumps due to noise from the slowly varying capacitance due to indentation by PSC. Table 3 shows the indentation rates obtained by linear regression to the noise-free periods of the indentation curve shown in Figure 4a. The rate varies from 6 to 47 nm/h. It is believed that these variations represent the changes in stress under the indenter and the dissolution of minor asperities of calcite. However, the most important result of this experiment is the possibility of measuring the rate of calcite deformation with a resolution never reached before and to obtain relevant indentation rates. As illustrated in Figure 4c, an increase in dead weight by a factor of 2.1 (from 80 to 167 g) has an immediate effect on the deformation rate, which increases by a factor of 2.7 (from 11 nm/h in the interval 3 8 - 4 6 h to 30 nm/h in the interval 4 6 - 5 4 h ) . Figure 4d also shows that if the dead weight is removed, the indentation process stops completely. The electrical disturbances make it difficult to view this effect clearly throughout the whole stress-free period.

88

S. ZUBTSOV ETAL. 3 E

~2

o l/j

1

0g

167g

67 g

~ o~

._

o

~o

~o

20

40

6o

80

60

80

100 1'20 time, h

1~40

1'60

1'80

200

140

160

180

1~40

160

180~00

45 40 ~5 L. -~ 35

L)

II ~ 26,6

~3o III ~

100

l

E N 25 J L _ J

120

time, h

2 0 ~ 0

60

80

1130 1120 time, h

E =-1.7 E el.E E d9

f

J

~1.~

Q_

~9 1.4

"O

167 g

1.!

8

,~0

42

,;4

~e

,~8

time, h i

go r

g2

54

1'15

116

i

~2.21~

2.2~ E

2.2z

~Q'212z ,_ -(3

2.2

/ 09

1'10

1'11

i

112

1'13

i

114

timo., h

Fig. 4. Experimental results of the first high-resolution pressure solution creep experiment. (a) Indentation curve: the dead weight was changed several times in the course of the experiment as shown on the figure. (b) Temperature during the experiment. (c, d) Zooms on the creep curve.

Nevertheless, the moderately long interval without electrical disturbances and without any movement of the indenter can be distinguished in Figure 4d (Table 3, interval 109.5-113.3 h). The indentation process starts again with the same rate as soon as the dead weight is replaced in its original position (Table 3, intervals 7 9 - 8 9 and 122-128 h). The results of the second experiment are shown in Figure 5a. The only difference from

the first experiment is a new indenter with a different number of spheres touching the crystal surface. Figure 5b shows the temperature of the experiment. The rates of indentation in the noise-free parts of the signal are 48 and 56 n m / h at the beginning and end of the experiment, respectively. Figure 6a shows the results of calcite dry indentation. The indentation curve only shows some electrically induced fluctuations between

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS Table 3. Rates of deformation at different time intervals of the first high-resolution PSC experiment (see Table 2) Time (h)

Indentation rate (nm/h)

Dead weight (g)

23.3 46.6 15.0 27.6 18.3 14.6 0.0 15.4 6.5

80 80 80 167 167 167 0 167 167

9-15 19-24 30-35 48-54 69-79 79-89 109.5-113.3 122-128 169-182

envelope and thus the relative height of the imaged surface at each pixel is determined with a resolution of 3 nm. The horizontal resolution depends on the lens used and with the highest magnification it is at the diffraction limit of about 0.5 pxm. The calcite crystal used in the first experiment was examined using this method. Figure 7a shows an indented part of a crystal surface. This image shows all the spherical imprints of the indenter and the scratches produced by the horizontal movement of the indenter. The depth of the holes is about 0.65 Ixm and that of the scratches is less than 0.30 ~m. These observations show that the change in dead weight during the experiment causes minor displacement of the indenter accompanied by surface scratching. When the spheres are placed under stress again, they restart crystal indentation until a new change in dead weight. This is the reason for the moderate depth of the holes measured by WLI, comparable to the values of deformation obtained in the experiment. Figure 7b presents a typical spherical dissolution imprint and its two-dimensional profile. The depth of the hole is about 1 o,m and its diameter is approximately 50 txm. The hole is surrounded by rings, which are about 200 nm higher than the crystal surface. These rings represent the reprecipitated calcite, Observations also indicate that the holes are produced by the

0 and 200nm, which are 10-15 times less than the deformation values obtained in the wet experiments for the same period of time.

Profiles of indented regions. The sample surfaces with indented regions are imaged by white light interferometer (WLI) microscopy (Wyko 2000 Surface Profiler from Veeco). The profiler is a microscope with a reference arm creating interference fringes with maximum intensity at equal optical path lengths of the imaging beam and reference beam. By vertical movement of the sample and simultaneous image capturing, the interference intensity

12 E

89

F

10

E

8

u9

6

~ 4

._

2 0 0

L

I

~

I

I

20

40

6O

8O

100

26.9

L

140

i

i

i

T

i

26.6

~

26.75

160

--

%

26.85

~

&

i

120

time, h

a

(J .

_

26,7

210 0

~ 40

J_ 60

, 80

i 1 O0

~ 120

140

160

time, h

Fig. 5. Experimental results of the second high-resolution pressure solution experiment. (a) Indentation curve; (b) temperature during the experiment.

S. ZUBTSOV ET AL.

90 0.2 E =.0.15 0.1 ~.. 0.05 ._

0 0

1'0

:~0

I

30

a

I

time, h

40

I

~0

60

26.7 ,o ~- 26.6s ~_ 26,61~

"~

E 26.64

.,.-,

26'620 b

1'0

20

I

I

30

40

50

dO

time, h

Fig. 6. Experimental results of the third (dry) high-resolution pressure solution experiment. (a) Indentation curve; (b) temperature during the experiment.

dissolution of calcite and not by subcritical crack growth and consequent brittle deformation. There is also evidence that some cracks formed under the indenter along the calcite cleavage planes. These cracks are produced by the action of high stress on the brittle mineral.

Stress estimate. For further calculations, the stress applied on the crystal surface in our experiments should be estimated. Given that the dead weight is 167 g, the diameters of the spherical imprints on the crystal surface are about 50 txm, the number of spheres touching the crystal surface is around 10, the stress could be assessed at 100 MPa.

Discussion Dominant deformation mechanism PSC indenter experiments with ex situ deformation measurement. The absence of any imprints on the crystal surface in the dry experiments strongly suggests that there is no crystal plastic deformation of the crystals and that the dominant mechanism is water-assisted deformation. The presence of dissolution imprints suggests that PSC is the main deformation mechanism.

High-resolution PSC experiments with continuous deformation recording. The deformation process stops in the absence of dead weight, that is, when no more stress is applied to the crystal surface. This fact indicates that, in the experiments conducted, deformation is effectively being measured and not electrical fluctuations. The spherical imprints of the indenters on the crystal surface (Fig. 7) indicate that the deformation measured is the deformation of the crystal surface itself and not of parts of the experimental device. The absence of any significant deformation in the dry experiments strongly indicates that the dominant mechanism of deformation is pressure solution. The presence of nearby grown calcite around the indentation imprint (Fig. 7b) also confirms this conclusion. Rate limiting process The rate of calcite monocrystal indentation in the high-resolution PSC experiments is about 1 jxm/ day for a stress of 100 MPa (see Table 2 and the results of the second experiment). On the other hand, the experiments with ex situ deformation measurements did not show any significant deformation after 42 days of indentation, even under a maximum stress of 200 MPa. The difference between these two types of experiment is the length of the diffusion path under the indenter which is six times shorter in the high-resolution

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS

91

Fig. 7. White light interferometry image of the indented surface of a calcite monocrystal from the first high-resolution pressure solution creep experiment. (a) Spherical imprints of the indenter and some scratches produced by the horizontal movement of the indenter on the whole calcite crystal. (b) Zoom on a typical imprint of a sphere on the calcite surface and its cross-sections, which can be used to determine the depth of the imprint and also shows the 200 nm height rims of calcite precipitation near the edge of the indentation (arrows).

PSC experiments than in the experiments with e x s i t u deformation measurement (the diameter of the spherical imprint on the crystal of the high-resolution experiment is 50 ~zm compared to the 300 Ixm width of the indenters in experiments with e x s i t u deformation measurement). Given that the rate of crystal indentation is independent of indenter diameter in the case of kinetic-controlled PSC (equation 4) and inversely proportional to the square of the indenter diameter in the case of diffusion-controlled PSC (equation 3), it could be suggested that diffusion is the limiting step in our experiments. To test this assumption, the rate of crystal indentation in the experiments with e x s i t u deformation measurement should be calculated for the diffusion-controlled case on the basis of the rate

in the high-resolution experiments. Starting from the difference in the length of diffusion paths, it is easy to see that this rate should be about 0.03 Ixm/day multiplied by a factor that represents the difference in the stress in these two experiments. This means that the final deformation after the end of the experiment (42 days) should be no more than several microns. This value is in good agreement with the experimental result - the absence of any measurable deformation after the end of the experiment. Thus, it may be concluded that calcite indentation by PSC under our experimental conditions is limited by the diffusion of solutes under the indenter. The diffusion control of calcite PSC in our experiments is a crucial difference between our

S. ZUBTSOV ET AL.

92

results and those of Zang et al. (2002). The plausible explanation of this difference is the indenter technique used here compared to the powder compaction method. Because of the very slow rate of calcite deformation by PSC, the maximum value of powder compaction obtained was only 1.2%. This means that the contact interface areas between two adjacent grains remain very small and irregular throughout the experiment. Therefore, the diffusion distances in this experiment are always extremely short. This may be the reason for kinetic control operating in the calcite powder compaction experiments. On the other hand, in the indentation experiments, significantly larger diameters of indenter impose a longer diffusion path, and the process is thus diffusion controlled.

Effect of the mineral solubility The experiments with ex situ deformation measurement with a saturated aqueous solution of calcite do not show any significant deformation. On the other hand, in the experiments conducted in a calcite-saturated aqueous solution of NH4C1, deformation of several tens of microns can be observed (Table 1). The reason for this change probably is the effect of ammonium chloride as a weak acid on the solubility of calcite. Solutions of weak acids should increase the rate of calcite deformation by PSC. For example, in an aqueous solution of NH4C1 saturated by calcite there is a dynamic equilibrium: 2CACO3 + 2NH4C1 + 2H20 = Ca(HCO3)2 + CaC12 + 2NHaOH

(10)

The chemical potential of the CaCO3 surface under the indenter is higher than that of the free surface. Therefore, in this region, the equilibrium is displaced to the right-hand side of equation (10) and dissolution of calcite takes place. This additional quantity of dissolved calcite diffuses to the free surface, where the concentration of dissolved calcite is lower, and precipitates there. The solubility of calcite in a 5% solution of NH4C1 is about 0.5 g/1 (Lafaucheux 1974), that is, 30 times higher than in pure water. This means that in this solution the final depth of the holes obtained after the end of the experiment should be 30 times higher than in the case of pure calcite solution. Based on the depth value of several microns in the case of pure calcite solution (see calculations in previous section), the depth of holes in the presence of NH4C1 may be estimated to be several tens of microns. This value is in agreement with the experimental results.

Effect of stress PSC indenter experiments with ex situ deformation measurement. Unfortunately, the scatter of the data (Fig. 3) makes it impossible to obtain a good approximation of this correlation by any simple function. On this figure, a straight line is represented as a guide for the eyes, but a power-law or an exponential function could also fit the data. It may be only noted that enhancement of the applied stress increases the rate of mineral indentation by PSC. High-resolution PSC experiments with continuous deformation recording. Figure 7a shows that the change in dead weight during the experiment causes some minor horizontal displacement of the indenter. When the spheres are under stress again, they restart crystal indentation until a new change in dead weight. During the first minutes of indentation, the contact area between the spheres and the surface is extremely small because of the small irregularities that are present even on a very well polished surface; therefore the spheres apply an excessively high stress on new areas of the crystal surface. As a consequence, a dramatic increase in deformation rate should be detected during the first minutes after the change in dead weight. This effect is clearly seen on the experimental curve when the dead weight changes from 0 to 167 g at t = 144 h (Fig. 4d). On the other hand, this is not the case when the dead weight changes from 80 to 167 g at t = 46 h (Fig. 4c). Maybe, in this second example, the indenter did not move horizontally during the change in dead weight and each sphere stayed in its initial hole. An increase in dead weight by a factor of 2.1 has an immediate effect on the rate of deformation, which increases by a factor of 2.7 (Fig. 4c). We propose that in our experiment, the rate of mineral deformation is almost directly proportional to the applied stress. Note that we only have one piece of data to support this conclusion. Comparison with PSC on halite PSC indenter experiments with ex situ deformation measurement. The experiments with a saturated solution of calcite in 5% solution of NH4C1 show indentation rates of about 1 p~m/ day (Table 1). This value is the same as in the experiments of Gratier (1993), where the rate of halite monocrystal deformation in the same type of experiments with diffusion control was found to be about 1 txm/day at the applied stress of about 50 MPa and with an indenter

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS diameter of 200 txm. Given that the solubility of calcite in an NH4C1 solution is 700 times lower than that of halite, that the rate of calcite indentation should also be slowed by a factor of 2 compared to halite because of the difference in the indenter diameters (300 txm in calcite experiments and 200 Ixm in halite ones), and that this rate should be enhanced several times because of the higher applied stress used in our experiments (the rate of mineral deformation by diffusion-controlled pressure solution is proportional to the applied stress in the case of moderate stress and in exponential proportion to the stress if the stress is high), it can be estimated that calcite deforms two orders of magnitude faster than it should when considering the difference in solubility between calcite and halite.

High-resolution PSC experiments with continuous deformation recording. The observed rate of calcite indentation is of the order of several tens of nm per hour (see Table 3). It should be noted that this rate does not change significantly during the experiment. This is an important difference between our results and those described in Dysthe et at. (2003) for similar experiments with halite single crystals with diffusion control, where the indentation rate varies by several orders of magnitude throughout the experiment: from 2 txm/h at the start, when stress is imposed, to 2 nm/h after several days. However, on most of the calcite indentation curve (from 10 to 100 h) the indentation rate falls in the range between 27 and 180 nm/h (Table 2). These data show that calcite deforms only 5 to 10 times slower than halite, whereas its solubility is 20 000 times lower. For our experiments, the stress may be estimated at 100MPa. This stress is 25 times higher than that used in the halite experiments. This difference enhances the rate of mineral indentation by two orders of magnitude. In addition, the difference in the lengths of diffusion paths (50 Ixm in our experiments compared to 30 Ixm in the halite indentation experiments) slows down this rate by a factor of 2.5. However, even with these explanations, the rate of calcite deformation remains two orders of magnitude higher than it should be if compared with halite when considering the difference in solubility. Comparison between the two types of experiments. If this method is used to compare the results of Gratier (1993) and Dysthe et al. (2002, 2003) for the rate of halite indentation, it would be found that these two rates are of the same order of magnitude provided that they

93

are scaled with respect to the stress and the indenter diameter. On the other hand, for the two types of calcite experiments, the deformation rate is two orders of magnitude higher than it should be when compared with halite considering the difference in solubility. Consequently, there are some additional effects that may be responsible for the fast rates measured on calcite.

Effect of microcracks For an additional explanation, the structure of the crystal surface should be considered. It is known that in diffusion-controlled PSC, the presence of microcracks in the region of dissolution under the indenter dramatically enhances the rate of mineral indentation by a significant reduction in solute transport distances in the trapped fluid phase (Gratz 1991; den Brok 1998; Gratier et al. 1999). The reason for this effect might be related to the substantial difference in transport properties between the trapped fluid phase along the grain contact and the bulk fluid (Rutter 1976). Because of this, the diffusion flux in microcracks (bulk fluid) is greater than along the trapped fluid phase. If the dissolution surface is cut by microcracks, the average diffusion distance in the trapped fluid phase is significantly reduced and the deformation rate is increased dramatically (den Brok 1998). This effect is believed to be critically important in our experiments, as a considerably high stress is applied on a brittle mineral. Cracks seen on the interferometry images of the indenter imprints in the high-resolution PSC experiments provide support for this assumption. The wide scatter of hole depths obtained in the same experimental conditions (Fig. 3) also confirms the microcrack hypothesis. The microcrack structure is unique for each experiment, which is the reason for the significant variation of indentation rates. Precise understanding of this interface structure needs further investigation. Microcracks might be formed by the release of elastic forces when the indenter is left on the crystal. Another effect is the formation of twinning that can be readily achieved in laboratory conditions. Twinning could lead to localized hardening and could promote further fracturation. This is a fast process that potentially happens in the early stages of the deformation process and controls the initial fracture spacing. At this stage, it can only be unequivocally assessed that the calcite interface structure under the indenter is radically different compared to that of halite. This result indicates that great care must be taken in extrapolating the kinetic

94

S. ZUBTSOV ETAL.

data of pressure solution creep obtained from one mineral to another.

Conclusions This article reports on the first experimental study of calcite monocrystal indentations by the pressure-solution creep. 9 The rate of mineral deformation was found to be inversely proportional to the square of the indenter diameter, which means that the process is diffusion-controlled. 9 When comparing the experiments with various calcite solubility (aqueous solution versus solution of NH4CI), the rate of calcite indentation was found to be proportional to its solubility. 9 The rate of calcite deformation was found to be almost proportional to the applied stress. 9 The observed rate of calcite indentation is two orders of magnitude higher than it should be when compared with the halite indentation rate in the same type of experiments, taking into account the difference in solubility of these two minerals. This effect could be explained by the presence of microcracks under the indenter phase. Understanding of this effect needs further studies of the interface structure. Our experimental data suggest that PSC deformation on calcite can be quite fast. This process should occur in nature (basin compaction, fault gouge) with a wide range of rates, depending on stress, fluid composition, and crack density. The presence of microcracks should enhance PSC in active faults after an earthquake as new surfaces are created and new 'indenters' are formed. This could explain the fast fault creep observed after a major earthquake (Donnellan & Lyzenga 1998). The project has been supportedby the ANDRA, the CNRS (Action Thrmatique Innovante), the Norwegian Research Council through the Fluid Rock Interaction Strategic University Program (grant 113354-420), and the Center for Advances at the Academy of Science of Norway. We would like to thank R. Guiguet, D. Tisserand, C. Pequegnat and L. Jenatton for their technical help, and D. Gapais, S. Reddy and B. den Brok for their constructive reviews.

References ALCANTAR, N., ISRAELACHVILI,J. & BOLES, J. 2003. Forces and ionic transport between mica surfaces: Implications for pressure solution. Geochimica and Cosmochimica Acta, 67, 1289-1394.

BAKER, P. A., KASTNER, M., BYERLEE, J. D. & LOCKNER, D. A. 1980. Intergranular pressure solution and hydrothermal recrystallisation of carbonate sediments - an experimental study. Marine Geology, 38, 185-203. DE MEER, S. & SPIERS, C. J. 1997. Uniaxial compaction creep of wet gypsum aggregates. Journal of Geophysical Research, 102, 875-891. DE MEER, S. & SPIERS, C. J. 1999. On mechanisms and kinetics of creep by intergranular pressure solution. In: JAMVEIT, B. & MEAKIN, P. (eds) Growth and Dissolution in Geo-Systems. Kluwer Academic Publishers, Dodrecht, 345-366. DEN BROK, S. W. J. 1998. Effect of microcracking on pressure-solution strain rate: the Gratz grainboundary model. Geology, 26, 915-918. DEN BROK, S. W. J., ZAH1D,M. & PASSCHIER, C. W. 1999. Pressure solution compaction of sodium chlorate and implications for pressure solution in NaCl. Tectonophysics, 307, 297-312. DONNELLAN, A. & LYZENGA, D. A. 1998. GPS measurements of fault afterslip and upper crustal deformation following the Northridge earthquake. Journal of Geophysical Research, 103, 21 28521 297. DYSTHE, D. K., PODLADCHIKOV, Y., RENARD, F., FEDER, J. G. & JAMTVEIT, B. 2002. Universal scaling in transient creep. Physical Review Letters, 89, 246 102-1-246 102-4. DYSTHE, D. K., RENARD, F., FEDER, J. G., JAMTVEIT, B., MEAKIN, P. & JOSSANG, T. 2003. High resolution measurements of pressure solution creep. Physical Review E, 68, 011603. GRATIER, J. P. & GUIGUET, R. 1986. Experimental pressure solution-deposition on quartz grains: the crucial effect of the nature of the fluid. Journal of Structural Geology, 8, 845-856. GRATIER, J. P. 1993. Experimental pressure solution of halite by an indenter technique. Geophysical Research Letters, 20, 1647-1650. GRATIER, J. P., RENARD, F. & LABAUME, P. 1999. How pressure solution creep and fracturing processes interact in the upper crust to make it behave in both brittle and viscous manner. Journal of Structural Geology, 21, 1189-1197. GRATZ, A. J. 1991. Solution-transfer compaction of quartzites: Progress towards a rate law. Geology, 19, 901-904. JONES, R. V. 1988. Instruments and Experience. Wiley, New York. LAFAUCHEUX, F. 1974. Contribution h l'dtude de ddfauts pMsent~s par des calcites hydrothermales de synthkse. Thrse d'Etat, Paris VI. MORSE, J. W. & ARVIDSON,R. S. 2002. The dissolution kinetics of major sedimentary carbonate minerals. Earth-Science Reviews, 58, 51-84. PATERSON, M. S. 1973. Nonhydrostatic thermodynamics and its geological applications. Reviews in Geophysics, 11, 355-389. PHARR, G. M. & ASHBY, M. F. 1983. On creep enhanced by a liquid phase. Acta Metallurgica, 31, 129-138. PLUMMER, L. N., WIGLEY, T. M. L. & PARKHURST, D. L. 1978. The kinetics of calcite

PRESSURE SOLUTION ON CALCITE MONOCRYSTALS dissolution in CO2-water systems at 5 ~ to 60 ~ and 0.0 to 1.0 atm C O 2. American Journal of Science, 278, 179-216. RAJ, R. 1982. Creep in polycrystalline aggregate by matter transfer through a liquid phase. Journal of Geophysical Research, 87, 4731-4739. RAMSAY,J. R. 1967. Folding and Fracturing of Rocks. McGraw-Hill, London. REDDY, M. M. & WANG, K. K. 1980. Crystallization of calcium carbonate in the presence of metal ions. I. Inhibition of magnesium ion at pH 8.8 and 25 ~ Journal of Crystal Growth, 50, 470-480. RUTTER, E. H. 1976. The kinetics of rock deformation by pressure solution. Philosophical Transactions of the Royal Society of London, 283, 203-219. RUTTER, E. H. 1983. Pressure solution in nature, theory and experiment. Journal of the Geophysical Society of London, 140, 725-740.

95

SKVORTSOVA,Z. N., TRASKIN,V. Y., LOPATINA,L. I. & PERTSOV, N. V. 1994. Adsorption-induced decrease of the recrystallisation creep rate. Colloid Journal, 56, 177-179. TADA,T. & SIEVER,R. 1986. Experimental knife-edge intergranular pressure solution of halite. Geochimica and Cosmochimica Acta, 50, 29-36. WEYL, P. K. 1959. Intergranular pressure solution and force of crystallization - a phenomenological theory. Journal of Geophysical Research, 64, 2001-2025. ZANG, X., SALEMANS,J., PEACH, C. J. • SPIERS, C. J. 2002. Compaction experiments on wet calcite powder at room temperature: evidence for operation of intergranular pressure solution. In: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. PENNOCK, G. M. (eds) Deformation Mecha-

nisms, Rheology and Tectonics: Current States and Future Perspectives. Geological Society, London, Special Publications, 200, 29-39.

Contrasting microstructures and deformation mechanisms in metagabbro mylonites contemporaneously deformed under different temperatures (c. 650 ~ and c. 750 ~ L. B A R A T O U X 1'2'5, K. S C H U L M A N N 3, S. U L R I C H t'4 & O. L E X A t

~Institute of Petrology and Structural Geology, Charles University, Albertov 6, 12843, Prague, Czech Republic (e-mail: Ika @natur.cuni.cz) 2UMR 5570, ENS and Lyon 1 University, 2 rue RaphaYl Dubois, 69622, Villeurbanne Cedex, France 3Universitd Louis Pasteur, EOST, UMR 7517, 1 Rue Blessig, Strasbourg, 67084, France 4Geophysical Institute, Czech Academy of Sciences, Bo(n{ 11/1401, 14131 Praha 4, Czech Republic 5Present address." Czech Geological Survey, Klarov 3, Praha 1, 11821, Czech Republic Abstract: Deformation mechanisms of amphibole and plagioclase were investigated in two metagabbroic sheets (the eastern and western metagabbros from the Stars M~sto belt, eastern Bohemian Massif), using petrology, quantitative microstructural and electron backscattered diffraction methods. After the gabbroic pyroxene was replaced by amphibole, both gabbroic bodies became progressively deformed. The eastern metagabbros were deformed under temperature of c. 650 ~ and the western metagabbros under c. 750 ~ Subgrain rotation and dislocation creep, characterized by strong crystallographic and shape preferred orientations, operated in plagioclase of the eastern belt during the early stages of deformation. Subsequent randomizing of plagioclase crystallographic preferred orientation is interpreted to be due to grain boundary sliding in the mylonitic stage. Large (50150 ixm) grain sizes during the mylonitic stages are interpreted to be due to low strain rates. Amphibole is stronger and deforms cataclastically, leading to important grain size reduction when the bulk rock strength drops substantially. In the western belt, plagioclase deformed by dislocation creep accompanied by grain boundary migration (possibly chemically induced) while heterogeneous nucleation and syndeformational grain growth in conjunction with dislocation creep were typical for amphiboles.

Microstructural and rheological behaviour of natural polyphase rocks is a complex problem, which has been studied in detail mostly in quartzo-feldpathic rocks (e.g. Gapais 1989; Handy 1990). In these rocks, feldspars are generally considered as strong minerals while quartz represents the weak phase. Handy (1994) proposed a scheme for rocks containing minerals of contrasting rheology with two end-members: load-beating framework (LBF) and interconnected weak layers (IWL), based on the assumption that at least one mineral (generally the weaker one) is deformed by the mechanism of dislocation creep. As pointed out by Jordan (1988) the LBF is not stable with increasing strain, resulting in mechanically induced compositional layering. The strength

of a two-phase material depends on the proportion of the weak mineral and on the rheological contrast between the two phases (Jordan 1988; Handy 1990, 1994). However, if the deformation mechanism switches from crystal plasticity to some grain size sensitive process, the strength of the bulk rock drops rapidly with respect to the preceding model due to strain softening (Knipe 1989). It was also shown that generally stronger feldspars may become even weaker than quartz if they deform by diffusional creep (e.g. Voll 1976; Simpson 1985; Gapais 1989; Martelat et al. 1999). Equivalent comparative microstructural study of plagioclase and hornblende in naturally deformed metabasic rocks as well as experimental studies of the rheology of amphibole-plagioclase

From: GAPA1S,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. DeformationMechanisms,Rheologyand Tectonics:from Mineralsto the Lithosphere. Geological Society, London, Special Publications, 243, 97-125. 0305-8719/05/$15.00 ~') The Geological Society of London 2005.

98

L. BARATOUX E T A L .

bearing rocks are rare (e.g. Hacker & Christie 1990; Wilks & Carter 1990). However, metabasites are considered to constitute a major part of the lower continental crust (Rutter & Brodie 1992) and the study of deformation microstructures and mechanisms is therefore crucial for understanding the deformation behaviour and rheology of mafic tectonites. Plagioclase rheology is believed to control the strength of the lower crust in several models (e.g. Carter & Tsenn 1987; O r d & Hobbs 1989). Hornblende is assumed to be a relatively strong phase and to behave passively during deformation unless it forms a load-supporting framework (Brodie & Rutter 1985). Many studies concerned with deformation mechanisms of plagioclase and hornblende have been published to date. Experiments suggest that plagioclase is deformed by dislocation creep under lower crustal conditions (e.g. Tullis & Yund 1987; Ji & Mainprice 1990; Kruse et al. 2001) with commonly active (010)[001] principal slip system (e.g. Olsen & Kohlstedt 1985; Kruhl 1987). Kruhl (1987) also observed a (001) [ 100] slip system in naturally deformed plagioclases and Stiinitz et al. (2003) have shown that slip on (001) and {111} in (110) direction are similarly active in experimentally deformed An60 crystals. Besides these examples, other less common slip systems were suggested by Marshall & McLaren (1977a, b), Montardi & Mainprice (1987), and Olsen & Kohlstedt (1984, 1985). In addition, grain size sensitive processes in plagioclases have been referred to in some studies (Boullier & Gu6guen 1975; Lapworth et al. 2002). Hornblende generally undergoes brittle deformation under low temperature conditions (e.g. Brodie & Rutter 1985; Nyman et al. 1992; Lafrance & Vernon 1993). Many studies document crystal plastic deformation of hornblende with a dominant (100) [001] slip system from experiments and natural rocks (e.g. Rooney et al. 1970; Cumbest et al. 1989a, b; Hacker & Christie 1990). Volume diffusion rates are, however, slow in amphiboles (e.g. Freer 1981) implying that dislocation climb is limited even at geological strain rates. Syndeformational chemical reactions between hornblende and plagioclase are common in metabasic rocks (Brodie 1981; Brodie & Rutter 1985) and involve deformation mechanisms such as chemically induced grain boundary migration (CIGM) (Cumbest et al. 1989a), on nucleation of new plagioclase (Rosenberg & Stfinitz 2003) or hornblende at plagioclase grain boundaries (Kruse & Sttinitz 1999). Metagabbros from the Star6 M6sto belt (eastern margin of the Bohemian Massif) represent

an example of a two-phase metabasic system deformed at different temperatures and strain gradients. The aim of this study is to show two types of progressive evolution of deformation microstructures and textures of dynamically recrystallized plagioclase-hornblende bearing metagabbros at amphibolite and upper amphibolite facies conditions with increasing bulk strain. Methods such as quantitative textural analysis, crystallographic preferred orientations (CPO) and study of mineral chemistry were used to constrain deformation mechanisms for both minerals. Finally, deformation mechanisms of rheologically contrasting plagioclase and hornblende are correlated and the mechanical behaviour of mafic rocks deformed under lower crustal metamorphic conditions is discussed.

Geological setting The Star6 M6sto (SM) belt in the eastern margin of the Bohemian Massif separates the high grade gneisses of thickened continental crust of the Lugian domain in the west from a NeoProterozoic continental margin in the east (Fig. l a). The SM domain was thinned during Cambro-Ordovician rifting and underwent granulite facies metamorphism (Stfpskfi et al. 2001). Subsequent Variscan tectonics resulted in N E SW trending structures dipping at relatively high angles to the west (Figs 1a and b). Strongly deformed 'western' metagabbros of Cambro-Ordovician protolith ages occur at the top of the SM belt (Fig. 1) (Kr6ner et aL 2000), The metagabbros were pervasively affected by a ductile shear zone along which was emplaced syntectonically a Carboniferous tonalite sill dated at 340 Ma (Stfpskfi et aL 2004). A carboniferous metamorphism of adjacent metagabbros reached higher amphibolite facies conditions because of the strong heat input from the tonalite sill. A leptyno-amphibolite complex of CambroOrdovician age comprising a sequence of alternating amphibolites and tonalitic gneisses occurs in the footwall of the tonalite sill. This complex is underlined by the 'eastern' metagabbro sheet, which is supposed to form part of the same lower crust as the western (upper) metagabbro sheet (Stfpskfi et aL 2001). Amphibolite facies Carboniferous metamorphic conditions of the eastern (lower) metagabbro are also attributed to the heat input of a more distant hangingwall tonalite intrusion. Rocks of Cambro-Ordovician age suffered two deformations: D~ of Cambro-Ordovician age and D2 of Carboniferous age. The latter one prevails in most lithologies of the SM belt, marked by a penetrative west-dipping $2 foliation bearing a

TEXTURES OF NATURALLY DEFORMED METAGABBROS

99

Fig. 1. (a) Location of the studied area in the frame of the European Variscides. Geological map of the Star6 M~sto belt is based on unpublished geological maps at 1:25 000 provided by courtesy of the Czech Geological Survey (Dr. M. Opletal, author). Important thrusts and faults as well as location of samples and a cross-section A-A' (Fig. lb) are indicated. (b) Schematic cross-section based on field observations shows major structures, lithology and major tectonic boundaries. (Vertical axis not to scale.) subhorizontal or gently inclined N - S trending mineral lineation. The eastern metagabbros in the hangingwall of the high-grade rocks of the Silesian d o m a i n were affected by localized D2 shear zones that developed under amphibolite

facies conditions attained during the peak of the Carboniferous m e t a m o r p h i s m . Identical geometry of structures and kinematics suggest that the eastern and western metagabbro belts were d e f o r m e d under a dextral transpressive shear

100

L. BARATOUX ETAL.

regime but under different thermal conditions (Stfpsk~i et al. 2001).

Methods and techniques Mineral chemistry

Minerals were analysed with a Cameca SX 100 electron microprobe equipped with four WDS spectrometers at Blaize Pascal University, Clermont-Ferrand, France. Operating conditions were 15kV, 10hA beam current, 2-5p~m beam diameter, 20 s counting time, and natural mineral standards. Some hornblende and plagioclase were analysed using a CamScan $4 scanning electron microscope and attached Link EDX microanalytical system, at Charles University, Prague, Czech Republic. Operating conditions were 20 kV, 1.8 nA beam current, 1-3 p,m beam diameter, 120 s counting time, and mineral standards Structure Probe Instruments (SPI). Ca maps from plagioclase were made at Claude Bernard University in Lyon, with operating conditions 15 kV, 15 nA beam current and spatial resolution of 512 x 512 pixels. Quantitative microstructural analysis

Quantitative microstructural analysis of grain boundaries was carried out on traced and digitized outlines of grains in ESRI ArcView 3.2 Desktop GIS environment. The map of grain boundaries was generated using ArcView extension Poly (Lexa 2003). The resulting polygons have been treated by MATLAB TM PolyLX Toolbox (http://petrol.natur.cuni.cz/~ ondro; Lexa 2003), in which grain shape and grain boundary preferred orientations (SPO and GBPO, respectively) were analysed using the moments of inertia ellipse fitting and eigenanalysis of bulk orientation tensor techniques (Lexa, 2003; modified SURFOR technique by Panozzo (1983) for GBPO). Their degree is expressed as the eigenvalue ratio (r = E1/E2) of the weighted orientation tensor of grain shapes or boundaries. The orientation is defined by V~ and V2 eigenvectors. The grain sizes of the minerals were calculated in terms of Ferret diameters of grain section without stereological corrections. Crystallographic preferred orientation

Amphibole and plagioclase crystallographic preferred orientations (CPO) were collected using a scanning electron microscope CamScan $4 in Prague and a JEOL JSM 5600 in Montpellier equipped with Channel5 electron backscatter

diffraction (EBSD) system from HKL Technology (Prior et al. 1999). Thin sections were polished using 0.25 ~m diamond paste. To remove all surface damage and minimize relief between minerals, sections were chemically polished using a colloidal silica suspension. All thin sections were carbon-coated. The coating reduces the quality of the electron backscatter diffraction patterns (EBSP) so that automatic indexing mode of the EBSP system could not be used. Most data were therefore collected manually. Operating conditions were 20 kV in Prague and 15 kV in Montpellier, 5.6 nA beam current, working distance 39 ram, and 2 - 5 ~m beam diameter. For each measurement, three Euler angles (vl, 05, v2) characterizing the lattice orientation as well as the nature of the mineral were determined and stored. Practice shows that plagioclase diffraction patterns do not change significantly from albite up to at least An65 (Lapworth et al. 2002). Therefore, the Anorthite48 database was used for indexing plagioclase. Pole figures and inverse pole figures were projected using the software developed by D. Mainprice (ftp://ftp.dstu.univmontp2.fr/pub/TPHY/david/pc). Projections of crystallographic axes ([a], [b] and [c]) or optical indicatrix (o~, /3, and 3/) are generally used for plagioclase (e.g. Jensen & Starkey 1985; Olsen & Kohlstedt 1985; Ji & Mainprice 1988, 1990; Prior & Wheeler 1999). The degree of CPO has been quantified using orientation tensor of crystallographic planes and directions (Mainprice, ftp://ftp.dstu.univmontp2.fr/pub.TPHY/david/pc). The intensity of CPO is given by the I parameter (Lisle 1985): 3

I = 15/2 x Z ( E i

- 1/3) 2

i=1

where E~, E2, and E3 represent the eigenvalues of the preferred orientation of poles to planes and directions of amphibole and plagioclase grains plotted in pole figures. The values of I range between 0 (no preferred orientation) and 5 (all fabric elements perfectly parallel one to another).

Microstructures The metagabbros from the eastern (lower) belt are affected by localized shear zones (Fig. lb) and all stages from non-deformed rock, protomylonite, mylonite, and up to highly strained ultramylonite are present (Figs 2a, b, c and d). In the western metagabbro (upper) belt, the protolith stage is not preserved and only two deformational stages can be distinguished

TEXTURES OF NATURALLY DEFORMED METAGABBROS

101

Eastern belt (-650 ~

Western belt (-750 ~

Fig. 2. Macro photographs of typical metagabbro structures from the eastern (a-d) and the western (e, t3 belts: (a) non-deformed metagabbro (E0); (b) protomylonite (El); (c) mylonite (E2); (d) ultramylonite (E3); (e) augen mylonite (W1); (f) banded mylonite (W2).

(Figs 2e and f): mylonites with hornblende porphyroclasts related most probably to the Cambro-Ordovician metamorphic event and completely recrystallized mylonites characterized by a monomineral layering.

Metagabbros of the eastern belt consist of plagioclase (40-60%), amphibole (40-60%), relict pyroxene in the cores of some large amphibole grains (less than 1%), and titanite (less than 1%). No substantial change in modal proportions

102

L. BARATOUX ETAL.

is observed with increasing deformation. Modal proportions of amphibole and plagioclase in the western metagabbros are more variable than those in the eastern belt. Amphibole proportion may vary between 20 and 80%, most likely due to original magmatic compositional variations. For the purpose of our study, samples composed of c. 50% of amphibole and c. 50% of plagioclase were chosen.

Deformation of the eastern (lower) metagabbro sheet Non-deformed metagabbro (EO). At low strain, plagioclase and hornblende exhibit euhedral randomly oriented crystals of 0.5-3 mm in size. Tapering mechanical twins and local undulatory extinction occur in plagioclase. There is no evidence for any kind of dynamic recrystallization or crystallization of new grains. Undulatory extinction locally attests to some bending of amphibole grains. Amphibole porphyroclasts show random spatial distribution and they are generally not interconnected.

Protomylonite (El). At higher strains, about 20-25% of the total volume of plagioclase grains but only 8-10% of amphibole show

evidence of strain, suggesting that the deformation was accommodated mostly by plagioclase recrystallization and associated grain-size reduction. Plagioclase grains show polysynthetic twins according to albite and pericline laws (Tullis 1983) and patchy undulatory extinction. Large plagioclase porphyroclasts of 2 - 5 mm are cut by brittle fractures (Fig. 3a) reducing the grain size to 0.5-1 mm. The fractures are filled with small twin-free recrystallized grains of 0.02-0.1 mm. Two recrystallization mechanisms have been identified: bulging and subgrain rotation recrystallization (Fig. 3b), as proposed by Poirier & Guillop6 (1979) or Fitz Gerald & StiJnitz (1993), leading to core-mantle structures. Core and mantle structures were observed with an intermediate zone of subgrains and new grains of similar size developed along porphyroclast boundaries. Boundaries between the neoblasts become progressively straight, meeting at triple junctions of 120 ~. Large porphyroclasts of hornblende reveal strong internal deformation such as kinking, bending leading to sweeping undulatory extinction and (100) twinning. The twin planes are not regular and they locally form finger-like structures. Brittle fractures transecting large grains are often present. Randomly oriented porphyroclasts of 2 - 5 m m in size rotate

Fig. 3. Drawing and micrographs(XPL) of the eastern metagabbro protomylonite. (a) Initial stage of plagioclase deformation characterized by brittle fractures cross-cutting large porphyroclasts. Digitizeddrawing was used for the quantitative textural analysis.Arrows mark three clasts derived by fracturing of one porphyroclast and show corresponding grains in the digitized drawing and micrograph. Plagioclaseis white, hornblendeis light grey, and opaque minerals are black. (b) New grainsdevelop preferentiallyalong fractures by a mechanismof subgrainrotation (SR). Note that the neoblasts are twin-free. (P) refers to porphyroclast.

TEXTURES OF NATURALLY DEFORMED METAGABBROS progressively their cleavage planes {110} parallel to the foliation. Small needle-like grains (0.020.05 x 0.1-0.3 mm in size), arranged parallel to the mylonitic foliation, are present in highstrain domains.

Mylonite (E2). Within the eastern metagabbros, plagioclase porphyroclasts are recrystallized into elongate aggregates or bands wrapped around hornblende clasts, and porphyroclasts represent only 5 - 1 0 % of the total plagioclase volume. Elongate needle-like hornblende grains are

103

arranged into aggregates or bands parallel to the main mylonitic foliation (Fig. 4a). Isolated hornblende needles scattered within plagioclase-rich domains are inclined at an angle of 2 0 - 3 0 ~ with respect to the main fabric, indicating noncoaxial deformation (Figs 4a and b). Plagioclase grains attain 0.1-0.25 mm in size. Increase in grain size with respect to the protomylonite stage and common optical zoning of recrystallized grains due to increase in anorthite content suggest syndeformational growth. Matrix grains are often twin-free; the twinned grains represent

Fig. 4. Digitized drawings and micrographs of the eastern mylonite. (a) Mylonitic foliation with alternating layers of amphibole and plagioclase aggregates (PPL). (b) Plagioclase of subequant shapes and grain size typical for the mylonitic stage (E2-pll). (c) High-strain plagioclase domain (E2-p12) adjacent to an amphibole porphyroclast is characterized by elongate shapes and strong SPO (XPL). (d) Amphibole porphyroclast (E2) cross-cut by microshear zones filled by small needle-like grains (XPL). Undulatory extinction documents strong internal deformation. Arrows mark corresponding grains in the digitized drawings and micrographs.

104

L. BARATOUX ETAL.

only 10-20% of the total plagioclase volume. Plagioclases generally exhibit subequant shapes with straight grain boundaries meeting at 120 ~ triple joints (Fig 4b). However, in the vicinity of amphibole porphyroclasts, the plagioclase grains are more elongate with higher aspect ratios up to 2.5 (Fig. 4c). Hornblende grains are arranged in anastomosing interconnected bands defining the mylonitic foliation. Most of the hornblende matrix grains, 0.1-0.3 mm in length, are elongate in XZ sections attaining high aspect ratios (up to 6) while they are lozenge-shaped in the YZ sections with aspect ratios of 1.5-3. A total of 10-20% of the hornblende is present in a form of locked-up porphyroclasts of 0.2-2 mm in size showing undulatory or patchy extinction characteristic of strong internal strain. These porphyroclasts are elongate parallel to the foliation. Porphyroclasts are often transected by microshear zones (Fig. 4d) and small needle-shaped grains consequently develop within these zones by activation of the weak cleavage system {110}.

Ultramylonite (E3). The ultramylonitic metagabbros of the eastern sheet are marked by very uniform mineral banding of fine-grained hornblende and plagioclase layers ranging between 1 and 5 mm in width (Fig. 5a). Asymmetric fabric of hornblende and plagioclase is no longer present. Plagioclase forms continuous bands including abundant interstitial grains of amphiboles

(Fig. 5b). In contrast, amphibole-rich bands are almost monomineral. Plagioclase is completely recrystallized, varying in size between 0.05 and 0.3mm. The grains have smooth rounded shapes locally elongate parallel to the foliation. Rounded grains of 0.5 mm in size, interpreted as relicts of original porphyroclasts, locally occur in the fine-grained matrix. Plagioclase grains are affected by strong retrograde sericitization and albitization. Interstitial and matrix grains of hornblende reach 0.03-0.1 x 0.1-0.3 mm in size. Lockedup porphyroclasts of amphibole (0.5-1 mm in size) occur parallel to the foliation. However, they are less common than in the mylonite (E2). All amphibole grains are characterized by very strong SPO (Fig. 5c). The mylonitic foliation is cross-cut by extensional veins filled by a mixture of epidote and amphibole. These veins show that fluid activity was high after the main mylonitic episode. Albitization and sericitization of plagioclase in some samples suggest that increased amount of water was present in these rocks also during late retrogression, which is most likely not related to the main process of mylonitization.

Deformation of the western (upper) metagabbro sheet Augen mylonite (W1). In the western metagabbro belt, the deformation was pervasive so that initial protomylonite stages of deformation

Fig. 5. Drawingsand micrographof the eastern ultramylonite.(a) Plagioclaseand amphibolebanding (PPL). Plagioclase is darker than amphiboledue to strong sericitization.(b) Plagioclaseband in ultramylonite(E3) with abundant interstitialhornblende. Small grain size and subequant shapes are typical for these plagioclases. (c) Amphibolelayer from ultramylonite(E3) is characterizedby very fine grain size and strong SPO.

TEXTURES OF NATURALLY DEFORMED METAGABBROS are absent. The least deformed rock represented by augen mylonite is marked by the presence of rare 2 - 3 mm large porphyroclasts of plagioclase (Fig. 6a) and by 1 - 2 r a m wide and 4 - 5 mm long porphyroclasts of hornblende within the recrystallized matrix (Fig. 6b). This rock type is preserved only in the southern part of the metagabbro belt. Plagioclase porphyroclasts represent only 5-10% of the total plagioclase volume and show deformational twins and undulatory extinction. Recrystallized plagioclase grains show optical zoning. Modal proportion of amphibole porphyroclasts cannot be estimated as they differ from the neoblasts only in their chemical composition, making them optically indiscernable. Amphibole porphyroclasts exhibit patchy or sweeping undulatory extinction but deformation twins have not been observed. The local occurrence of twins is assumed to be growth-related, the twin planes being always perfectly straight. Prismatic porphyroclasts attain higher aspect ratios than more rounded neoblasts. All rocks in the western belt are affected by amphibolite facies re-equilibration partly destroying the granulite facies peak metamorphic assemblages.

105

Banded mylonite (W2). Distinct monomineral bands of 1-10 mm in width are typical of these highly deformed mylonites. Small hornblende grains of irregular shape are common within the plagioclase layers, but plagioclase is rarely included within amphibole bands. Both minerals are completely recrystallized. Hornblende is in some places replaced by later prisrnatic cummingtonite, which grows either parallel or obliquely to the mylonitic foliation and is locally included within the plagioclase layers. Two types of plagioclase microstructures occur within one plagioclase band. The first type is characterized by grain sizes ranging from 0.3 to 1 mm in diameter, subequant or slightly elongate, but always irregular, grain shapes with serrated boundaries (Fig. 7a), and growth-related optical zoning. Mechanical twins cross-cut the optical zoning in some cases (Fig. 7b), suggesting that at least some of them formed later. The second type is characterized by narrow (c. 0.5 mm) high-strain zones trending parallel or slightly oblique to the layering (Fig. 7c). Here, the grain size is reduced to 0.05-0.2 ram. Plagioclase grains attain higher aspect ratios (up to 2) and exhibit undulatory extinction and subgrain boundaries elongate parallel to the high-strain zone. The contact between these two deformational domains is either sharp or transitional. Hornblendes of 0.1 - 1 mm in size have straight grain boundaries locally meeting at high angles suggesting a high degree of textural equilibration (Vernon 1976) (Fig. 7d). No undulatory extinction occurs and twinning is very rare. Where present, the twin planes are always very straight, indicating that twinning is growth-related rather than deformational. Aspect ratios vary between 1.5 and 2.5, but some grains with aspect ratio close to 1 occur. Grains with high aspect ratios are not always parallel to the mylonitic foliation.

Bulk rock chemistry, mineral chemistry and zoning Bulk rock chemistry

Fig. 6. Drawingsand micrographsof the western augen mylonite. (a) Digitizeddrawing of plagioclase(W1) used for the quantitative textural analysis.Large porphyroclasts are surrounded by a mantle of finegrained matrix grains. (b) Hornblende(W1) shows variable grain size and strong SPO. Plagioclaseis white, hornblende is light grey, and opaque mineralsare black.

Table 1 presents bulk rock chemical analyses studied by X-ray fluorescence at Claude Bernard University, Lyon, France. The eastern and western metagabbro sheets are regarded to be a part of the same lower crustal unit prior to their involvement into the Variscan tectonics (Stfpskfi et al. 2001). However, their chemistry is not identical and bulk rock analyses (Table 1) show that western metagabbros are characterized by higher contents in A1 and Ca and lower contents in Na and Fe than their

106

L. BARATOUX ET AL.

Fig. 7. Drawings and micrographs of the western banded mylonite. (a) Lobate boundaries document grain boundary migration (GBM) in coarse-grained plagioclases of banded mylonite (XPL). (b) Mechanical twins (MT) transect the growth-related zoning (GZ) in banded mylonites documenting a later deformation phase (XPL). (c) Digitized drawing of two plagioclase domains (W2-pll and W2-p12) used for quantitative textural analysis. (d) Hornblendes (W2) are characterized by straight grain boundaries meeting occasionally at triple points at variable angles. Many small grains indicating high nucleation rate are present. SPO is rather low. Plagioclase is white, hornblende is light grey, pyroxene is dark grey, and opaque minerals are black in all drawings.

eastern counterpart. The increase in Na and depletion in Ca in both metagabbro belts is related either to mylonitization or to primary variations due to postmagmatic processes. Trace elements show very similar trends for the eastern and western metagabbros, suggesting that they may originate from the same source. Metagabbros from both belts display a slightly negative Nb anomaly normalized to MORB (Sun & McDonough 1989). Augen mylonites from the western belt are depleted in Zr and Ti compared to eastern mylonites. Contents of Rb, Ba, and K are strongly variable in both belts, which is most probably related to postmagmatic alterations and/or mobility of these elements during metamorphism.

Mineral chemistry and zoning Syndeformational chemical reactions are common in metabasic rocks (Brodie 1981; Brodie & Rutter 1985). Variations in plagioclase composition are shown in Figures 8 and 9. Amphibole compositions according to the classification of Leake et al. (1997) are plotted in Figure 10. Representative analyses of mineral compositions are listed in Tables 2 and 3.

Mineral abbreviations are according to Kretz (1983). It is likely that the studied metagabbros experienced a metamorphic event prior to the main deformation and metamorphism studied in this work. This is supported by existence of incompletely amphibolized pyroxenes in undeformed rocks of low-strain domains. In addition, the compositions of cores of large amphibole porphyroclasts (actinolites to actinolitic magnesiohornblendes) deviate from those of recrystallized magnesio-hornblendes in the eastern belt. Such compositions of primary amphiboles may indicate early greenschist to amphibolite facies conditions preceding the main higher grade Carboniferous deformation. The peak metamorphic assemblages within the eastern metagabbro consist of P1 + Hbl + Qtz + Ttn _+ Ilm __%Mag. The composition of plagioclase porphyroclasts in protomylonite (El) varies between Anso and An6o (Fig. 8a). Small recrystallized grains show a composition similar to the mother host grain corresponding to Anso-6o. More albitic compositions (An4o_45) occur along grain boundaries or triple junctions of recrystallized grains (Fig. 9a). The geometry and sharp gradient in mineral zoning crosscutting several grains is attributed to post-peak

TEXTURES OF NATURALLY DEFORMED METAGABBROS

107

Table 1. Representative analyses of bulk rock compositions Eastern belt

Western belt

Protolith (E0) Mylonite (El) Ultramyl. (E3) Augen myl. (Wl) Banded myl. (W2) Banded myl. (W2) SiO2 TiO2 A1203 FeO T MnO MgO CaO Na20 K20 P205 LOI H20Total

48.61 0.71 17.76 6.14 0.12 9.31 12.47 2.64 0.16 0.04 0.88 0.06 98.90

49.20 0.82 19.74 6.32 0.10 6.56 9.22 3.29 1.52 0.03 1.91 0.12 98.83

52.69 0.94 18.04 7.33 0.13 5.81 8.60 4.99 0.10 0.03 0.57 0.05 99.28

46.23 0.10 21.57 4.86 0.10 7.91 13.15 0.82 1.60 0.01 2.76 0.19 99.30

50.20 0.35 21.43 4.10 0.08 6.20 12.08 2.98 0.51 0.02 1.21 0.09 99.25

48.92 0.37 24.20 3.84 0.06 4.76 11.58 3.55 0.25 0.07 1.65 0.10 99.35

Ba Rb Sr Zr Nb Y V Cr Ni CO Sc Cu Pb Ga XMg*

32.1 1.6 409.4 33.0 1.3 12.2 185.9 550.5 122.8 36.8 42.3 8.0 0.8 13.8 0.730

149.0 55.1 742.0 34.7 1.3 9.4 160.1 214.7 60.6 35.4 24.9 3.6 4.5 15.0 0.649

52.0 4.9 226.9 39.3 1.7 20.2 203.9 222.4 58.8 29.2 30.4 19.1 2.5 12.2 0.586

374.3 74.7 219.1 7.8 0.6 3.3 115.0 136.0 48.9 32.0 41.3 2.9 6.0 12.8 0.744

52.3 26.9 346.3 10.7 1.3 7.2 131.9 622.9 87.3 26.6 33.6 20.4 2.4 16.9 0.729

54.6 11.9 397.4 29.6 1.2 6.4 92.0 269.9 90.9 25.4 16.8 33.3 2.3 17.1 0.688

Major elementsare given in wt%, trace elementsare given in ppm. *XMg= Mg/(Fe2+ + Mg).

a) 12-

Porphyroclasts

E a s t e r n belt

b)

Western

Porphyroclasts

belt

trogression

g o8

~_ 6

i

IIIN!I I].

.

.

I] ,. I

I

,I

Recrystallizedgrains I Late retrogression

2HI 10

20

30

40

50

% An

i

i

i

,

i

,

i

i

III

,

IIII i

i

,

i

i

~ protomylonite(El) Recrystallizedgrains ~[Z] mylonite(E2) r~ ultramylonite(E3) I

~I---1 Banded Augenmylonite rnylonite(W1)1 (W2)l

I,

111 t h

60

70

80

90

10

20

30

40

50

% An

60

70

80

90

Fig. 8. Composition of plagioclase in (a) the eastern and (b) the western belts. Porphyroclasts are depicted separately from matrix grains. Deformation grade increases from dark grey to white colour.

108

L. BARATOUX ET AL.

Fig. 9. BSE images of plagioclaseand amphibole and representative chemical maps of plagioclase compositional variations. (a) Plagioclasein protomylonite from the eastern belt (El) suffersfluidrelated chemical variationsmodifyingthe composition at triple points and grain boundaries. (b) Growth-related chemical zoning in mylonitefrom the eastern belt (E2). (c) Growth-related zoning in coarse-grainedplagioclase of banded mylonitefrom the western belt (W2).

fluid-phase activity along grain boundaries (McCaig & Knipe 1990). Recrystallized matrix plagioclases in the mylonite (E2) have andesitic compositions (An3o-5o) and show increases in anorthite content towards their rims (Fig. 9b), which is commonly ascribed to increasing metamorphic conditions. Rare rounded porphyroclast, as well as matrix grains in ultramylonites (E3), underwent strong and irregularly developed late albitization. Matrix grains correspond to An3o_43 although oligoclase and albite compositions (An6_ 15) are also common along fractures, cleavages and grain boundaries. These irregular compositional variations are attributed to late retrogression unrelated to the main deformation process. Amphibole porphyroclasts in protomylonites (El) are magnesio-hornblendes with variable content in Si but constant XMg (Fig. 10a).

Amphiboles with more actinolitic composition replacing former magmatic pyroxenes can be found in the cores of some grains. Small amphibole grains show similar compositions compared to the porphyroclasts, some of them being more tschermakitic. The amphiboles in mylonites (E2) are magnesio-hornblendes with slightly lower Si contents than those in the protomylonites. High Si/low A1 domains documenting incomplete transformation of pyroxenes into hornblendes are present in the cores of large porphyroclasts as well. Locked-up grains in the ultramylonites (E3) vary between magnesiohornblende and tschermakitic composition. Small new grains have slightly lower A1 contents compared to the less deformed stages. In the western belt, the peak metamorphic assemblages correspond to P1 § Hbl _ Cpx _+ Opx +_ Grt ___Ttn + Ilm. A granulite facies mineral assemblage, comprising also sapphirine and corundum, is not stable and re-equilibrated during subsequent upper amphibolite facies reworking. In the south, granulite facies assemblages are absent and probably completely re-equilibrated. In the north, a late metamorphic phase is documented by the presence of prismatic orthoamphibole (gedrite). Plagioclase porphyroclast compositions are very similar to those of the eastern belt and vary between An4o and An65. Recrystallized matrix grains in the augen mylonite (W1) show locally increasing anorthite content up to An6o_9o compared to the porphyroclasts, indicating that chemical reactions took place during recrystallization. Sharp limited domains of An4o_45 composition along grain boundaries are also common, suggesting late fluid activity (Knipe & McCaig 1994). Recrystallized plagioclase grains in banded mylonites (W2) show stronger variability in compositions (An35-9o). Growth-related zoning due to increasing metamorphic grade is documented in Figure 9c. The cores have more albitic composition (An34) compared to the rims (An6o_65) and are attributed to the early metamorphic stage. Modifications of plagioclase composition characteristic of fluid-related compositional changes are also developed in the banded mylonite. Amphibole porphyroclast compositions (Fig. 10b) were analysed in two different thin sections of augen mylonite (W1) and the difference in their mineral chemistry may be explained by local bulk rock chemical variations. The amphibole porphyroclasts correspond to magnesiohornblendes with similar compositions to those in the eastern belt. Porphyroclast rim compositions show a decrease in Si content with respect to the cores. Neoblasts are marked

TEXTURES OF NATURALLY DEFORMED METAGABBROS 9 9 b, o

1,0 0.9

Core of porphyroctast Rim of porphyroclast Core of neoblast Rim of neoblast

I

1.0--

Magnesio-homblende

Tremolite Aclinolite

9

Tschermakite

9 A 9 A,~ Ultramylonite(E3)

foo:o

I ]-schermakite

Mylonite (E2)

I Core of yroclast

o

~

Magnesio-hornblende

Tremolite Actinolile

0.9

109

o

!

~rrotschermakite

5.75

Tschermakite

i Protomylonite (El) Core of

Ai

porphyroclast

o irrotschermakite } . - - - -

[

__Mg~ (Mg2++Fe2+)

{

....

5,75

! i I

O.C Ferro-actinolite 7.50 8.00

6.50

1.0 i 0.9i

1.0 I 0.9

5,75

Magnesio-hornblende

Tremolite Actinolite

Magnesio-hornblende

Tremolite Actino}ite

~,~.~ o o ,.1~,~

~ . . _ ~

oO~oo ....

_

~ o o 0,9o--

Ferrotschermakite -

(Mg2++Fe2+)[

8.00

Tschermakite ~

Tschermakite

Porphyroclast 2

:00

oo

Banded mylonite (W2) Porphyroctast 1

_Ng~

Ferrotschermakite i

Ferro-hombtende

~.50

5,75

Augen mylonite (Wl) Ferro*actinolite

Ferro-homblende

7.50)

Ferrotschermakite

6.50

5.75

4

Fig. 10. Composition of amphiboles in (a) the eastern and (b) the western belts. (Classification of Leake et was used, (Na + K)A < 0.5.)

by sharp compositional differences expressed by decreasing Si and XMg contents. In the banded mylonite (W2), larger grains show higher Si content than small neoblasts (very low Si content), both being of tschermakitic composition. Temperature (T) was calculated using the hornblende- plagioclase thermometer by Holland & Blundy (1994; thermometer B). Estimated temperatures are based on 30 and 40 amphibole-plagioclase couples in the eastern

al.

(1997)

and western belt, respectively. In the eastern (lower) sheet, calculated T was estimated to be 650 4-50~ while T in the west was 750 + 50 ~ Deformation is interpreted to have taken place under or close to these metamorphic conditions. Pressure (P) could not be estimated as garnet is absent in the mineral assemblage of the metagabbros. However, P of 9 kbar was estimated at the lower contact between the metagabbro sheet and tonalitic sill, where garnets have been formed (St/pskfi e t al. 2001).

110

L. BARATOUX ETAL.

r~

t'-,I

MdMdddd~dd r~

< r.-~

t"-

m

9

~ d ~ d d d d d d ~

TEXTURES OF NATURALLY DEFORMED METAGABBROS

111

e4

M

o ('4

,,.0 ~

I'~ ~0 oo e4 oo ~1- P--. e~h

~"~ ',4~ ~

',~ t"q ~0 ',,0 0", tt3 c4 0", ,--~

9

~-~.~

~8

4-

~dd

dd,~,~d~

~dd,~ddd4~ddd II

-F +

d 4-

o .1...o

0

~

~8 r~

.~

. ~

.qq

~

. q ~

. ~

.~

L. BARATOUX ET AL.

112

45 ixm (El) (Fig. l la). Matrix grains in mylo-

Quantitative textural analysis

Grain size distribution The average grain sizes are presented as median values of the Ferret diameter and the grain size spread is expressed by the difference of third and first quartile (Q3- Ql). The statistical values of the quantitative microstructural and grain size analyses are summarized in Table 4.

Plagioclase.

Plagioclase in XZ section of the

protomylonitic stage in the eastern metagabbro belt is characterized by an average grain size of

nites are marked by an increase of the average grain size, reaching 74 pore in less (E2-pll) and 58 Ixm in more (E2-p12) strained domains. Grain coarsening in the mylonite is consistent with the syndeformational growth of grains originated from subgrain rotation in the protomylonitic stage. This is supported by a greater amount of larger grains in grain size frequency histograms. The grain size is r e d u c e d to ~ 4 5 Ixm in the ultramylonite (E3). Plagioclase grains in the augen mylonite of the western metagabbro belt (W1) are 43 Ixm on average, which is the statistical value

Table 4, Statistical values of the quantitative textural analysis Eastern belt

Western belt

Protomyl. Mylonite Mylonite Ultramyl. Augen myl. Banded myl. Banded myl. pl amp, pll p12 amp, pl amp, pl amp, pll p12 GBPO Eigenvalue r = el/e2 Eigenvector V1 Orientation (~

Pl-pl 1.23 Amp-amp Amp-plt 1.41 Pl-pl 7 Amp-amp Amp-pl * 16 SPO P1 1.45 Eigenvalue Amp r = el/ea Amp t 1.81 Eigenvector V1 P1 - 14 Orientation (~ Amp Amp t 18 Aspect ratio P1 1.59 (median) Amp Amp ~ 2.67 Grain size-Ferret diameter Median (Ixm) P1 Amp Amp t Q1 (~m) P1 Amp Amp t Q3 (~m) P1 Amp Amp t Q3 - QJ (~m) P1 Amp Amp t Skewness "+ PI Amp Kurtosis * P1 Amp R2 P1 correlation Amp coef.

45

1.30 2.50 1.94 24 11 12 1.46 1.83 2.67 21 16 9 1.56 4.46 2.59

1.58 2.60 23 18 1.73 3.43 22 21 1.82 3.02

1.44 3.53 2.18 17 2 7 1.85 4.40 3.26 8 2 8 1.84 3.79 3.00

74 58 45 49 41 49 39 43 33 33 52 43 30 36 30 38 29 31 24 63 103 80 72 68 62 64 52 60 44 30 51 37 42 33 32 26 23 29 20 1.44 - 0.34 - 0.08 - 0.09 0.62 0.25 7.30 2.86 2.77 2.53 4.86 3.12 0.9151 0.9909 0.9979 0.9951 0.9768 0.9942

1.39 2.26 1.55 3 4 -5 1.86 2.89 1.96 1 8 - 1 1.68 2.36 2.28

1.24 1.96 1.50 - 3 - 2 1 1.38 2.21 1.76 2 - 2 9 1.50 2.12 1.84

43 119 59 30 62 34 63 211 102 32 150 68 1.15 0.02 6.99 2.43 0.9442 0.9899

92 94 50 57 61 35 147 142 74 90 81 39 0.01 0.36 2.98 2.74 0.9977 0.9883

*Positive values are oriented anticlockwisewith respect to the horizontal line. tin plagioclasedomains. *The R2 values, skewnessclose to zero, and kurtosis close to 3 signify well-fittedlognormaldistribution.

1.59 5 1.64 8 1.72

72 50 106 56

TEXTURES OF NATURALLY DEFORMED METAGABBROS

113

b) Amphibole ~.. ~ ~'-----2~--~--

'

~

'Plagioclase ~ East '

I.

Amphibole, Amphibolein plagioclasedomains

ttl

o

\

m

Protomylonite - E1

Amphibole - East

~

Mylonit~-E2

[~

Ultramylonite-E3

b

3

~I

~

Mylonite-E2

91 1 ~

I

I Ultramylonite-E3

A Plagioclase

--

5

0

0

50 100 150 Grain size - Ferret diameter (pm)

9 "~ A LU z~ /x

Protomylonite- E1 Mylonite- E2-pll Mylonite- E2-pl2 Ultramylonite- E3

~9 r

"

t

2E

Amphiboleinplagioclase domains

9 Augen mylonite- Wl o Banded mylonite- W2-pll o Banded mylonite- W2-p12

200

18 16 14

fillL. .

Plagioclase

llilJL -grained p l a g i o c l a s e

~8 6 4 2 0

Amphibole - West

- West

~21]il

-

dlllllllllllllllilLlUL J~nn200n~l][]nl250IJ]n~3000 50 100 150 Grain size - Ferret diameter (~m)

50

m

AugenmyloniteW1

[~

Bandedmylonite- W2

100 150 200 250 300 350 400

Grain size - Ferret diameter (p.m)

Fig. 11. Grain size evolution in the eastern (triangles) and western (circles) metagabbros. Degree of defomlation increases from dark to white colour. (a) Calculatedmedians and standard deviations. (b) Histograms showing the characteristic frequenciesof grain size for each deformational stage.

representing the size of the recrystallized matrix grains. Plagioclase grain size is substantially higher in the coarse-grained domain (W2-pll) of the banded mylonite (92 ~zm). The grain size is reduced to 72 ~m on average in the highstrain domain (W2-p12). The grain size distributions (Fig. 1 lb) in protomylonites of the eastern belt (El) show slightly bimodal distribution due to presence of large porphyroclasts surrounded by small recrystallized grains. A higher amount of large grains characterizes the western banded mylonite (W2) compared to the augen mylonite (W1) and eastern mylonites (El-E3). Lognormal distributions correlate with intensity of deformation in both belts.

Amphibole. The amphibole grain size is 49 p~m in the eastern metagabbro mylonite (E2) and 41 ~m in the ultramylonite (E3). However, the grain size spread is always lower than that of plagioclase in the same rock. The size of amphibole grains included within plagioclase domains decreases systematically with increasing deformation, from 49 ~m in the least deformed sample (El) to 39 ~m and 43 ~m in the mylonite (E2-pll and E2-p12, respectively) and 33 p~m in the ultramylonite (E3). The grain sizes are always slightly lower than those of amphibole matrix grains in the same thin section.

Amphibole grain size in the western metagabbro belt is highest in the sample including porphyroclasts (W1), reaching 119 Fm on average. Very high variance is typical of this mylonite, which is due to a relatively low number of new grains and a high number of large porphyroclasts. In the banded mylonites (W2), grain size decreases to 94 Fm on average. In the eastern belt, amphibole grain size distributions (Fig. 11b) are similar to those of plagioclase. They are characterized by well-developed lognormal distribution in both mylonite and ultramylonite stages. In contrast, the distribution of amphibole grain size in the western belt exhibits weaker fits of lognormal distribution with increasing deformation, which is attributed to a higher number of large grains of variable size (Table 4).

Shape preferred orientation PIagioclase. Recrystallized plagioclase grains in the protomylonite of the eastern belt (El) and mylonite (E2-pll) have aspect ratio (AR) of 1.59 and 1.56, respectively (Fig. 12a). Both aspect ratio (AR = 1.82) and SPO increase in the analysed high-strain plagioclase aggregate adjacent to the rigid amphibole porphyroclast in mylonite (E2-pl2). The ultramylonite (E3) exhibits further strengthening of the SPO and similar high aspect ratios (AR = 1.84).

L. BARATOUX ETAL.

114

Plagioclase grains in augen mylonites of the western belt (W l) have the strongest SPO but similar aspect ratios (AR = 1.68). The SPO decreases significantly in the coarse-grained plagioclase layers of banded mylonite (W2-pll), which also have the lowest aspect ratio of all analysed samples (AR = 1.50). Both aspect ratio and SPO further increase in the high-strain domain of banded mylonite (W2-p12).

4 1 plagioclase

3 o ._

I

F] West

o &

'

,<

Amphibole. Amphibole grains in the mylonite of

1.5!1

i

i 1.4

1

ll29

J

SPO

i 1.6

i

i 1.8

the eastern metagabbro belt (E2) have high aspect ratios (AR --- 4.46) but SPO is the lowest of all analysed amphibole domains (ram p = 1.83) (Fig. 12b). Aspect ratio in the ultramylonite (E3) is lower (AR = 3.79) but SPO is much h i g h e r (ram p --- 4.40) and interpreted as due to progressive rotation and elongation of most of the grains parallel to the mylonitic foliation. In the western belt, amphiboles are characterized by low aspect ratios in the augen ( A R - 2.36; W1) and banded mylonites (AR = 2.12; W2). The SPO also decreases from augen (ramp = 2.89) to banded mylonite (ram p = 2.21).

i

11 amphibole 10 East A 9 Wesl 8 .~ 7 n" 6

•[_.•

~5

24

3

Grain boundary preferred orientation

2 1

1

1.5

2

2.5

3 SPO

3,5

4

4.5

5

c 0 ~

Amphibole-amphibole

I A -------------

~

0

____.----~ A Amphibole-plagioclase in plagioclasedomains

~ &

pll

o p12

~...e

pllA/

A

I plagioclase-plagioclase

Apl2

1.5

2.5 GBPO

315

Fig. 12. Plot of grain shape preferred orientation (SPO) vs. aspect ratio for plagioclase (a) and for amphibole (b), and diagram of grain boundary preferred orientation (GBPO) of amphibole and plagioclase like-like and unlike boundaries (c). (a) and (b) Results are summarized in a boxplot-type diagram of aspect ratios vs. SPO expressed as eigenvalue ratio of bulk matrix of inertia of individual phases. Individual boxes show median, 1st and 3rd quartile of aspect ratio. The whiskers represent statistical estimate of the data range. Outliers are not plotted. Boxes of the western metagabbros are shaded. (c) GBPO is expressed as eigenvalue ratio of bulk matrix of inertia of individual phase boundaries. Triangleseastern belt; circles-western belt; Degree of deformation increases from dark to white colour.

The degree of grain boundary preferred orientation (GBPO, Fig. 12c) of plagioclase-plagioclase boundaries (rpl-p0 increases from 1.23 for the protomylonite (E 1) to 1.44 for the ultramylonite (E3). A maximum value (rpl_pl = 1.58) is achieved in the high-strain domain adjacent to the amphibole porphyroclast in mylonite (E2-p12). The orientation of the maximum eigenvector (VI) progressively deviates from the foliation direction with increasing strain (Table 4), which indicates an increasing effect of noncoaxial deformation. The r p l _ p l in the western metagabbro belt is quite low in the augen mylonite (W1; rpl_pl = 1.39) and further decreases in the coarse-grained matrix of banded mylonite (W2-pll; rpl_pl--1.24). The rpl_pl increases again in the localized sheared zone (W2-p12). Orientations of the maximum eigenvector V~ are more or less parallel to the foliation indicating a predominantly coaxial deformation (Table 4). The degree of preferred orientation of amphibole- plagioclase boundaries (ramp_pl) shows a similar evolution but higher values of ramp-p1 compared to rp]_p] in both belts. The degree of amphibole-amphibole GBPO (ramp-amp) in the eastern belt reaches the highest values, which is due to the aggregate distribution and elongate shape of grains. The ramp_am p increases from 2.5 in the mylonite (E2) to 3.53 in the ultramylonite (E3). The

TEXTURES OF NATURALLY DEFORMED METAGABBROS maximum eigenvector V1 is only slightly inclined with respect to the foliation, suggesting minor component of non-coaxial deformation. In the western belt, ramp_am p weakens from 2.26 in the augen mylonite (W1) to 1.96 in the banded mylonite (W2), being still stronger than rpl-pl in similar rock types. The eigenvector V~ is parallel to the foliation in both types of mylonites.

Crystallographic preferred orientation Five thin sections, covering all the deformation stages, were analysed using the EBSD technique. In the eastern mylonite (E2) and the western banded mylonite (W2), two plagioclase domains were investigated. Besides pole figures of [100] directions and poles to (010) and (001) planes, inverse pole figures, calculated parallel to stretching lineation and pole to foliation, are used in the eastern belt in order to identify orientation maxima, which could correspond to less common slip systems reported in the literature. The inverse pole figures may be employed because the orientations of lineation and foliation follow strong SPO and CPO of amphibole. For the description of inverse pole figures, only the slip systems previously described in the literature will be mentioned (for summary see e.g. Kruse et al. 2001 ; Sttinitz et al. 2003). The I parameters representing the intensity of CPO are shown in Table 5. Plagioclase

Within the protomylonite of the eastern belt (El), the plagioclase CPO is not random (Figs 13 and 14a). In the inverse pole figure, the lineations are distributed subparallel to the horizontal plane with a maximum around [i 10]. Poles to foliation are arranged in a N - S girdle with maxima around (001) and a* (100). The (001) [1 i0] slip system (Olsen & Kohlstedt 1984, 1985) could

Table 5. Intensity I of crystallographic preferred orientation Hornblende

Plagioclase

(100) [001] {110} [1001 (010) (001) E1 E2-pl 1 E2-p12 E3

1.51t

1.926 0.897 0.116 0.095 0.108 2.906 1.591 1.454 0.083

0.666 0.109 0.347 0.033

0.319 0.100 0.171 0.178

Wl 1.422 1.854 0.846 0.111 0.074 0.020 W2-pll 0.948 1.606 0.653 0.457 0.552 0.948 W2-p12 0.590 0.299 0.741

115

thus be possibly active. The orientation of porphyroclasts is different from that of recrystallized grains, suggesting that the recrystallization was not host-controlled (e.g. Ji & Mainprice 1990; Kruse et al. 2001). The mylonite (E2-pll and E2-p12) shows weak CPO in both plagioclase domains. A weak maximum of lineations is situated around [110] in the E2-pll _domain, and some clusters of lineation around [1_12] directions and poles to foliation close to (201) can be observed in the high-strain domain in sample E2-p12. These orie_ntationsmay indicate the possible activation of (201)1/21112] (Marshall & McLaren 1977a, b) in the high-strain fine-grained domain adjacent to the hornblende porphyroclast (Fig. 4d). The ultramylonite CPO (E3) displays maximum for poles to foliation clustering near the (021) and maximum for lineations near the [112] direction. Such CPO would suggest activation of (021)1/21112] slip system (Olsen & Kohlstedt 1984; Montardi & Mainprice 1987). The I parameters attain lower values in more deformed stages for [100], (010), and (001), respectively. The CPO of the augen mylonite in the western belt (W1) (Fig. 15) is very weak and no slip systems could be extracted even by examining the inverse pole figures, which are not shown here. Within banded mylonite (W2-pll and W2p12), both coarse-grained (pll) and fine-grained (p12) plagioclase domains exhibit strong maxima of [100] directions close to the lineation. Poles to (010) and (001) planes are distributed close to the pole of the foliation, which is defined by compositional layering. The domain W2-p12 represents late shear zones marked by important grain refinement (Fig. 7c). The 1 parameters are higher in the banded mylonite (W2) compared to the augen mylonite (W1), indicating strengthening of plagioclase CPO. It is likely that this late shearing was not entirely coaxial with respect to the main flow represented by coarse-grained aggregate fabrics (W2-pl 1). Therefore, the difference of (010) and (001) maxima in pole figures could result from slightly oblique orientation of late shear plane with respect to dominant compositional layering. The shear zone (W2-p12) is marked by an inclination of [100] maximum of about 10 ~ with respect to the lineation, which could be attributed to non-coaxial deformation (Ji et al. 1988). Activity of slip systems (001)[100] (Marshall & McLaren 1977a, b) and (010)[100] (Montardi & Mainprice 1987) may be inferred from such orientation patterns. Amphibole

Small recrystallized hornblendes adjacent to their host grain in the eastern protomylonite

L. BARATOUX ETAL.

116

Plagioclase - eastern belt (010)

[1001

.,7,

~_s

......

..

.

kQ...'........,....._~...:/r

.

.

.

.

.,.., "E

~,,.

.

....' .:.-,~.~ :9 M D = 3.36

.e-...........,,

(001)

:~':.

.'

..

-.~ '...:

,'

9 M D = 6.24

9 M D = 3.63

9 M D = 3.28

9 MD = 2 , 8 3

9 MD = 2.54

9 M D = 3.33

9 MD = 3.99

9 M D = 3.33

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

:? ,::..,...

--

-.; ~&~-41 Lower hemisphere

9 M D = 2.71

. --.,.."" :'. ".,. ......., ,,~ ..

f" """~.-. ):: ....

9 M D = 2.67

9 M D = 2.63

Fig. 13. Point and contour pole figures of plagioclase CPO in the eastern belt (lower hemisphere, equal area projections). Diagrams are contom'ed as multiples (0.5, 1.0, 1.5, 2.0, 2.5 .... x ) of uniform distribution. MD is maximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineation is trending E-W. (El) show progressive reorientation of the crystallographic planes (100) towards the foliation direction and c-axes [001] towards the lineation (Fig. t6a). Cleavage planes {110} cluster parallel to the foliation as well. The asymmetry of the CPO fabric is consistent with the dextral shear zone cross-cutting the porphyroclast under the angle of ~ 2 0 ~ with respect to the mylonitic foliation. The CPO of hornblende in the ultramylonite (E3) displays a strong maximum of poles to the (100) and {110} planes (Fig. 16b) clustering normal to the foliation ( I - - 2 . 9 for (100)). The c-axes trend in the direction of lineation. Such CPO patterns suggest that the (100)[001] slip system could be activated. Inclination of data with respect to the mylonitic foliation due to dextral shear is present. Similar fabric asymmetry, interpreted as the result of a non-coaxial strain, has been described in eclogite facies glaucophanites (Zucali et al. 2002). Hornblende CPO in the augen mylonite (W 1) from the western belt form a cluster of (100)

planes subparallel to the mylonitic foliation and c-axes parallel to the lineation. However, a few grains occur with a different orientation (Fig. 16c). In the banded mylonite (W2), the (100) planes and c-axes reveal similar orientation as in the augen mylonite but the maxima are less pronounced (Fig. 16d). The I parameters are slightly lower in the banded mylonite (W2) compared to the augen mylonite (W1). This CPO pattern is consistent with activation of the (100)[001] slip system. No asymmetry with respect to the shear plane occurs. Discussion

and

conclusions

Recrystallization and deformation mechanisms of plagioclase The early deformation stage in plagioclase in the eastern amphibolite facies metagabbro belt is marked by undulatory extinction, twinning, and fracturing, which are typical for low temperature

TEXTURES OF NATURALLY DEFORMED METAGABBROS

117

Plagioclase - eastern belt Poles to foliations a-coo) Lineations a-(~oo) a*(100) a*(100)

a) upper hemisphere

=219

Protomylonite

{.-'..:: "_~.'.~.~

E1 ~

M

D a*(100)

= 3.91 9

a'(100)

N=160

a*(100)

D = 3.60, 9

a*(100)

N

Mylonite

:'.~-...:. , .....

E2-pl 1

~!-'".:=o .; D=3,53 9

D=2.83 9

a*(lOO)

a*(lO0) ,

Mylonite

N=I

",

o-

a*(lOO)

a*(lOO)

= 160

~

0o

E2-p12 MD = 2,95 9 a'(~00)

D = 3.20 9 a'(10o)

a'0 00)

a'(100)

= 200

N

Ultramylonite

E3 D=4,

b)

~

D=4.

IPF Reference frame 1 2 ~ ~ 0 _ Poles to planes(hkl) /o/o21O1 ~'1~1~10,~

Directions [uvw] ~

9

O

l

O

O l O ~ o i -

o

Fig. 14. (a) Inverse pole figures of plagioclase CPO in the eastern belt (upper hemisphere, equal area projections). Projections of lineations are presented as point and contour diagrams in the left, projections of poles to foliations are in the fight side of the figure. N is the number of plotted points. Diagrams are contoured as multiples (0.5, 1.0, 1.5, 2.0, 2.5.... • ) of uniform distribution. MD is maximum density of data in the contour diagrams. Large circle and square in the protomylonite represent two porpbyroclasts. (b) The crystallographic reference frame.

and/or high differential stress (Tullis & Yund 1987, 1992; Rosenberg & Sttinitz 2003). Small grains concentrated in these fractures exhibit typical features of dynamic recrystallization such as bulging and subgrain rotation. However, some of them might have developed by crushing of the host grain or by heterogeneous nucleation (StiJnitz et al. 2003; Rosenberg & Sttinitz 2003). Twin-free grains adjacent to twinned host grains (see Fig. 3b), compositional

resemblance of recrystallized grains to host grains, and strong CPO (Lister et al. 1978) may indicate a recrystallization mechanism by subgrain rotation. This recrystallization mechanism is also supported by small average grain sizes, small aspect ratios and weak SPO of new grains. A fairly small degree of GBPO for plagioclase-plagioclase boundaries is in agreement with subequant shapes of equidimensional new grains, thus favouring the hypothesis of subgrain

118

L. BARATOUX E T A L . P l a g i o c l a s e - w e s t e r n belt

(010)

N = 150

9 M D = 3.64

9 MD = 5.23

Lower hemisphere

9 MD = 6.00

(001)

9 M D = 3.93

9 MD = 3 . 1 6

9 MD = 4.39

9 M D = 5.41

9 M D = 4.38

9 MD = 4.77

Fig. 15. Point and contour pole figures of plagioclase CPO in the western belt (lower hemisphere, equal area projections). Diagrams are contoured as multiples (0.5, 1.0, 1.5, 2.0, 2.5.... • ) of uniform distribution. MD is maximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineation is trending E-W. rotation recrystallization mechanisms. The CPO is strong and indicates a possible activity of the (001)[110] slip system. The mylonitic stage is characterized by subequant shapes of recrystallized plagioclase grains (low aspect ratios) with straight boundaries, weak shape preferred orientation and slightly higher degree of GBPO compared to the protomylonite. Recrystallized grains show increasing grain sizes compared to protomylonitic stage. Compositional zoning in plagioclase indicates syndeformational growth of matrix grains (Sodre Borges & White 1980). The plagioclase CPO in the mylonite is fairly weak. These observations suggest that grain boundary sliding (Boullier & Gu6guen 1975; Lapworth et al. 2002), accompanied by some component of crystal plastic deformation in plagioclase, was the dominant deformation mechanism. Mixing of amphibole and ptagioclase grains could be another argument for the presence of grain boundary sliding. Rosenberg & Sttinitz (2003) suggested for the syntectonically cooled and deformed Bergell tonalite that mixing of plagioclase and biotite in a fine-grained matrix together with weakening of CPO implied a mechanism of diffusion-accommodated grain boundary sliding. A switch of deformation mechanism from dislocation creep to grain size sensitive (GSS) flow

due to strain softening and grain size reduction is well known from quartzo-feldspathic rocks (e.g. Walker et al. 1990; Tullis & Yund 1991). Very small grain size is required for activation of GSS deformation: 'less than 10 ~m' by Boullier & Gu~guen (1975), 2-16 &m by Tullis & Yund (1991), and 3-30 p~m by Stfinitz & Fitz Gerald (1993). However, some studies suggest grain boundary sliding deformation for grain sizes of 100-150 I-~m for plagioclase (Jensen & Starkey 1985), 24-41 I~m for mixed plagioclase-hornblende layers (Kruse & Stfinitz 1999), and 100-250 p,m for plagioclase (Lapworth et al. 2002). In our samples with weak or random CPO, the grain size varies between 30 and 150 txm. The occurrence of grain-boundary diffusion creep for such grain size may be explained by very low strain rates (Fliervoet et al. 1999; Lapworth et al. 2002). Fine-grained plagioclases from high-strain domains adjacent to amphibole porphyroclasts and matrix plagioclase grains in ultramylonites show a decrease in average grain size coupled with increased aspect ratios, strong SPO and GBPO. The CPO is strong for the ultramylon_ite sample, suggesting activation of a (021) 1/2 [ 112] slip system. These features may indicate that at very high strains the grain size sensitive flow becomes less important in plagioclase.

119

TEXTURES OF NATURALLY DEFORMED METAGABBROS

[a] (lOO) a)

[c] [001]

Hornblende Lower [a] {110} hemisphere (100)

Protomylonite - E 1

b)

[c] [001]

{110}

Ultramyionite - E3

-

a,~,~

-.

.~-

|

9

.

.

."

N=98

N = 209

9 MD = 7,69

c)

9MD = 9.57

9 M D = 4.61

Augen mylonite - W l

N=

m MD = 7.93

9

9 MD = 15.98

d)

9 MD = 12.08

Banded mylonite - W2

N=

105

9 MD = 9,94

=MD

= 5.11

9 MD = 5.55

=MD

= 5.15

107

= MD = 9,46

9 M D = 4.91

Fig. 16. Point and contour pole figures of amphibole CPO in (a and b) the eastern belt and (c and d) the western belt (lower hemisphere, equal area projections). Contours are counted as multiples (1, 2, 3, 4 .... • ) of uniform distribution. MD is maximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineation is trending E-W. The CPO of porphyroclast is marked by a circle.

A possible switch from diffusion-dominated flow in the mylonite to dislocation creep in the ultramylonite may be due to a decrease in temperature of deformation as described from felsic granulites by Martelat et al. (1999). The switch from the GSS to the GSI regime for decreased grain size is possible and implies an increase of strain rate and/or a decrease of temperature in the shear zone (Handy 1990). A similar transition from diffusion-dominated creep in plagioclase aggregates in augen mylonites to dislocation creep in banded ultramylonites was reported by Schulmann et al. (1996) from deformed metagranitoids.

Dominant (010)[001] slip was identified in plagioclase from both CPO patterns and/or TEM observations (e.g. Olsen & Kohlstedt 1984, 1985; Montardi & Mainprice 1987; Ji & Mainprice 1988, 1990). However, the observed CPO patterns for amphibolite facies metagabbros are not compatible with activation of the classic slip system. Various other slip systems, which may be activated in plagioclase also, were described by Marshall & McLaren (1977a, b), Olsen & Kohlstedt (1984, 1985), and Montardi & Mainprice (1987). In addition, St(initz et al. (2003) concluded that during experimental deformation of plagioclase several slip systems

120

L. BARATOUX ET AL.

are contemporaneously activated. The presence of several maxima in inverse pole figures of our samples may indicate that several slip systems probably operated simultaneously and their unequivocal identification from CPO patterns is therefore difficult. In the western augen mylonites, plagioclase recrystallizes dynamically by subgrain rotation. Aggregates of plagioclase matrix grains show microstrnctures indicative for dislocation creep, such as subgrain formation, strong SPO, high aspect ratios and relatively strong GBPO of plagioclase boundaries. Large differences in chemical compositions between porphyroclast and matrix may imply that some of the grains originated by heterogeneous nucleation (Rosenberg & Sttinitz 2003). Surprisingly, the CPO of recrystallized grains is very weak, which may suggest that some GSS process took place at this stage of deformation. In the banded mylonites, the plagioclase matrix grains show strongly migrated boundaries, and lower SPO, aspect ratios and GBPO compared to the augen mylonite. Plagioclase is characterized by a relatively high average grain size and a common growth-related, chemically zoned composition. The plagioclase cores are overgrown by a new plagioclase with a more calcic composition at upper amphibolite to granulite facies conditions. A similar mechanism was numerically modelled by Jessell et al. (2003) showing that some original grain cores are directly in contact with rims of new grains due to preferential growth of grains at the expense of others. This sharp compositional difference must necessarily introduce a strong chemical potential between these two grains. Serrated boundaries indicate a mechanism of grain boundary migration. A process called chemically induced grain boundary migration (CIGM) was introduced by Hillert & Purdy (1978) and in the case of naturally deformed amphiboles described by Cumbest et al. (1989a). Such a process requires a strong chemical potential and is driven by a reduction in the free chemical energy associated with the chemical change (Yund & Tullis 1991). However, Yund & Tullis (1991) observed important compositional changes in experimentally deformed plagioclase, which were not assumed to be the driving force for associated dynamic recrystallization. The CPO in banded mylonites is relatively strong and (001)[100] and (010)[100] slip may be inferred from the pole figures. Martelat et al. (1999) reported a strong CPO in plagioclase deformed by high-temperature diffusion creep in coarse-grained felsic granulites. We propose, however, that a difference in chemical potential

could have been one of the main driving forces for grain boundary migration in the banded mylonites. Fine-grained shear zones, cross-cutting the coarse-grained plagioclase bands, developed later at somewhat lower temperatures. This is supported by the mechanical twins crosscutting chemically zoned plagioclases from these shear zones. Higher aspect ratio, stronger SPO and GBPO, and strong CPO suggest dominant activity of dislocation creep with activation of (001)[100] slip in these domains. Deformation mechanisms of amphibole

Microstructural evidence for cataclastic flow was presented by Nyman et al. (1992) from amphibolites that were previously interpreted to have been deformed plastically. Amphiboles are relatively resistant to crystal plastic deformation even at high metamorphic grades (Brodie & Rutter 1985). The dominant slip system (100)[001] in amphibole was inferred from both experimental and natural studies (Dollinger & Blacic 1975; Rooney et al. 1975; Cumbest et al. 1989b; Hacker & Christie 1990; Skrotzki 1990). The (hk0)[001] slip system was determined by Biermann & Van Roermund (1983) and Morrison-Smith (1976). The (010)[100] and (001)[100] slip systems were observed by Morrison-Smith (1976) in experimental studies but not in nature. All these studies concluded that the only possible active Burgers vector operating in amphiboles WaSoin [001], which is also the shortest one (5.299 A). [100] and [010] Burgers vectors (9.885 A and 18.169 A, respectively) were thought to be too long to be activated (Rooney et al. 1970). Hornblende microstructures from the eastern metagabbro belt (protomylonite and mylonite) exhibit twinning on (100) planes (Rooney et al. 1975), kinking, and strong undulatory extinction related to bending/twins of the crystalline lattice. Microshear zones transect large porphyroclasts into bookshelf-like segments and these segments subsequently rotate with their c-axes parallel to the shear direction. This is consistent with pole figures showing CPOs of small grains that are different from the host porphyroclast. In YZ sections, the grains have lozenge shapes, with their grain boundaries parallel to the {110} cleavage. The SPO and GBPO are relatively weak for these strongly elongate grains when compared to grains from more advanced stages of deformation. This is interpreted to be the result of development of new grains along intracrystalline microshear zones and along margins of porphyroclasts.

TEXTURES OF NATURALLY DEFORMED METAGABBROS Some observations are consistent with brittle deformation of amphibole. For instance, hornblende host grains are often rimmed by 'subgrains' forming core-mantle-like structures. Grain boundaries between new grains and porphyroclasts are very sharp, indicating that new grains formed by fracturing rather than by subgrain rotation (Nyman et al. 1992). Interstitial plagioclase decorate grain boundaries between small recrystallized grains demonstrating opening of spaces between amphiboles. However, the operating deformation mechanism in hornblende cannot be unequivocally determined from optical observations and EBSD data alone. The strong CPO patterns may be interpreted by (100)[001] slip with dislocation creep being the major operative deformation mechanism, by translation gliding along {110} planes indicating dominant microfracturing or by mechanical rotation of anisotropic grains. The strong asymmetry of CPO reflects the development of antithetic microshear zones operating during bookshelf-like rotations associated with dextral shear. Hornblende crystals in ultramylonites are characterized by higher aspect ratios, more uniform grain sizes, and stronger SPO and GBPOs compared to amphiboles from proto mylonites and mylonites. Strong CPO of (100) and {110} cleavage planes subparallel to the foliation and c-axes subparallel to the lineation are consistent with activation of the (100)[001 ] and {110}[001] slip systems. As mentioned above, a strong CPO of amphibole may also be produced by rigid body rotation. The most important observation in the western augen mylonitic hornblendes is that the original porphyroclasts and new grains have different chemical compositions (magnesio-hornblende and tschermakite with higher Na and K content, respectively). Chemical differences between old and new grains suggest that new grains originated by heterogeneous nucleation under higher metamorphic conditions (e.g. Binns 1965; Laird & Albee 1981), which may be chemically and/ or deformationally induced (Berger & Sttinitz 1996). The porphyroclasts show high aspect ratios and SPO, GBPO, and CPO maxima oriented parallel to the mylonitic foliation. A rigid body rotation of large porphyroclasts in the weaker plagioclase matrix is the most likely mechanism to explain the strong preferred orientation of these grains. New grains have significantly lower aspect ratios, strong SPO, CPO and high degrees of GBPO of like-like boundaries. The orientation of CPO maxima is consistent with the possible activation of (100)[001] slip.

121

In the banded mylonites, primary hornblende porphyroclasts are missing; however, slight chemical zoning within large and smaller new grains indicates their syndeformational growth. Amphiboles, in monomineral hornblende bands, show straight boundaries decorated with small amounts of tiny interstitial plagioclase. Amphibole grains show no signs of internal deformation and have straight to lobate grain boundaries, which is indicative of high degrees of textural equilibration (Brodie & Rutter 1985). The grain sizes and the degree of SPO and GBPO decrease with respect to less deformed rocks (augen mylonite). However, the CPO is strong and very similar to that from less deformed augen mylonite, indicating possible activation of (100) [001 ] slip. Hornblende grains, which occur in plagioclase aggregates, are marked by significantly smaller grain sizes compared to host matrix crystals (Fig. 7c). They are often located at highenergy surfaces or unstable triple points, and show globular shapes typical for a very different structure of host and inclusion grain. In some places, where the minor phase becomes abundant, it is forced to occupy less favourable planar interfaces (Vernon 1976). All these criteria suggest that these grains developed by heterogeneous nucleation, namely in a high surface energy plagioclase aggregate (Spry 1969; Dallain et al. 1999). The mechanism of syndeformational growth is more pronounced in hornblende layers. This is supported by shapes of grain size distribution histograms that show important numbers of large grains. Possible mechanisms to explain strong CPO in the studied rocks is a nucleation and growth of new hornblende grains in a non-lithostatic stress field as was proposed for pyroxenes by Helmstaedt et al. (1972) or Mauler et al. (2001).

Implications for the theology of a two-phase amphibole-plagioclase system

Experimental deformation of amphibolite and granulite facies natural samples of metabasites (Wilks & Carter 1990) reveals that the rheology of such polyphase systems depends on temperature, rock composition, deformation mechanisms, water content, and to a minor extent, pressure. The relationship between these factors and differential stress and strain is, however, complex and Wilks & Carter (1990) concluded that it is necessary to estimate the contribution of each phase to the creep rate of the bulk rock in order to establish a convenient flow law. Rosenberg & Sttinitz (2003) proposed that interconnected monomineralic networks of

122

L. BARATOUX ET AL.

recrystallized plagioclase in middle or deep crustal metabasic rocks are rare and restricted to rocks with very high plagioclase modal abundances such as anorthosites (Ji et al. 1988). However, as mentioned by other authors (e.g. Brodie & Rutter 1985), banded metabasic mylonites are common, which is confirmed by this work. We show that plagioclase may easily form interconnected weak layer networks in hornblende gabbros under upper amphibolite facies conditions. Our microstructural study of an amphibolite facies metagabbro (650 _+ 50 ~ shows that the load-bearing framework structure (Handy 1990, 1994) is restricted to the lowest deformation intensities. The non-deformed metagabbro shows coarse-grained ophitic structure composed of randomly distributed hornblende and plagioclase, where amphibole grains are only locally in contact. With ongoing strain, the deformation is mostly concentrated in the plagioclase, which is at the transient region between brittle and plastic behaviour, leading to development of a fine-grained matrix (Tullis & Yund 1987). Amphibole grains showing high internal strain and local fracturing behave as rigid bodies surrounded by interconnected layers of plagioclase grains. Such microstructures may be interpreted as an interconnected weak layer structure (IWL) with a high viscosity contrast between rigid clasts of amphibole and weaker, finegrained plagioclase layers (Handy et al. 1999). In the amphibolite facies metagabbroic mylohires, the monomineralic hornblende layers are observed, while plagioclase-rich layers show almost perfect mixing with hornblende. The grain size of plagioclase and amphibole in the plagioclase-rich matrix areas is fairly similar, the latter showing slightly higher elongation. The phase distribution of plagioclase-hornblende mixture, the absence of CPO in plagioclase, and its aspect ratio suggest that the dominant mechanism is granular flow (grain boundary sliding). Based on this interpretation, we suggest that the phase mixing is probably a mechanical process. The microstructures and CPO of amphibole forming monomineralic layers indicate either dislocation creep or cataclastic flow. The absence of boudinage and progressive mixing of plagioclase and amphibole suggest that the diffusion-dominated flow process operating in plagioclase aggregates is mechanically as efficient as dislocation or cataclastic flow in the hornblende layers. The final structure resembles the interconnected weak layer structure with low viscosity contrast (Handy 1994). A switch in deformation mechanism from dislocation creep towards a grain size

sensitive process is thought to be responsible for the convergence of mechanical properties of amphibole and plagioclase in the mylonite, resulting in a drop of bulk rock strength (Etheridge & Wilkie 1979; Kirby 1985; Rutter & Brodie 1988). Two deformation stages were observed in the upper amphibolite facies (750 -t- 50 ~ metagabbros: (1) augen mylonite with locally preserved porphyroclasts of both plagioclase and hornblende; and (2) banded mylonites with completely recrystallized amphibole and plagioclase, each arranged in monomineralic layers. The initial stages of deformation are characterized by tectonic grain size reduction of plagioclase while hornblendes represent strong objects floating in the weak plagioclase matrix. The deformation of the metagabbro is interpreted to have occurred via dislocation creep accompanied by diffusion mass transfer mechanisms, responsible for moderate mixing of plagioclase and amphibole. The banded fabric, only developed at high strains, can be defined as an interconnected weak layer structure with low viscosity contrast (Handy et al. 1999). The layered structure shows that the strengths of amphibole monornineralic aggregates and plagioclase-rich bands are similar, suggesting convergence of rheologies of both minerals at high strains (Jordan 1988; Handy 1994). In conclusion, amphibolite and upper amphibolite facies metagabbroic mylonites are characterized by layered low-viscosity IWL structures. This indicates that at high strains this banded structure in metagabbros forms a so-called steady-state foliation (Means 1990). Mechanical mixing of phases is more important in lower temperature (eastern belt) than in higher temperature (western belt) amphibolite facies mylonites. The bulk strength of amphibolite and upper amphibolite facies mylonitic metagabbros is controlled by an equal contribution of both rock-forming minerals showing contrasting but equally efficient deformation mechanisms.

We are indebted to D. Mainprice for his help with the EBSD analyses. Fruitful discussionswith F. Holub, D. J. Prior, H. Stiinitz and J. Wheeler are gratefully acknowledged. We thank P. T~)cov~,J. Haloda, P. Grandjeanand P. Capiez for the help with microprobeand bulk rock analyses. K. Brodie, H. Van Roermund and D. Gapais are thanked for thorough reviews, which improved significantly the original manuscript. The project was funded by grants of Czech National Grant Agency No. 42-201204 to K.S. and 42-201-318 to P. Stipskfi,by Czech Geological Service assignmentNo. 6327 to P. Mixa, and by a PhD financialsupport attributedby the French government to L.B.

TEXTURES OF NATURALLY DEFORMED METAGABBROS

References BERGER, A. & STI~INITZ,H. 1996. Deformation mechanisms and reaction of hornblende: examples from the Bergell tonalite (central Alps). Tectonophysics, 257, 149-174. BIERMANN, C. & VAN ROERMUND, H. L. M. 1983. Defect structures in naturally deformed clinoamphiboles; a TEM study. Tectonophysics, 95, 267-278. BINNS, R. A. 1965. The mineralogy of metamorphosed basic rocks from the Williama Complex, Broken Hill district, New South Wales, Part I, Hornblendes. Mineralogical Magazine, 35, 306-326. BOULLtER, A. M. & GUEGUEN,Y. 1975. SP-mylonites; origin of some mylonites by superplastic flow. Contributions to Mineralogy and Petrology, 50, 93-104. BROD~E, K. H. 1981. Variation in amphibole and plagioclase composition with deformation. Tectonophysics, 78, 389-402. BRODIE, K. H. & RUTTER, E. H. 1985. On the relationship between deformation and metamorphism, with special reference to the behavior of basic rocks. In: THOMPSON, A. B. & RUBIE, D. C. (eds) Meta-

morphic Reactions; Kinetics, Textures, and Deformation. Springer-Verlag, New York, MY, United States, 138-179. CARVER, N. L. & TSENN, M. C. 1987. Flow properties of continental lithosphere. Tectonophysics, 136, 27-63. CUMBEST, R. J., DRURY, M. R., VAN ROERMUND, H. L. M. & S1MPSON,C. 1989a. Dynamic recrystallization and chemical evolution of clinoamphibole from Senja, Norway. Contributions to Mineralogy and Petrology, 101, 339-349. CUMBEST, R. J., Van ROERMUND, H. L. M., DRURY, M. R. & SIMPSON, C 1989b. Burgers vector determination in clinoamphibole by computer simulation. American Mineralogist, 74, 586-592. DALLAIN, C., SCHULMANN, K. & LEDRU, P. 1999. Textural evolution in the transition from subsolidus annealing to melting process, Velay Dome, French Massif Central. Journal of Metamorphic Geology, 17, 61-74. DOLLINGER, G. & BLACIC, J. D. 1975. Deformation mechanisms in experimentally and naturally deformed amphiboles. Earth and Planetary Science Letters, 26, 409-416. ETHERIDGE, M. A. & WILKIE, J. C. 1979. Grainsize reduction, grain boundary sliding and the flow strength of mylonites. Tectonophysics, 58, 159-178. FITZ GERALD, J. D. & STI)NITZ, H. 1993. Deformation of granitoids at low metamorphic grade; I, Reactions and grain size reduction. Tectonophysics, 221, 269-297. FLIERVOET, T. F., DRURY, M. R. & CHOPRA, P. N. 1999. Crystallographic preferred orientations and misorientations in some olivine rocks deformed by diffusion or dislocation creep. Tectonophysics, 303, 1-27. FREER, R. 1981. Diffusion in silicate minerals and glasses: a data digest and guide to the literature.

123

Contributions to Mineralogy and Petrology, 76, 440-454. GAPAIS, D. 1989. Shear structures within deformed granites; mechanical and thermal indicators. Geology, 17, 1144-1147. HACKER, B. R. & CHRISTIE, J. M. 1990. Brittle/ductile and plastic/cataclastic transitions in experimentally deformed and metamorphosed amphibolite. In: DUBA, A. G., DURHAM, W. B., HANDIN, J. W. & WANG HERBERT, F. (eds) The Brittle-Ductile Transition in Rocks. American Geophysical Union, Washington DC, Geophysical Monograph, 56, 127-147. HANDY, M. R. 1990. The solid-state flow of polymineralic rocks. Journal of Geophysical Research, B, 95, 8647 - 8661. HANDY, M. R. 1994. Flow laws for rocks containing two non-linear viscous phases; a phenomenological approach. Journal of Structural Geology, 16, 287-301. HANDY, M. R_, WISSING, S. B. & STREIT, L. E. 1999. Frictional-viscous flow in mylonite with varied bimineralic composition and its effect on lithospheric strength. Tectonophysics, 303, 175-191. HELMSTAEDT, H., ANDERSON,O. L. & GAVASCI,A. T. 1972. Petrofabric studies of eclogite, SpinelWebsterite, and Spinel-Lherzolite xenoliths from kimberlite-bearing breccia pipes in Southeastern Utah and Northeastern Arizona. Journal of Geophysical Research, 77, 4350-4365. HILLERT, M. & PURDY, G. R. 1978. Chemically induced grain boundary migration. Acta Metallica, 26, 333-340. HOLLAND, T. J. B. & BLUNDY, J. 1994. Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry. Contributions to Mineralogy and Petrology, 116, 433 -447. JENSEN, L. N. & STARKLY,J. 1985. Plagioclase microfabrics in a ductile shear zone from the Jotun Nappe, Norway. Journal of Structural Geology, 7, 527-539. J~SSELL, M. W., KOSTEN~:O,O. & JAMTVEIT, B. 2003. The preservation potential of microstructures during static grain growth. Journal of Metamorphic Geology, 21, 481-491. JI, S. & MAINPRICE, D. 1988. Natural deformation fabrics of plagioclase; implications for slip systems and seismic anisotropy. Tectonophysics, 147, 145-163. JI, S. & MAINPRICE, D. 1990. Recrystallization and fabric development in plagioclase. Journal of Geology, 98, 65-79. JI, S., MAINPRICE, D. & BOUDIER, F. 1988. Sense of shear in high-temperature movement zones from the fabric asymmetry of plagioclase feldspars. Journal of Structural Geology, 10, 73-81. JORDAN, P. G. 1988. The rheology of polymineralic rocks - an approach. Geologische Rundschau, 77, 285-294. KIRBY, S. H. 1985. Rock mechanics observations pertinent to the theology of the continental lithosphere and the localization of strain along shear zones. Tectonophysics, 119, 1-27.

124

L. BARATOUX ETAL.

KNIPE, R. J. 1989. Deformation mechanisms, recognition from natural tectonites. Journal of Structural Geology, 11, 127-146. KNIPE, R. J. & MCCAIG, A. M. 1994. Microstructural and microchemical consequences of fluid flow in deforming rocks. In: PARNELL, J. (ed) Geofluids; Origin, Migration and Evolution of Fluids in Sedimentary Basins. Geological Society, London, Special Publications, 78, 99-111. KRETZ, R. 1983. Symbols for rock-forming minerals. American Mineralogist, 68, 277-279. KRONER, A., STIPSKA, P., SCHULMANN, K. & JAECKEL, P. 2000. Chronological constraints on the pre-Variscan evolution of the northeastern margin of the Bohemian Massif, Czech Republic. In" FRANKE, W., HAAK, V., ONCKEN, O. & TANNER, D. (eds) Orogenic Processes; Quantification and Modelling in the Variscan Belt. Geological Society, London, Special Publications, 179, 175-197. KRUHL, J. H. 1987. Preferred lattice orientation of plagioclase from amphibolite and greenschist facies rocks near the Insubric line (Western Alps). Tectonophysics, 135, 233-242. KRUSE, R. & STf0NITZ, H. 1999. Deformation mechanisms and phase distribution in mafic hightemperature mylonites from the Jotun Nappe, southern Norway. Tectonophysics, 303, 223-249. KRUSE, R., STONITZ, H. & KUNZE, K. 2001. Dynamic recrystallization processes in plagioclase porphyroclasts. Journal of Structural Geology, 23, 1781-1802. LAFRANCE, B. & VERNON, R. H. 1993. Mass transfer and microfracturing in gabbroic mylonites of the Guadalupe igneous complex, California. In: BOLAND, J. N. & FITZ GERALD, J. D. (eds) Defects and Processes in the Solid State; Geoscience Applications; The McLaren Volume. Developments in Petrology, 14, 151-167. LAIRD, J. & ALBEE, A. L. 1981. Pressure, temperature, and time indicators in mafic schists: their application to reconstructing the polymetamorphic history of Vermont. American Journal of Science, 281, 127-175. LAPWORTH, T., WHEELER, J. & PRIOR, D. J. 2002. The deformation of plagioclase investigated using electron backscatter diffraction crystallographic preferred orientation data. Journal of Structural Geology, 24, 387-399. LEAKE, B. E., WOOLLEY, A. R. et aI. 1997. Nomenclature of amphiboles; report of the subcommittee on amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names. The Canadian Mineralogist, 35, 219-246. LEXA, O. 2003. Numerical approach in structural and microstructural analyses. PhD thesis, Charles University, Prague. LISLE, R. J. 1985. The use of the orientation tensor for the description of statistical fabrics. Journal of Structural Geology, 7, 115-117. LISTER, G. S., PATERSON, M. S. & HOBBS, B. E. 1978. The simulation of fabric development in plastic

deformation and its application to quartzite; the model. Tectonophysics, 45, 107-158. MARSHALL, D. B. & MCLAREN, A. C. 1977a. Deformation mechanisms in experimentally deformed plagioclase feldspars. Physics and Chemistry of Minerals, 1,351-370. MARSHALL, D. B. & MCLAREN, A. C. 1977b. The direct observation and analysis of dislocations in experimentally deformed plagioclase feldspars. Journal of Material Science, 12, 893-903. MARTELAT, J. E., SCHULMANN,K., LARDEAUX,J. M., NICOLLET,C. & CARDON,H. 1999. Granulite microfabrics and deformation mechanisms in southern Madagascar. Journal of Structural Geology, 21, 671-687. MAULER, A., GODARD, G. & KUNZE, K. 2001. Crystallographic fabrics of omphacite, rutile and quartz in Vendee eclogites (Armorican Massif, France); consequences for deformation mechanisms and regimes. Tectonophysics, 342, 81 - 112. McCAI6, A. M. & KNIPE, R. J. 1990. Mass-transport mechanisms in deforming rocks; recognition using microstructural and microchemical criteria. Geology, 18, 824-827. MEANS, W. D. 1990. Kinematics, stress, deformation and material behavior. Journal of Structural Geology, 12, 953-971. MONTARDI, Y. & MAINPRICE,D. 1987. A transmission electron microscopy study of the natural plastic deformation of calcic plagioclases (An 68-70). Bulletin de Mineralogie, 110, 1-14. MORRISON-SMITH, D. J. 1976. Transmission electron microscopy of experimentally deformed hornblende. American Mineralogist, 61, 272-280. NYMAN, M. W., LAW, R. D. & SMELIK, E. A. 1992. Cataclastic deformation mechanism for the development of core-mantle structures in amphibole. Geology, 20, 455-458. OLSEN, T. S. & KOHLSTEDT, D. L. 1984. Analysis of dislocations in some naturally deformed plagioclase feldspars. Physics and Chemistry of Minerals, 11, 153-160. OLSEN, T. S. & KOHLSTEDT, D. L. 1985. Natural deformation and recrystallization of some intermediate plagioclase feldspars. Tectonophysics, 111, 107-131. ORD, A. & HOBBS, B. E. 1989. The strength of the continental crust, detachment zones and the development of plastic instabilities. Tectonophysics, 158, 269-289. PANOZZO, R. H. 1983. Two-dimensional analysis of shape-fabric using projections of digitized lines in a plane. Tectonophysics, 95, 279-294. POIRIER, J. P. & GUILLOPt~, M. 1979. Deformation induced recrystallization of minerals. Bulletin de Mindralogie, 102, 67-74. PRIOR, D. J. & WHEELER,J. 1999. Feldspar fabrics in a greenschist facies albite-rich mylonite from electron backscatter diffraction. Tectonophysics, 303, 29 -49. PRIOR, D. J., BOYLE, A. P. et al. 1999. The application of electron backscatter diffraction and orientation contrast imaging in the SEM to textural problems in rocks. American Mineralogist, 84, 1741-1759.

TEXTURES OF NATURALLY DEFORMED METAGABBROS ROONEY, T. P., RIECKER, R. E. & Ross, M. 1970. Deformation twins in hornblende. Science, 169, 173-175. ROONEY, T. P., RIECKER, R. E. & GAVASCI, A. T. 1975. Hornblende deformation features. Geology, 3, 364-366. ROSENBERG, C. L. & STONITZ, H. 2003. Deformation and recrystallization of plagioclase along a temperature gradient: An example from the Bergell tonalite. Journal of Structural Geology, 25, 389-408. RUTTER, E. H. & BRODIE, K. H. 1988. Experimental 'syntectonic' dehydration of serpentinite under conditions of controlled pore water pressure. Journal of Geophysical Research, B, 93, 49074932. RUTTER, E. H. 8,: BRODIE, K. H. 1992. Rheology of the lower crust. In: FOUNTAIN,D. M., ARCULUS,R. & KAY, R. W. (eds) Continental Lower Crust. Developments in Geotectonics, Elsevier, Amsterdam, 23, 201-267. SCHULMANN, K., ML(~OCH, B. 8,: MELKA, R. 1996. High-temperature microstructures and rheology of deformed granite, Erzgebirge, Bohemian Massif. Journal of Structural Geology, 18, 719-733. SODRE-BORGES, P. & WHITE, S. H. 1980. Microstructural and chemical studies of sheared anorthasites, Roneval, South Harris. Journal of Structural Geology, 2, 273-280. SIMPSON, C. 1985. Deformation of granitic rocks across the brittle ductile transition. Journal of Structural Geology, 7, 503-511. SKROTZKI, W. 1990. Microstructure in hornblende of a mylonitic amphibolite. In: Knipe ROBERT, J. & RUTTER, E. H. (eds) Deformation Mechanisms, Rheology and Tectonics. Geological Society, London, Special Publications, 54, 321-325. SPRY, A. 1969. Metamorphic Textures. Pergamon Press, Oxford. STiPSKA, P., SCHULMANN, K., THOMPSON, A. B., JE~EK, J. & KRONER, A. 2001. Thermo-mechanical role of a Cambro-Ordovician paleorift during the Variscan collision; the NE margin of the Bohemian Massif. Tectonophysics, 332, 239-253. STfPSKA, P., SCHULMANN, K. & KRONER, A. 2004. Vertical extrusion and middle crustal spreading of omphacite granulite: a model of synconvergent exhumation (Bohemian Massif, Czech Republic). Journal of Metamorphic Geology, 22, 179-198. ST15NITZ, H. & F/TZ GERALD,J. D. 1993. Deformation of granitoids at low metamorphic grade; II, Granular flow in albite-rich mylonites. Tectonophysics, 221, 299-324. STUNITZ, H. FITZ GERALD, J. D. & TULLIS, J. 2003. Dislocation generation, slip systems, and dynamic recrystallization in experimentally deformed

125

plagioclase single crystals. Tectonophysics, 372, 215-233. SUN, S. S. & MCDONOUGH, W. F. 1989. Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes. In: SAUNDERS, A. D. & NORRY, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications, 42, 313-345. TULLIS, J. 1983. Deformation of feldspars. In: RmBE, P. H. (ed) Feldspar Mineralogy. Reviews in Mineralogy. Mineralogical Society of America, Washington, DC, 2, 297-323. TULLIS, J. & YUND, R. A. 1987. Transition from cataclastic flow to dislocation creep of feldspar; mechanisms and microstructures. Geology, 15, 606-609. TULLIS, J. & YUND, R. A. 1991. Diffusion creep in feldspar aggregates; experimental evidence. Journal of Structural Geology, 13, 987-1000. TULLIS, J. & YUND, R. 1992. The brittle-ductile transition in feldspar aggregates; an experimental study. In: EVANS, B. & WONG TENG, F. (eds) Fault Mechanics and Transport Properties of Rocks; A Festschrift in Honor of W. F. Brace. Academic Press, San Diego, United States, 89-117. VERNON, R. H. 1976. Metamorphic Processes. Murby, London, Wiley, New York. VOLL, G. 1976. Recrystallization of quartz, biotite and feldspars from Erstfeld to the Leventina Nappe, Swiss Alps, and its geological significance. Schweizerische Mineralogische und Petrographische Mitteilungen, 56, 641-647. WALKER, A. N., RUTTER, E. H. & BRODIE, K. H. 1990. Experimental study of grain-size sensitive flow of synthetic, hot-pressed calcite rocks. In: KNIPE, R. J. & RUTTER, E. H. (eds) Deformation Mechanisms, Rheology and Tectonics. Geological Society, London, Special Publications, 54, 259-284. WILKS, K. R. • CARTER, N. L. 1990. Rheology of some continental lower crustal rocks. Tectonophysics, 182, 57-77. YUND, R. & TULLIS, J. 1991. Compositional changes of minerals associated with dynamic recrystallization. Contributions to Mineralogy and Petrology, 108, 346-355. ZUCALI, M., CHATEIGNER, D., DUGNANI, M., LETTEROTTI, L. & OULADDIAF,B. 2002. Quantitative texture analysis of glaucophanite deformed under eclogite facies conditions (Sesia-Lanzo zone, Western Alps): comparison between X-ray and neutron diffraction analysis. In: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. & PENNOCK, G. M. (eds) Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 219-238.

Stress-strain rate history of a midcrustal shear zone and the onset of brittle deformation inferred from quartz recrystallized grain size F. G U E Y D A N t, C. M E H L 2 & T. P A R R A 3

IG~osciences Rennes, UMR CNRS 6118, Universit~ de Rennes 1, Campus Beaulieu, 35042 Rennes Cedex, France (e-mail: frederic, gueydan @ univ- rennes l.fr) 2Laboratoire de Tectonique, Universit~ Pierre et Marie Curie, UMR CNRS 7072, Case 129, 4 place Jussieu, 75252 Paris Cedex 05, France 3Laboratoire de G~ologie, UMR CNRS 8538, Ecole Normale Supdrieure, 24 rue Lhomond, 75005 Paris, France Abstract: The quantification of quartz shear stress and strain rate within a midcrustal shear

zone provides a mechanical frame to describe the evolution from penetrative ductile deformation to localized deformation and the onset of brittle deformation. The quantification is based on the relationships between the quartz recrystallized grain size, the quartz shear stress (piezometric relation) and the strain rate (dislocation creep flow law). Increasing strain is accompanied by a general decrease of quartz recrystallized grain size and a decrease in grain size scattering. These are interpreted as a result of a complex loading history. The evolution from penetrative ductile deformation toward strain localization, marked by an increase of the strain rate by one order of magnitude, is inferred from grain size memory. Brittle deformation is triggered for quartz shear stress of the order of 70 MPa and strain rate close to 10 -12 s -1. This relative low value of the quartz shear stress necessary to trigger faulting implies a less important strength for the midcrust compared with strengths predicted by classical rheological envelopes.

Metamorphic core complexes are often marked by extensional midcrustal shear zones (Davis 1980; Wernicke 1985; S6ranne & S6guret 1987; Lister & Davis 1989; Andersen et al. 1991; Jolivet et al. 1991, 1998). During the exhumation of lower crustal metamorphic rocks, these shear zones ultimately lead to the formation of detachment faults (i.e. low angle normal fault) at the brittle-ductile transition. Greenschist facies extensional shear zones form in a thickened continental crust at depths close to 15 km (400 ~ and 500 MPa). At these depths, thermally activated viscous creep dominates, leading to a macroscopic ductile behaviour. Brittle deformation mechanisms are required to form the detachment fault, at the later stage of strain localization (Lister & Davis 1989). However, the transition in time and space between ductile strain localization and the brittle deformation along the detachment remains unexplained and is studied here. A quantification of stress and strain rate should explain the rheological evolution of the shear zone from ductile strain to

brittle deformation. The example studied is that of the extensional detachment zone associated with the Tinos metamorphic core complex in the Greek Cyclades. Twiss (1977) first established theoretical relationships between shear stress and either subgrain size or dynamically recrystallized grain size during steady-state creep, corresponding to a power relation between the shear stress and the recrystallized grain size. Experimental calibration of the piezometric relation for quartz (Mercier et al. 1977; Koch 1983) has been used to quantify shear stress within shear zones (Weathers et al. 1979; White 1979a, b; Christie & Ord 1980; Kohlstedt & Weathers 1980). Recently, Stipp & Tullis (2003) have proposed a new calibration of the quartz grain size piezometer for midcrustal conditions. A combination of a piezometric relation and dislocation creep flow laws has also been used to estimate shear stress and strain rate from quartz recrystallized grain size during mylonitization in metamorphic core complexes (Hacker et al. 1990, 1992).

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 127-142. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

128

F. GUEYDAN ET AL.

The relationships between quartz recrystallized grain size evolution within a kilometre-scale shear zone and the related strain localization mechanism remains however unexplored. The aim of this paper is to apply a piezometric relation (Stipp & Tullis 2003) combined with a dislocation creep flow law (Paterson & Luan 1990) to the Tinos Island (Cyclades, Greece) kilometre-scale shear zone, to provide a mechanical frame (quantified in stress, strain rate and temperature) for the evolution from penetrative ductile deformation toward strain localization and the onset of brittle deformation. The microphysical model that relates the recrystallized grain size, the shear stress, the strain rate (at a constant temperature) and the temperature (at a constant strain rate) is first presented. The application of these relations to a microscopic scale then provides explanation for scattering in recrystallized grain size in terms of grain size memory. These findings are finally applied to the Tinos Island extensional shear zone, where the recrystallized grain size and the grain size scatter decrease with increasing strain. The quartz shear stress and the strain rate are increasing towards the core of the kilometre-scale shear zone, whereas the temperature is decreasing. These observations are consistent with an evolution to strain localization. The observed stress variations across a shear zone are in discrepancy with mechanical equilibrium (constant shear stress). These stress variations should therefore be related to different loading events, each event being characterized by a constant shear stress across the shear zone (Weathers et al. 1979). The observed grain size scattering is thus a grain size memory of these successive loading events. Grain size memory provides a mechanical scheme, quantified in stress and strain rate, for the evolution from penetrative ductile deformation towards strain localization.

the steady-state shear stress "rD (in MPa) by the equation: "rD = BD -p

(1)

This relation reflects equilibrium between the elastic energy stored in a grain and the surface energy (Twiss 1977). The values of the exponent p and the constant B depend on the dynamic recrystallization mechanism, either grain boundary migration or subgrain rotation (Guillop6 & Poirier 1979). Derby & Ashby (1987) have suggested that dynamic recrystallization reflects equilibrium between grain size reduction (in the dislocation creep regime) and grain growth (during diffusion creep). The D - r relation should therefore mark a balance between grain size sensitive (grain growth) and grain size insensitive (grain size reduction) flow (de Bresser et al. 1998). The piezometric relation is thus not only stress sensitive (Eq. 1) but also temperature sensitive, leading to significant differences in stress estimate. On this basis, Gueydan et al. (2001) have proposed a grain size sensitive rheology for the lower crust, corresponding to a series association between grain size sensitive and insensitive flow. Grain size sensitive flow is activated for temperature greater than 500 ~ and depths greater than 20 km. Thus, the balance hypothesis may not reflect the dynamically recrystallized grain size for depths close to the brittle-ductile transition. Moreover, recent studies (Drury 2005) have shown that the temperature dependence of the recrystallized grain size is small and not as large as predicted by the model discussed by de Bresser et al. (2001). Therefore, and for the sake of simplicity, the piezometric relation D - rD (Eq. 1) is used in this study to compute a dislocation creep flow stress consistent with the measured dynamic recrystallized grain size D. Quantification o f the strain rate

Microphysical model This section provides the relations between the recrystallized grain size with a steady flow shear stress and strain rate. These relations will then be applied to the Tinos midcrustal shear zone, which is presented in the next section.

Recrystallized grain size p i e z o m e t e r

Experimental work on metals, ceramics and rocks has shown that the recrystallized grain size, denoted hereafter D (in ia,m and usually referred to as 'the grain size') is related to

and the temperature

Having computed a shear stress from the recrystallized grain size (Eq. 1), the strain rate or the temperature can be derived from quartz dislocation creep flow law, which reads: e = A exp - ~ Q - ~rj' r] n

(2)

where/~, T and r are the strain rate (in s-1), the temperature (in K) and the shear stress (in MPa), respectively. Assuming a constant temperature T, the strain rate e is directly related to the shear stress by the flow law given in Equation (2).

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE The combination of this equation with the piezometric relation thus defines the relationships between the strain rate kD and the grain size D, called (D - eo) relation:

I QrlO-"'

eo = AB n exp - RTo_I

(3)

where To is the selected value of the temperature. Similarly, assuming a constant strain rate, the temperature To could be derived from Equations (2) and (1). The temperature-grain size relation (D - T/)) thus reads:

where e0 is the selected value of the strain rate. The subscript D in Equations (3) and (4), stands for the value of the strain rate or temperature estimated for the recrystallized grain size D. These two expressions can only be applied for microstructure formed in the dislocation creep regime of quartz. It is thus necessary to determine the deformation mechanism of Tinos microstructures prior to the application of these relations. The Tinos Island extensional shear zone

This section provides PT conditions for the formation of the Tinos extensional shear zone and the characteristic microstructures of the shear zone, in order to select the appropriate flow laws and paleopiezometer. Geological setting: P T conditions and timing f o r shear zone f o r m a t i o n

The Tinos Island belongs to the 45 Ma highpressure belt of the Hellenides (Jolivet & Patriat 1999). The island is marked by a N W SE elongate dome structure, with a shallow dip of the schistosity on the NE side (Fig. 1). Three tectonic units have been recognized on the island (Avigad et al. 1992; Jolivet & Patriat 1999), from top to bottom (Fig. la): (1) an upper unit, equivalent to the Pelagonien nappe ophiolites and metamorphosed around 70 Ma in the greenschist facies, (2) a lower Blueschist unit (mixture of metapelites and metabasites) affected by eclogitic metamorphism around 45 Ma (Br6cker et al. 1993; Br6cker & Franz 1998) and overprinted by greenschist metamorphism around 20-25 Ma, and (3) a massive dolomite unit of Triassic age, cropping out mainly in NW Tinos. Note that this third unit was not reported in the schematic geological map (Fig. 1). The upper and lower units are

129

separated by a shallow dipping contact, interpreted as an extensional ductile shear zone (Gautier & Brun 1994; Jolivet et al. 1994; Jolivet & Patriat 1999). During the late stage of HP-LT rock exhumation, a shallow dipping brittle detachment fault formed between the upper and the lower unit. An early Miocene (c. 19 Ma, Andriessen et al. 1979; Altherr et al. 1982) complex of granodiorite intruded both the lower and upper unit, mostly in the North Easthern part of the island, and is responsible for contact metamorphism. This granitic intrusion thus marked the end of the extensional shear zone formation and activity. The extension in the Aegean region started 30 Ma ago (Jolivet & Faccenna 2000) and the onset of greenschist overprint around 25 Ma ago. The exhumation rate for the lower unit was such that it reached a depth of c. 28kin (800MPa, Parra et al. 2002) 30 Ma ago and reached 7 km (200 MPa, Parra et al. 2002) 19 Ma ago. The vertical exhumation rate is thus estimated to be approximately 2 ram/a, corresponding to 5 - 6 m m / a along the c. 20 ~ NE dipping Tinos extensional shear zone. From SW to NE, a strain gradient is observed, together with a gradient of greenschist retrogression in both basic and pelitic rocks (Matthews et al. 1999; Parra et al. 2002). Using chloritephengite local equilibria, Parra et al. (2002) estimated PT conditions for extensional shear zone formation and related detachment faulting at around 2 - 4 kbar and around 400 ~ A sharp decrease of phengitic substitution is observed below the detachment (samples 1-4) and in the centre of the island (sample 14-16, Fig. lc, dark grey boxes). This is related to a pressure decrease during the exhumation metapelites and to a local increase in strain, marked by two kilometre-scale shear zones (Fig. lb). Metapelite samples of the Cycladic Blueschists collected every 50 to 200 m along the Ormos Isternia-Panormos Bay cross-section (Fig. 1) are used in this study. Quartz recrystallized grain sizes were optically measured in each sample (around 75 measurement per thin section). These samples were also used for PT estimates using chlorite-phengite local equilibria (Parra et al. 2002). A comparison between the intensity of retrogression (phengitic substitution) and the intensitiy of recrystallization (recrystallized grain size) could therefore be made.

Selected quartz piezometer and f l o w law f o r the Tinos shear zone

Quartz grains in Tinos Island shear zones are extremely flattened and exhibit optically visible

130

F. G U E Y D A N E T AL.

a- Simplified .qeological map of Tinos Island

b- Simplified cross section and sample location

c- Phengitic substitution

Fig. 1. (a) Structural map of the Tinos Island (after Gautier & Brun 1994; Patriat & Jolivet 1998; Jolivet & Patriat 1999), showing the superposition of two structural units (upper unit and lower unit), separated by a detachment. The samples used in this study are located on the map. The two insets show typical microstructures of sample 23 (bottom left) and sample 1 (top right). (b) Schematic cross section showing the sample location and the position of the extensional shear zone (grey zone). (c) Decrease of the degree of phengitic substitutions (Parra et al. 2002) towards the detachment shows an increase of overprinting from SW to NE. Note an increase of the overprinting in the middle of the island (samples 14 and 16), consistent with the presence of a major shear zone shown in (b).

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE subgrains, as well as recrystallized grains along grain boundaries (Fig. la, sample 23). These microstructures are typical of regime 2 of dislocation creep (Hirth & Tullis 1992), where dynamic recrystallization is accommodated by subgrain rotation. At a closer distance to the detachment, in samples more intensely strained, quartz grains are completely recrystallized, forming an oblique foliation, a feature characteristic of regime 3 (Fig. 1a, sample 1). Tinos Island microstructures, created during shear zone formation at greenschist conditions, are thus typical of the transition from regime 2 to regime 3. The application of the piezometric relation to the continental crust is based on the calibration of Equation (1) using microstructures from experimentally deformed quartz aggregates (Mercier et al. 1977; Twiss 1977; Koch 1983; Stipp & Tullis 2003). The Twiss piezometer, which has been widely used to estimate shear stress in naturally deformed rocks, is accurate for conditions around the transition between regime 2 and regime 3 of dislocation creep (Hirth & Tullis 1992). This piezometer has been however, criticized for an incorrect application of thermodynamics equilibrium (de Bresser et al. 2001). Piezometers from Mercier et al. (1977) and Koch (1983) give shear stress estimates lower than the Twiss piezometer, because of a lower grain size exponent. The Stipp and Tullis piezometer (Stipp & Tullis 2003) was also recently calibrated for regimes 2 and 3, and provides stress estimates very close to those obtained with the Twiss piezometer. Since Tinos Island microstructures are characteristic of regime 2 and 3 dislocation creep, the Stipp & Tullis (2003) piezometer appears as the more relevant shear stress estimate for our study and is the one selected here. Note that recrystallized grain size measurements were conducted within quartz aggregates (Fig. l a, sample 1) in order to prevent pinning effects due to the presence of phyllosilicate layers, which could lead to non-piezometric (e.g. non-steady-state) recrystallized grain size. Dislocation creep flow laws for quartz, quartzite and quartz aggregates are relatively abundant (Koch et al. 1989; Paterson & Luan 1990; Luan & Paterson 1992; Gleason & Tullis 1995). Recently, Hirth et al. (2001) discussed the validity and applicability of these flow laws using a combination of microstructural, geochronological and geological observations. Following Hirth et al. (2001), dislocation creep parameters of Paterson & Luan (1990) and Luan & Paterson (1992) are selected here and given in Table 1.

131

Table 1. Dislocation creep (Paterson & Luan 1990) and piezometer (Stipp & Tullis 2003) parameters used to compute the shear stress (Eq. 1), the strain rate (Eq. 3) and the temperature (Eq. 4)from the measured recrystallized grain size. The value of the strain rate eo used to estimate the temperature is set to 10 -15 s -1. The temperature To, used to estimate the strain rate, is set to 400 ~ owing to the PT conditions determined using chlorite-phengite local equilibria (Parra et al. 2002)

Parameter values Dislocation creep A (MPa-ns -1) Q (kJ/mol) n Piezometer B (Mpa. ixm-p) p Strain rate e0 (s -1) Temperature To (~

10 9.4 135 4 3631 1.26 10 15 400

Based on the PT conditions determined by Parra et al. (2002), the temperature is set to To = 400 ~ for the estimate of the strain rate ko (Eq. 3). For the temperature estimate To (Eq. 4), the strain rate is set to e 0 - 10-aSs-1 (Table 1). Having defined the appropriate flow law and piezometer, the measured recrystallized grain size could be related to a shear stress, strain rate and temperature, using Equations (1) to (4).

Grain size evolution across a microscopic shear band: Stress history of ductile strain localization The aim of this section is to apply the D - ~-, D - kD and D - To relations (Eqs. 1, 3 and 4, respectively) at a microscopic scale. This microscopic study will provide explanation for the presence of scattering in D values within a microscopic shear band. These interpretations will be used at macroscopic scale within the Tinos shear zone. D e c r e a s e o f recrystallized grain size with increasing strain

Ten optical measurements of the minimum width of quartz grain were conducted in three sites of a microscopic shear band (Fig. 2a), cropping out in the vicinity of the detachment fault (sample 3, Fig. 1). The scale of each site of measurement is of the order of 1 mm (Fig. 2a, grey boxes). Ten optical measurements of recrystallized grain size of c. 100 Ixm is thus the maximum number

F. GUEYDAN ET AL.

132

a- Microscopic shear bands and interpretative scheme

b- Quartz recrystallized grain size

c- Quartz shear stress

d- Strain rate

e- Temperature

f-Two step history

Fig. 2. Microscopic study: (a) microscopic shear band (sample 1, Fig. 1) and an interpretative scheme, where the sites of recrystallized grain size measurements have been reported; (b) recrystallized grain size; (c) shear stress (estimated by Eq. 1), (d) strain rate (Eq. 3), and (e) temperature (Eq. 4) variation as a function of the distance to the microscopic shear band; (f) A two-step history could explain the scattering in grain size, shear stress, strain rate and temperature. of measurements in this micro-scale study. The number of measured recrystallized grains for the kilometre-scale study presented in the next section has been significantly increased (75 measurements). Only optically undeformed quartz grains (i.e. recrystallized grains free of undulatory extinction) were measured.

The mean value of recrystallized grain sizes decreases from 100 ~m to 30 Ixm in the core of the microscopic shear band. A scattering in D values is observed in each site of measurement. Far from the shear band, D values range between 160 Ixm and 40txm (grey box, Fig. 2b), while in the core of the shear band,

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE recrystallized grain sizes vary from 40 Ixm to 20 Ixm (dark grey box, Fig. 2b). The maximum and minimum value of the recrystallized grain size thus decreases with increasing strain. The scattering is thus decreasing with increasing strain. The shear stress and strain rate values (Fig. 2c, d) are computed from the recrystallized grain sizes, using Equations (1) to (3) with constant values of Table 1. Both shear stress and strain rate are increasing towards the core of the shear band. The increase in shear stress is sharper than the decrease in grain size, because of the power dependency between the grain size and the shear stress (Eq. 1). Quartz shear stress ranges between 4 and 32MPa far from the shear bands, whereas it reaches values of up to 60 MPa in the shear band core. Strain rate starts between 1 0 - 1 7 S -1 a n d 1 0 - 1 4 s -1 and increases to values closes to 10 -13 s -1 within the shear band. This increase in stress and strain rate is compatible with the increase in strain within the shear bands. The scattering in D values induces a scattering in shear stress and strain rate. The increase in shear stress could also be related to a decrease in temperature (Fig. 2e, Eq. 4). Temperature TD reaches values as low as 275 ~ in the core of the shear band, whereas TD is of the order of 425 ~ outside the band. This significant decrease of temperature, at a constant strain rate, could also favour localization. The estimated values of the strain rate and temperature should thus be seen as maximum and minimum estimates, respectively. The coupled estimate of strain rate and temperature variation related to shear stress changes cannot be made analytically and would require a numerical simulation, which was out of the scope of the present work. Within the shear band, simple shearing mode of deformation is expected to be dominant. Mechanical equilibrium for simple shear implies a condition of constant shear stress. The scattering in shear stress and the increase of the mean shear stress towards the core of the shear band is not consistent with mechanical equilibrium across the shear band. This discrepancy could however be explained by the grain size memory and therefore by a change in time of the applied shear stress, as discussed next.

Recrystallized grain size memory Two types of processes are commonly evoked to explain the scattering in D values, as observed in each sample. First, scattering could correspond to a grain size distribution. Experimental studies

133

have indeed shown that dynamic recrystallization at a given shear stress commonly leads to a grain size distribution, characterized by a mean value and a standard deviation, and therefore may not lead to a single value of the grain size D (Ranalli 1984). The decrease of the dispersion, AD towards the core of the shear zone is consistent with experimental results showing that the standard deviation of the grain size distribution decreases with increasing strain (Ter Heege 2002). Second, the recrystallized grain size scattering could also reflect a complex loading history. As pointed out by Weathers et al. (1979), a second stress increase during a relatively short time lapse will induce the formation of smaller grain size. The short time lapse implies that this second event will not lead to a total recrystallization into smaller grains, so that the smaller grain size reflect the final stress level (highest stress value) and the larger grain size corresponds to the earlier lower stress level. Such explanation of the grain size scattering in terms of stress/loading history is not in conflict with the hypothesis of grain size distribution. Indeed, a bimodal grain size distribution also indicates two phases of stress pulse (Weathers et al. 1979). In the absence of a grain size distribution paleopiezometer in the literature, the quartz piezometer (Eq. 1) is used in this study, with the interpretation of the grain size scattering in terms of loading history. The necessity of having a constant shear stress within the shear zone at a given time (mechanical equilibrium) provides also evidence for grain size memory, each recrystallized grain size being related to a loading event, characterized by a constant shear stress across the shear zone. On these bases, the grain size distribution as a function of the distance to the shear zone can be divided into at least two successive stress events. The first stress event is marked by grain size values between 160 Ixm and 40 Ixm, observed in all sites, shear stress between 4 and 32 MPa and strain rate between 10 -17 and 1 0 - 1 4 S -1 (grey box, Fig. 2). The second stress event leads to grain size values as low as 2 0 - 4 0 p~m, only observed in the vicinity of the shear bands, and to shear stress and strain rate values of up to 60MPa and 10 -13 s -1, respectively (dark grey box, Fig. 2). Being present in all samples, the first event should be related to penetrative ductile deformation (Fig. 2f). The scattering in D during this event could correspond to different stress loading, varying between 10 MPa and 50 MPa, and also to local stress heterogeneities, due for example to the presence of shadow zones. The second event marks ductile strain localization, which leads to the formation

134

F. GUEYDAN ETAL.

of the microscopic shear band (Fig. 2f). Again, the scattering in D values within this second event reflects small variation of stress between 32 and 60 MPa. Sites that suffered event 2 also show grain size associated with event 1. This is explained by the balance between the timing of change in stress and the timing of dynamic recrystallization. If stress increases faster than the time necessary for a complete recrystallization, grain size will keep in memory an earlier stage of deformation. Recrystallized grain size scattering can be interpreted in terms of grain size memory. The scattering in D can thus be related to the time evolution of the ductile flow. In summary, the scattering in recrystallized grain size provides evidence for grain size memory. The evolution in stress, strain rate and temperature from penetrative ductile deformation towards strain localization could therefore be quantified from the measurement of the recrystallized grain size. Penetrative ductile deformation affects all the material and is thus marked by a similar distribution of recrystallized grain size in all sites. In contrast, ductile strain localization is confined within the shear band and is marked by smaller grain sizes. Grain size evolution within the Tinos shear zone: M e c h a n i s m of the formation of midcrustal shear zone and detachment fault The microscopic scale findings interpreted in terms of grain size memory are now applied to a kilometer-scale shear zone, in order to quantify the stress and strain rate related to ductile strain localization and detachment faulting.

boxes, Fig. 3a), a sharp decrease of the minimum and maximum values of the grain size is observed with respect to samples collected at larger distance to the detachment (samples 18 to 23, grey box, Fig. 3a). The distribution of the frequency of the recrystallized grain size for four samples (23, 19, 12 and 3) is given in Figure 4. Table 2 gives the mean values, the standard deviation, the peak value and the dispersion (AD) for the whole set of measurements. For the four samples (23, 19, 12 and 3), the distribution is approximately lognormal. With increasing strain (from sample 23 to sample 3), the mean value of the lognormal distribution is moving towards smaller grain size (143.59 p~m to 57.24 p~m, Table 2) and the standard deviation o- is decreasing (from 43.19 ~m to 18.65 ~zm, Table 2). The scattering in D values is thus significantly decreasing with increasing strain, as already shown in the microscale study. Again, the scattering in D values could correspond to a grain size distribution (Ranalli 1984), which could therefore explain the decrease of the standard deviation with increasing strain (Ter Heege 2002). This scattering could also correspond to a complex loading history, as discussed above at microscopic scale (Weathers et al. 1979). Three main sets of grain size, and associated loading events can thus be distinguished: (1) grain sizes between 260 p~m and 60 ~m (grey box, Fig. 3a), (2) grain sizes between 60 ~m and 25 p~m (dark grey box, Fig. 3), and (3) grain sizes lower than 25 p~m (black box, Fig. 3). Step 1 affects the entire studied domain whereas steps 2 and 3 are observed in more localized zones. The Tinos Island shear zone therefore appears to have suffered these three successive events, starting with event 1 (in all samples), then with event 2 and finally with event 3.

Recrystallized grain size evolution Within the Tinos Island extensional shear zone (Fig. 1), 75 optical measurements of quartz recrystallized grain size were randomly performed in thin sections (13 samples, Fig. 3). For a given sample, a wide range of grain size value D is observed. At large distance from the detachment, D ranges between 260 ~m and 60 p~m, and below the detachment D varies between 150 l.zm and 15 p~m. Despite the D scattering, the mean value of the grain size decreases towards the detachment, from 160 p~m to 20 p~m. The maximum and minimum values of D decrease from 260 p~m to 150 p~m and from 60 ~m to 15 p~m, respectively. The dispersion thus decreases from AD = 200 p~m to AD = 135 ~m. For distances less than 1600 m from the detachment (samples 16 to 1, two dark grey

Shear stress, strain rate and temperature evolution The quartz shear stress, the strain rate and the temperature, computed from the recrystallized grain size, are plotted in Figures 3b, 3c and 3d, respectively. The sharp decrease of the grain size near the detachment implies a sharp increase in shear stress and strain rate, reaching values close to 100 MPa and 2 • 10 -12 s -1, respectively. As reported at microscopic scale, the shear stress is not constant within the kilometer-scale shear band. This observation violates mechanical equilibrium in the shear zone, which requires a constant shear stress. Shear stresses computed across the shear zone are therefore not contemporaneous but relate to at least three

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE

135

Fig. 3. Kilometre-scale study: (a) quartz recrystallized grain size, (b) quartz shear stress (Eq. 1), (c) strain rate (Eq. 3), and (d) temperature (Eq. 4) as a function of the horizontal distance to the detachment fault. A three-step history is proposed to explain the grain size, stress, strain rate and temperature evolution: (1) penetrative ductile deformation (grey boxes), (2) ductile strain localization (dark grey boxes), and (3) onset of brittle deformation (black boxes).

successive steps. The first stress loading (60 txm < D < 260 txm, grey boxes, Fig. 3) corresponds to quartz shear stresses less than 22 MPa and strain rate less than 3 • 10 -15 s -1. Event 1 is therefore characterized by a mean grain size, shear stress and strain rate of 160 txm, 11 MPa and 1.5 • 10 -15 s -1. Since the entire domain has suffered event 1, penetrative ductile deformation should have dominated during this first loading event. This loading event is marked by the pervasive development of conjugate shear bands in the southern part of the Island (Jolivet & Patriat 1999), which is consistent with a penetrative ductile deformation. During the second step (25 i~m < D < 60 p~m, dark grey boxes,

Fig. 3) quartz shear stress and strain rate are between 2 2 M P a and 4 2 M P a and 3 x 10 -15 s -1 and 5 x 10 -14s -1, respectively, leading to mean values of grain size, shear stress and strain rate of 42.5 ~m, 32 MPa and 2.6 x 10-14s-1. The strain rate during this second loading event is approximately 17 times larger than the strain rate during event 1. Strain localization has thus been activated during this secondary event. Consistently, only a restricted zone of 1600 m in width (samples 1 to 16) has suffered this second event. This tendency for localization is marked in the field by the greenschist overprint and the development of the extensional shear zone (Figs la and b, where the grey

F. GUEYDAN ET AL.

136

Fig. 4. Grain size distribution for (a) sample 23, (b) sample 19, (e) sample 12, and (d) sample 3 (located on Fig. 1) shows the frequency (number of measured grains) as a function of the recrystallized grain size. In total, 75 measurements have been made for each sample. The distribution is almost lognormal. The mean value and the standard deviation is decreasing with increasing strain (sample 23 to 3). Table 2 gives the mean value and the standard deviation for the studied samples.

Table 2. Mean value, standard deviation, peak value (grain size obtained f o r the maximum number of measurements), minimum value, maximum value and dispersion AD = max-rain, in tzm, for the whole set of measurements. Sample location is given in Figure 1 Sample 23 21 19 18 16 14 12 8 5 4A 4B 3 2 1

Mean value (~m)

Std. dev. o- (~m)

Peak value (/.zm)

Min (~m)

Max (/.~m)

Dispersion AD (~m)

143.59 114.38 136.91 121.93 98.67 87.51 77.72 61.65 61.35 45.67 45.11 57.24 49.27 48.38

43.19 33.43 37.83 41.32 33.49 28.55 26.03 25.57 25.06 20.13 21.74 18.65 17.61 17.32

102.13 85.11 127.66 102.13 85.11 85.11 76.60 42.55 59.57 42.55 25.53 51.06 34.04 42.55

68.09 51.06 59.57 17.02 51.06 42.55 34.04 25.53 25.53 17.02 17.02 17.02 25.53 17.02

238.30 187.23 252.00 255.32 204.26 165.00 140.00 126.00 148.00 102.00 104.00 103.00 100.00 105.00

170.21 136.17 192.43 238.30 153.19 122.45 105.96 100.47 122.47 84.98 86.98 85.98 74.47 87.98

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE colour stands for the Tinos extensional shear zone). The domain where step 2 is marked is also characterized in the field by the two major shear bands that are observed on the island (Fig. 1b). Note that the computed shear stress is not the overall shear stress, which should decrease in the shear zone due to softening responsible for localization, but the partial shear stress for quartz. The increase of the quartz partial stress should therefore be balanced by an increase of the other phase partial shear stress (e.g. mica) to enforce mechanical equilibrium (constant overall shear stress). Finally, the third step (15 txm < D < 25 txm, mean value of 20 Ixm) is marked by an average quartz stress of 72 MPa and an average strain rate of 10 -12 s -1. Compared with the second event, the strain rate has been increased by almost two orders of magnitude and a thinning of the deformation zone is observed (250 m thick zone below the detachment, compared to 1600 m during step 2). These features are consistent with a more pronounced tendency toward localization. This third event is moreover marked by microcracks within the small recrystallized quartz grain and by semi-brittle faults at all scales (Patriat & Jolivet 1998; Jolivet & Patriat 1999). Brittle deformation has thus been triggered during event 3. The shear stress increase could also be related to a decrease in temperature (Fig. 3d, Eq. 4). The mean temperature indeed decreases from c. 4 8 0 ~ (step 1) to c. 330 ~ (step 2) and reaches low value during step 3 (c. 260 ~ This temperature decrease could also favour strain localization. The evolution from step 1 to step 3 should therefore be marked by both an increase in strain rate and a decrease in temperature. The quantification of the simultaneous variation in strain rate and temperature related to the variation in shear stress at the scale of the island cannot be made analytically and required a numerical model, which was out of the scope of the present work.

Validation of the proposed method for the quantification of stress, strain rate and temperature Comparison of the estimated strain rates with the exhumation rate. The PTt data, computed for Tinos Island metapelites by Parra et al. (2002), Brrcker et al. (1993) and Brrcker & Franz (1998), have provided an estimate of the exhumation velocity to be of the order of 5 - 6 ram/ a along the c. 20 ~ NE dipping Tinos extensional shear zone. Independently, the product of the

137

estimated mean strain rate & the width of the deformation zone during the three successive loading events could also give an estimate of the velocity that prevailed at shear zone boundary. Such an estimate cannot be made during step 1 since the width of the deformation domain is poorly constrained (much greater than the island). During step 2, a mean strain rate of 2 . 6 • 10 -15s -1 and a width of the order of 3 km, led to a velocity of 1.3 mm/a, whereas during step 3, the mean strain rate and the width have been estimated around 10-12 s-1 and 200 m, respectively, which yields a velocity of 7.9 mm/a. The average value of the estimate velocity during steps 2 and 3 is therefore 4.6 mm/a, a value very close to that deduced from PTt data ( 5 - 6 ram/a). Moreover, the velocity estimate inferred from strain rates shows an increase during step 3, which is consistent with the trend toward brittle deformation. Thus, these results permit us to validate a posteriori the method used in this study for the quantification of the strain rate from the measurement of recrystallized grain size.

Comparison of the estimated temperature variation with the PT path. Parra et al. (2002) have used chlorite-phengite local equilibria, computed in the samples used in this study, to estimate the PT path of the Tinos metapelites. From 30 Ma ago to 20 Ma ago, the temperature is found to decrease from 550 ~ to 300 ~ (Parra et al. 2002, p. 61, Fig. 14). During steps 1 and 2, the temperature is found here to decrease from 600 ~ to 300 ~ (Fig. 3d), a tendency that is quite consistent with the estimate of Parra et al. (2002). Temperature then reaches value as low as 225 ~ during step 3. This third step, which reveals an evolution towards brittle deformation, was not calibrated by local equilibria, probably because of the absence of mica crystallization in the brittle domain. The good consistency between our estimate of temperature variation and that independently computed with mineral local equilibria validates our method of quantification of the temperature that prevailed across the shear zone. Comparison with the evolution of phengitic substitutions. A third validation of the method used to quantify shear stress and strain rate is given by the analysis of the phengitic substitution (Fig. lc). Phengitic substitutions decrease with the increase in strain. During penetrative ductile deformation (step 1, grey box, Figs lc and 3), phengitic substitutions between 3.65 and 3.35 are recorded in the entire studied area.

138

F. GUEYDAN ET AL.

Strain localization (step 2, dark grey boxes, Figs lc and 3) is marked by phengitic substitutions lower than 3.35, because of the larger retrogression in the shear zones. Note that the evolution towards brittle deformation (step 3) does not modify the phengitic substitution. Again, this could be explained by the absence of mica crystallization in the brittle domain. The good correlation between the intensity of retrogression, outlined by the distribution of the phengitic substitutions, and the strain localization, recorded with recrystallized grain size, validates the use of piezometer combined with dislocation creep flow in order to quantify stress and strain rate.

Discussion The three loading events recorded in Tinos Island metapelites have been plotted on a deformation mechanism map in order to constrain in terms of shear stress and grain size the fields of strain localization and brittle deformation (Fig. 5). On the map, dislocation creep flow stresses at a given strain rate are represented as horizontal lines, marking the grain size insensitivity of the flow. Increasing strain rate leads to an increase of the quartz flow stress, because of the power dependence between shear stress and strain rate (Eq. 2). The Stipp and Tullis piezometer (Eq. 1) corresponds to a line in this log-log diagram, because of the power law relation between the

shear stress and the recrystallized grain size. Measured recrystallized grain sizes are plotted along this line, as boxes for the three steps. The protolith has a quartz grain size of the order of 1 mm (black box, Fig. 5). The evolution from the protolith to the shear zone (step 2) and the onset of brittle deformation (step 3) are marked by a decrease of the recrystallized grain size. The decrease in grain size could therefore be used, as well as the phengitic substitution, as a good indicator of the strain intensity. Strain localization is marked by an increase in strain rate by at least a factor of 17, with respect to the less deformed rocks. The quartz shear stress increase during localization is less pronounced. The onset of brittle deformation is here observed for relatively low values of the quartz shear stress (c. 70 MPa) but large values of the strain rate (c. 10 -13 to 10 -12 s-J). This value of the shear stress is low compared with the predicted strength of the midcrust (Brace & Kohlstedt 1980; Carter & Tsenn 1987; Le Pichon & Chamot-Rooke 1991; Ranalli 1995, 2O00). The computed value of the shear stress only depends on the selected piezometer. The Stipp and Tullis piezometer (Stipp & Tullis 2003) that was selected here does not lead to the largest values of the shear stress for small grain size, compared to the Twiss piezometer (Twiss 1977), also reported on Figure 5. Using the Twiss piezometer, the shear stress estimate for

Fig. 5. Deformation mechanism map (log-log scale), where quartz shear stress is plotted as a function of grain size. Setting the temperature to 400 ~ dislocation creep flow stress for a given strain rate corresponds to horizontal dashed lines, marking the grain size insensitivity of the flow law. Piezometric lines of Twiss (1977) and Stipp & Tullis (2003) are drawn. The dimension of the boxes for the three steps, in stress, strain rate and grain size, is deduced from estimates shown in Figure 3. During step 3, the onset of brittle deformation is observed, which permits us to schematically locate a brittle deformation domain (dashed region). The protolith quartz grain size (black box) is approximately 1 mm.

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE grain sizes lower than 70 ixm will be a little larger. Stipp & Tullis (2003) have however shown with little ambiguity that their piezometer is well calibrated for small grain size, compared to the Twiss piezometer. These two piezometers give also the largest stress estimate compared to other piezometers (Mercier et al. 1977; Koch et al. 1989). The low values of the quartz shear stress computed in this study are therefore maximum estimates. The values of strain rate depend on the piezometer, the flow law and the temperature To (Eq. 3). Following Hirth et al. (2001), the quartz dislocation creep flow law of Luan & Paterson (1992) gives the best estimate of quartz strength for midcrustal conditions. Using the Gleason & Tullis (1995) flow law yields significantly lower strain rate estimations, because of the lower pre-exponential A value (Eq. 2). For example, strain rate in step 2 would be of the order of 10 -15 S-1 instead of 1 0 - 1 4 S -1 computed here. However, the tendency observed in this study and the strain rate ratio from step 1 to step 3 will be the same. The temperature was set to 400 ~ according to the PT estimate of Parra et al. (2002). However, during strain localization, an increase in the intensity of the retrogression (Fig. lc) should be not only related to strain increase but also to a progressive exhumation, leading to a temperature decrease. This temperature decrease is well estimated when setting a constant strain rate (Fig. 3d). In this study, and for the sake of simplicity, strain rate increase was computed at a constant temperature (Eq. 3), and temperature decrease was computed at a constant strain rate (Eq. 4). However, a decrease in temperature during steps 2 and 3 would then lead to a less pronounced increase of the estimated strain rate. A temperature of 350 ~ within the shear zone, instead of 400 ~ outside, could for example explain the increase in shear stress in the shear zone at a constant strain rate. Shear zones must however be marked by an increase in strain rate. Therefore, the evolution from penetrative ductile deformation to strain localization should be related to an increase in strain rate, probably less pronounced than computed in this study, and to a temperature decrease due to the progressive unroofing of the midcrust. However, a quantification of the simultaneous increase of strain rate and decrease of temperature related to the shear stress increase would require numerical simulations. The low value of 70 MPa for the onset of fracturing at midcrustal depths (10-15 km) corresponds to a friction coefficient between 0.1 and 0.2. Such low values of the friction coefficient are consistent with the metapelitic composition

139

of the Tinos tectonites (Byerlee 1978) and with potential fluid effects that are likely to occur at midcrustal depths (Brace & Kohlstedt 1980). For the Tinos Island detachment fault, field evidence suggests fluid circulation at these depths, leading to the formation of numerous quartz veins (Patriat & Jolivet 1998). The role of mica-rich shear zone or fault gouge could also significantly decrease the midcrust strength, leading to stress values close to 6 0 - 7 0 MPa, as pointed out by Bos et al. (2000) with experimental study and by Gueydan et al. (2003) with a numerical model.

Concluding remarks The aim of this paper was to estimate stresses, strain rates and temperatures from recrystallized grain sizes D, across a kilometre-scale extensional shear zone, bounded by a brittle detachment formed at a later stage of exhumation (Gautier & Brun 1994; Jolivet & Patriat 1999). Quartz stresses, strain rates (at constant temperature) and temperatures (at constant strain rate) are computed from D using the Stipp & Tullis (2003) piezometer and a dislocation creep flow law (Luan & Paterson 1992). Quartz recrystallized grain size (D) decreases towards the intensely strained zone, reaching values close to 15 ~m within the detachment. Consequently, quartz shear stress and strain rate increase up to values of 100MPa and 10 -12 S--1 close to the detachment. A scattering in the recrystallized grain size, and therefore in the shear stress and strain rate, is interpreted here as related to a grain size memory resulting from a complex loading history applied to the Tinos Island metapelites. Three main loading events, successive in time, are observed and reported in a crustal cross section in Figure 6. A penetrative ductile deformation, characterized by D c. 160 ~m, a shear stress of 11 MPa and a strain rate of 1.5 x 10 -15 s -1, corresponds to a first step that affects the whole midcrustal rocks. This first event was followed by strain localization, marked by a decrease of D (c. 42.5/xm), and an increase of the shear stress (32MPa) and strain rate (2.6 • 1 0 - 1 4 s - 1 ) . This second event, observed across a 1600 m thick deformation zone, marks the formation of extensional shear zones at midcrustal conditions. Strain localization corresponds to an increase of the strain rate by at least a factor of 17 within the shear zone. It was finally followed by the onset of brittle deformation, with c. 20 p~m recrystallized grain sizes, corresponding to a shear stress of 72 MPa, and a strain rate of up to 10 -12 S -1. This final step corresponds to the

140

F. GUEYDAN ETAL.

Fig. 6. Schematic crustal evolution from midcrustal strain localization to detachment formation. Corresponding mean values of quartz shear stresses and strain rates for (1) penetrative ductile deformation, (2) strain localization, and (3) onset of brittle deformation are from Figure 3.

formation of the d e t a c h m e n t fault during a late stage of extension. The main information concerning this final step is the low value of the shear stress necessary for the onset of brittle deformation. This low value corresponds to a friction coefficient of the order of 0.2, which is relevant for midcrustal conditions if fluid effects are invoked (Brace & Kohlstedt 1980). This value o f the midcrust strength is also consistent with an upper crustal strength governed by weak faults (Jackson 1987; W e r n i c k e 1992, 1995). The reviews ofF. Renard and M. Drury as well as the considerable works made by the Associate Editor, D. Gapais, helped us in preparing the final version of this paper.

References ALTHERR, R., KREUZER, H., WENDT, I., LENZ, H., WAGNER, G. & KELLER, J. 1982. A late oligocene/early miocene high temperature belt in the attic-cycladic-crystalline complex (SE Pelagonian, Greece). Geologisches Jahrbuch, E23, 96-164. ANDERSEN, T. B., JAMTVEIT, B., DewEY, J. F. & SWENSSON, E. 1991. Subduction and eduction of continental crust: major mechanism during continent-continent collision and orogenic

extensional collapse, a model based on the south Norwegian Caledonides. Terra Nova, 3, 303-310. ANDR1ESSEN, P. M., BOELRIJK, N. A. I. M., HEBEDA, E. H., PRIEM, H. N. A. & VERDUMEN, E. 1979. Dating the events of metamorphism and granitic magmatism in the alpine orogen of Naxos (Cyclades, Greece). Contribution to Mineralogy and Petrology, 69, 215-225. AVIGAD, D., MATTHEWS, A., EVANS, B. W. & GARFUNKEL,Z. 1992. Cooling during the exhumation of a blueschist terrane: Sifnos (Cyclades, Greece). European Journal of Mineralogy, 4, 619-634. Bos, B., PEACH,C. J. & SPIERS, C. J. 2000. Frictionalviscous flow of simulated fault gouge caused by the combined effects of phyllosilicates and pressure solution. Tectonophysics, 327, 173-194. BRACE, W. F. & KOHLSTEDT, D. L. 1980. Limits on lithospheric stress imposed by laboratory experiments. Journal 03" Geophysical Research, 85, 6248-6252. BROCKER, M. & FRANZ, L. 1999. R b - S r isotope studies on Tinos Island (Cyclades, Greece): additional time constrains for metamorphism, extent of infiltration-controlled overprinting and deformation activity. Geological Magazine, 135 (3), 368-382. BROCKER, M., KREUTZER, H., MATTHEUWS, H. & OKRUSH, M. 1993.4~ and oxygen isotope studies of polymetamorphism from Tinos Island, Cycladic blueshist belt, Greece. Journal of Metamorphic Geology, 11, 223-240.

MECHANICAL HISTORY OF A CRUSTAL SHEAR ZONE BYERLEE, J. D. 1978. Friction of rocks. Pure and Applied Geophysics, 116, 615-626. CARTER, N. L. & TSENN, M. C. 1987. Flow properties of continental lithosphere. Tectonophysics, 136, 27 -63. CHRISTIE, J. M. & ORD, A. 1980. Flow stress from microstructures of mylonites: example and current assessment. Journal of Geophysical Research, 85, 6253-6262. DAVIS, G. A. 1980. Problems of intraplate extensional tectonics, western United States. In: Continental Tectonics. National Academy of Science, Washington, DC. DE BRESSER, J. H. P., PEACH, C. J., REIJS, J. P. J. & SPIERS, C. J. 1998. On dynamic recrystallization during solid state flow: effects of stress and temperature. Geophysical Research Letters, 25, 3457-3460. DE BRESSER, J. H. P., TER HEEGE, J. H. & SPIERS, C. J. 2001. Grain size reduction by dynamic recrystallization: can it result in major rheological weakening? International Journal of Earth Sciences, 90, 28 -45. DERBY, B. & ASHBY, M. F. 1987. On dynamic recrystallization. Scripta Metallica, 21, 879-884. DRURY, M. R. 2005. Dynamic recrystallization and strain softening of olivine in the laboratory and the lithosphere. In: GAPAIS, D., COBBOLD, P. R. & BRUN, J. P. (eds) Deformation Mechanisms, Rheology and Tectonics: .from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 143-158. GAUTIER, P. & BRUN, J. P. 1994. Crustal-scale geometry and kinematics of late-orogenic extension in the central Aegean (Cyclades and Evvia island). Tectonophysics, 238, 399-424. 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.

GUEYDAN, F., LEROY, Y. M. & JOLIVET, L. 2001. Grain size sensitive flow and shear stress enhancement at the brittle-ductile transition of the continental crust. International Journal of Earth Sciences, 90, 181 - 196. GUEYDAN, F., LEROY, Y. M., JOLIVET, L. & AGARD, P. 2003. Analysis of continental midcrustal strain localization induced by reaction-softening and microfracturing. Journal of Geophysical Research, 108, 2064, doi: 10.1029/2001JB000611. GUILLOPI~,M. & POIRIER, J. P. 1979. Dynamic recrystallization during creep of single-crystalline halite: an experimental study. Journal of Geophysical Research, 84 (B), 5557-5567. HACKER, B., CHRISTIE, J. & SNOKE, A. 1990. Differential stress, strain rate, and temperature of mylonitization in the Ruby Mountains, Nevada: Implications for the rate and duration of uplift. Journal of GeophysicaI Research, 95, 8569-8580. HACKER, B., YIN, A., CHRISTIE, J. & DAVIS, G. 1992. Stress magnitude, strain rate, and rheology of extended middle continental crust inferred from quartz grain sizes in the Whipple Mountains, California. Tectonics, 11, 36-46.

141

HIRTH, G. & TULLIS, J. 1992. Dislocation creep regimes in quartz aggregates. Journal of Structural Geology, 14, 145-159. HIRTH, G., TEYSSIER, C. & DUNLAP, W. J. 2001. An evaluation of quartzite flow laws based on comparisons between experimentally and naturally deformed rocks. International Journal of Earth Sciences, 90, 77-87. JACKSON, J. A. 1987. Active normal faulting and continental extension. In: COWARD, M. P., DEWEY, J. F. & HANCOCKEY,P. L. (eds) Continental Extension. Geological Society, London, Special Publications, 28, 3-17. JOLIVET, L. & PATRIAT, M. 1999. Ductile extension and the formation of the Aegean Sea. In: DURAND, B., JOLIVET, L., HORVATH, F. & SERANNE, M. (eds) The Mediterranean Basins: Tertiary Extension Within the Alpine Orogen. Geological Society, London, Special Publications, 156, 427-456. JOLIVET, L. & FACCENNA, C. 2000. Mediterranean extension and the Africa-Eurasia collision. Tectonics, 19 (6), 1095-1106. JOLIVET, L., DANIEL, J. M. & FOURNIER, M. 1991. Geometry and kinematics of extension in Alpine Corsica. Earth and Planetary Sciences Letters, 104, 278-291. JOLIVET, L., BRUN, J. P., GAUTIER, P., LALLEMANT, S. & PATRIAT, M. 1994.3-D kinematics of extension in the Aegean from the Early Miocene to the Present, insight from the ductile crust. Bulletin de la Soci~t~ G~ologique de France, 165, 195-209. JOLIVET, L., FACCENNA, C. et al. 1998. Midcrustal shear zones in postorogenic extension: example from the northern Tyrrhenian Sea. Journal of Geophysical Research, 103, 12123-12160. KOCH, P. S. 1983. Rheology and microstructures of experimentally deformed quartz aggregates, PhD thesis, University of California. KOCH, P. S., CHRISTIE,J. C., ORD, A. & GEORGE, R. P. J. 1989. Effect of water on the rheology of experimentally deformed quartzite. Journal of Geophysical Research, 94, 13975-13996. KOHLSTEDT, D. L. & WEATHERS, M. S. 1980. Deformation-induced microstructures, paleopiezometers, and differential stresses in deeply eroded fault zones. Journal of Geophysical Research, 85, 6269-6285. LE PICHON, X. & CHAMOT-ROOKE, N. 1991. Extension of continental crust. In: MULLER, D. W., MCKENZIE, J. A. & WEISSERT, H. (eds) Controversies in Modern Geology. Academic Press Ltd, London 313-338. LISTER, G. S. & DAVIS, G. A. 1989. The origin of metamorphic core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, U.S.A. Journal of Structural Geology, 11, 65-94. LUAN, F. C. & PATERSON, M. S. 1992. Preparation and deformation of synthetic aggregates of quartz. Journal of Geophysical Research, 97, 301-320. MATTHEWS, A., LIEBERMAN, J., AVIGAD, D. & GARFUNKEL, Z. 1999. Fluid rock interaction and thermal evolution during thrusting of an Alpine

142

F. GUEYDAN ET AL.

metamorphic complex (Tinos Island, Greece). Contribution to Mineralogy and Petrology, 135, 212-224. MERCIER, J. C., ANDERSON, D. A. & CARTER, N. L. 1977. Stress in the lithosphere: inference from the steady state flow of rocks. Pure and Applied Geophysics, 115, 199-226. PARRA, T., VIDAL, O. & JOLIVET, L. 2002. Relation between the intensity of deformation and retrogression in blueschist metapelites of Tinos Island (Greece) evidenced by chlorite-mica local equilibria. Lithos, 63, 41-66. PATERSON, M. S. & LUAN, F. C. 1990, Quartzite rheology under geological conditions. In: KNIPE, R. J. & RUTTER, E. H. (eds) Deformation Mechanisms, Rheology and Tectonics. Geological Society, London, Special Publications, 54, 299-307. PATRIAT, M. & JOLIVET,L. 1998. Post-orogenic extension and shallow-dipping shear zones, study of a brecciated decollement horizon in Tinos (Cyclades, Greece). Comptes Rendus de l'Academie des Sciences de Paris, Sciences de la Terre et des Planetes, 326, 355-362. RANALLI, G. 1984. Grain size distributiuon and flow stress in tectonites. Journal of Structural Geology, 6, 443-447. RANALLI, G. 1995. Rheology of the Earth. Chapman & Hall, London. RANALLI,G. 2000. Rheology of the crust and its role in tectonic reactivation. Journal of Geodynamics, 30, 3-15. SERANNE, M. t~ SI~GURET, M. 1987. The Devonian basins of Western Norway: tectonics and kinematics of extending crust. In: COWARD, M. P., DEWEY, J. F. & HANCKOCK,P. L. (eds) Continental

Extensional Tectonics, Geological Society, London, Special Publication, 28, 537-548. STIPP, M. & TULLIS, J. 2003. The recrystallized grain size piezometer for quartz. Geophysical Research Letters, 30 (21), 2088, doi:10.1029/ 2003GL018444. TER HEEGE, J. H. 2002. Relationship between dynamic recrystallization, grain size distribution and rheology, PhD thesis, Universiteit Utrecht. Twlss, R. J. 1977. Theory and applicability of a recrystallized grain size paleopiezometer. Pure and Applied Geophysics, 115, 227-244. WEATHERS, M. S., BIRD, J. M., COOPER, R. F. & KOHLSTEDT, D. L. 1979. Differential stress determined from deformation-induced microstructures of the Moine thrust zone. Journal of Geophysical Research, 84, 7495-7509. WERNICKE, B. 1985. Uniform-sense normal simple shear of the continental lithosphere. Canadian Journal of Earth Sciences, 22, 108-125. WERNICKE, B. 1992, Cenozoic extensional tectonics of the U.S. cordillera. In: BURCHFIEL, B. C., LIPMAN, P. W. & ZOBACK, M. L. (eds) The Cordilleran Orogen: Conterminous U.S. Geological Society of America, Boulder, Colorado, 553-581. WERNICKE, B. 1995. Low-angle normal faults and seismicity: a review. Journal of Geophysical Research, I00, 20 159-20 174. WHITE, S. 1979a. Grain and sub-grain size variations across a mylonite zone. Contribution to Mineralogy and Petrology, 70, 193-202. WroTE, S. 1979b. Paleo-stress estimates in the Moine Thrust Zone, Eriboll, Scotland. Nature, 280, 222-223.

Dynamic recrystallization and strain softening of olivine aggregates in the laboratory and the lithosphere M A R T Y N R. D R U R Y

Department of Earth Sciences, Faculty of Geosciences, PO Box 80.021, 3508TA Utrecht, The Netherlands (e-mail martynd@ geo. uu. nl)

Abstract: The effects of dynamic recrystallization on the deformation mechanisms and rheology of olivine aggregates in the laboratory and the lithosphere are reviewed in this paper. The low-strain rheology of olivine is well documented; however, deformation in the lithosphere often involves large strains. Large strain experiments show that recrystallization can result in both hardening and softening during deformation. Moderate strain softening in experimental shear and torsion can be explained by the operation of dislocationaccommodated grain boundary sliding in bands of fine recrystallized grains. Data on the temperature dependence of recrystallized grain size are needed to extrapolate the effects of dynamic recrystallization to the lithosphere. Theories of dynamic recrystallization suggest that grain size is strongly stress dependent and moderately temperature dependent. A re-analysis of experimental grain size data indicates that the recrystallized grain size is temperature independent for olivine aggregates with low water content (< 300 ppm H/Si). Rheological regime maps have been constructed for the lithospheric mantle. The maps suggest that grain size sensitive power law creep, involving both grain boundary sliding and dislocation creep, will produce strong strain softening, greater than found so far in experimental studies, in dry and wet lithosphere shear zones.

Deformation of the lithospheric mantle will usually be controlled by olivine, which is the most abundant and weakest mineral in upper mantle rocks. In this paper the deformation mechanisms and rheology of olivine during low- and high-strain deformation are reviewed, concentrating on the role of recrystallization during deformation (dynamic recrystallization). The effect of dynamic recrystallization on deformation mechanisms and rheology depends on the grain size produced by recrystallization and how this grain size depends in turn on the deformation conditions. A new analysis is presented of the recrystallized grain size data for wet and dry olivine to investigate how the grain size depends on stress and temperature. The extrapolation of laboratory data to natural conditions is explored by construction of rheological regime maps appropriate to dry and hydrated lithosphere. Several recent papers have reviewed the experimental low-strain rheology of olivine (Karato & Wu 1993; Kohlstedt et al. 1995; Drury & Fitz Gerald 1998; Hirth 2002; Karato & Jung 2003; Hirth & Kohlstedt 2003). The results of the low-strain deformation studies will be briefly reviewed before going on to consider dynamic recrystallization and highstrain deformation.

Rheology of olivine aggregates in the laboratory Low-strain rheology The low-strain rheology of olivine single crystals and aggregates is now reasonably well documented. At high temperatures, the deformation of olivine can be described by a flow law of the form

= Alo'nd - m exp[-(E* + PV*)/RT]

(1)

where e is the strain rate, o- is the stress, d the grain size, T the temperature, P the pressure, n the stress exponent, m the grain size exponent, E* the apparent activation energy, V* the apparent activation volume, and R the gas constant. At high stress and lower temperature an exponential flow law applies (Goetze 1978). k = A2 exp{ - [(E* + PV*)/RT][1

-

(o'/O'p)] 2 } (2)

where A2 and O-p are constants. In the exponential creep field the flow strength may be grain size dependent with finer grained rocks being

From: GAPAIS, D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. DeformationMechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 143-158. 0305-8719/05/$15.00 cC~The Geological Society of London 2005.

146

M.R. DRURY

at shear strains between 0.1 and 1.1 (Fig. lb, c). The strain softening is accompanied by grain size reduction and the development of a strong CPO. The development of a strong CPO in large grains, the elongation of old grains, and measured stress exponents of 3.3 suggest that dislocation creep accommodates a significant amount of the deformation at high strains (Bystricky et al. 2000). However, the occurrence of weak CPO, low dislocation density and four grain junctions, suggestive of grain switching, indicate that significant grain boundary sliding occurred in the very fine recrystallized material of some highstrain samples (Lee et al. 2002). Recently, Holyoke & Tullis (2003) have performed highstrain direct shear and compression tests at lower temperatures (950-1100 ~ and they report much greater strain softening (23-70% weakening) than found at 1200-1300~ by Zhang & Karato (1995) and Bystricky et al. (2OOO). The experimental studies on single crystals and aggregates show that dynamic recrystallization has an important effect on the rheology of olivine, producing an increase of effective viscosity in single crystals at very high temperatures and an effective viscosity decrease in olivine aggregates deformed at 950-1300 ~ How can these different effects be rationalized and extrapolated to natural conditions? Deformation mechanism maps (Frost & Ashby 1982) are a useful way to assess how recrystallization may influence rheology during large strain deformation. Such maps are constructed from theoretical flow laws for particular deformation mechanisms constrained by experimental data. In the laboratory studies on olivine aggregates, empirical flow laws have been derived from experimental data and the association between rheological regimes and deformation mechanisms is not always established. Deformation maps based on empirical rheological flow laws should be termed rheological regime maps (Paterson 1987). As noted by Frost & Ashby (1982), the uncertainties in constructed deformation mechanism and theological regime maps are often large. Uncertainties in flow law constants translate into compounded uncertainties in the position of field boundaries in the map. When these uncertainties are extrapolated to natural conditions then uncertainties of several orders of magnitude in the position of field boundaries can occur. Another limitation is that the number of rheological regimes or deformation mechanisms included in the map may not be complete. Rheological regime maps for dry olivine constructed for experimental conditions showing

regimes for GSS-linear creep, GSI-power law creep and GSS-power law creep are shown in Figure 2. The maps are based mainly on the flow laws of Hirth & Kohlstedt (2003). The maps show that changes in grain size can produce transitions between rheological regimes. Transitions from GSI-power law creep to GSS mechanisms are of particular interest because significant strain softening can be produced if drastic grain size reduction is possible. Grain size reduction can produce softening by inducing a change from GSI-power law creep to GSSpower law creep (Hirth & Kohlstedt 1995b; Drury & Fitz Gerald 1998; Jin et al. 1998) or from GSI-power law creep to GSS-linear creep (Rutter & Brodie 1988). The viability of these transitions depends on the grain size produced during deformation, which in monophase materials at high temperatures is controlled by dynamic recrystallization (Twiss 1977). Rheological regime maps f o r dry olivine

In the very high temperature experiments of Karato et al. (1982) creep at low strains is controlled by slip on the [a](001) and [c](100) systems. After recrystallization, the strain rate becomes similar to the rate predicted for creep on the 'hard' [c](010) slip system. This is consistent with the idea that all slip systems in olivine must be activated so that general strains can be accommodated in olivine aggregates (Beeman & Kohlstedt 1993; Hirth & Kohlstedt 1995b). Thus, in this case the strain rate decrease associated with recrystallization is related to a transition from creep on weak slip systems to multiple slip on weak and strong slip systems. The rheological regime map constructed in Figure 2a for the Karato et al. (1982) experiments shows that the dynamic recrystallized grain size actually plots in the GSS-power law creep field. For example, at a creep stress of 50 MPa after recrystallization to a grain size of about 100 Ixm, the total strain rate is predicted to be 2.6 x 10 - 4 s - I , with 18% of the strain accommodated by GSI-power law creep and 82% of the strain accommodated by GSSpower law creep. The strain rate reduction caused by the change to multislip dislocation creep is partly offset by the increase in strain rate accommodated by grain boundary sliding in the medium-grained aggregate. The deformation map suggests that under these conditions recrystallization of a coarse aggregate would be associated with an increase of strain rate during creep by a factor of 4.4 or with softening by a factor of 1.5 during constant strain rate deformation.

DYNAMIC RECRYSTALLIZATION AND STRAIN sOFrENING IN OLIVINE AGGREGATES

high-strain experiments at high confining pressure have been c o n d u c t e d on olivine aggregates by Z h a n g & Karato (1995) and Bystricky et al. (2000). These constant strain rate experiments show that strain softening, with 1 5 - 2 0 % weakening, occurs after a peak stress is attained

natural strains of 2.0. In these creep experiments (Fig. la) the single crystals were converted to fine-grained aggregates by d y n a m i c recrystallization (Karato et al. 1980) and, importantly, recrystallization was associated with a strain rate decrease (Karato et al. 1982). Pioneering

Olivine 1650~ -2 ~ % 1

145

(Karato et al. 1982) .~

1100% 120MPa

=

-3

85 Mpa

9

76 MPa -4 0.5

1.0

1.5

strain Olivine 1200-1300~

(Zhang et al. 2000) 10%

200

1473K

50%

100

9

1573K

f/ I

L

I

1.0

2.0

shear strain Olivine 1200~

(Bystricky et al. 2000)

~ 200--

50%

1.2 x 10-4 6 x 10-5

95%

loo-

I

I

I

I

1.0

2.0

3.0

4.0

shear strain Fig. 1. Creep curves and stress-strain curves for high-strain experimental deformation of olivine. (a) Olivine single crystals deformed in uniaxial compression. During deformation the single crystals recrystallize to fine-grained aggregates. The volume fraction of recrystallization is shown by the dashed lines, with 1% marking the onset and 100% the completion of recrystallization. (b) Direct shear of olivine deformed at shear strain rates about l x 10 -5 s -1. The volume fraction of recrystallization is marked in %. At 1200 ~ recrystallization is limited and strain hardening occurs up to shear strains of 1.1. At higher temperatures recrystallization is more rapid and strain softening occurs at high strains. (c) Torsion experiments at 1200 ~ The volume fraction of recrystallization is marked in % along the stress-strain curve.

146

M.R. DRURY

at shear strains between 0.1 and 1.1 (Fig. lb, c). The strain softening is accompanied by grain size reduction and the development of a strong CPO. The development of a strong CPO in large grains, the elongation of old grains, and measured stress exponents of 3.3 suggest that dislocation creep accommodates a significant amount of the deformation at high strains (Bystricky et al. 2000). However, the occurrence of weak CPO, low dislocation density and four grain junctions, suggestive of grain switching, indicate that significant grain boundary sliding occurred in the very fine recrystallized material of some highstrain samples (Lee et al. 2002). Recently, Holyoke & Tullis (2003) have performed highstrain direct shear and compression tests at lower temperatures (950-1100 ~ and they report much greater strain softening (23-70% weakening) than found at 1200-1300~ by Zhang & Karato (1995) and Bystricky et al. (2OOO). The experimental studies on single crystals and aggregates show that dynamic recrystallization has an important effect on the rheology of olivine, producing an increase of effective viscosity in single crystals at very high temperatures and an effective viscosity decrease in olivine aggregates deformed at 950-1300 ~ How can these different effects be rationalized and extrapolated to natural conditions? Deformation mechanism maps (Frost & Ashby 1982) are a useful way to assess how recrystallization may influence rheology during large strain deformation. Such maps are constructed from theoretical flow laws for particular deformation mechanisms constrained by experimental data. In the laboratory studies on olivine aggregates, empirical flow laws have been derived from experimental data and the association between rheological regimes and deformation mechanisms is not always established. Deformation maps based on empirical rheological flow laws should be termed rheological regime maps (Paterson 1987). As noted by Frost & Ashby (1982), the uncertainties in constructed deformation mechanism and theological regime maps are often large. Uncertainties in flow law constants translate into compounded uncertainties in the position of field boundaries in the map. When these uncertainties are extrapolated to natural conditions then uncertainties of several orders of magnitude in the position of field boundaries can occur. Another limitation is that the number of rheological regimes or deformation mechanisms included in the map may not be complete. Rheological regime maps for dry olivine constructed for experimental conditions showing

regimes for GSS-linear creep, GSI-power law creep and GSS-power law creep are shown in Figure 2. The maps are based mainly on the flow laws of Hirth & Kohlstedt (2003). The maps show that changes in grain size can produce transitions between rheological regimes. Transitions from GSI-power law creep to GSS mechanisms are of particular interest because significant strain softening can be produced if drastic grain size reduction is possible. Grain size reduction can produce softening by inducing a change from GSI-power law creep to GSSpower law creep (Hirth & Kohlstedt 1995b; Drury & Fitz Gerald 1998; Jin et al. 1998) or from GSI-power law creep to GSS-linear creep (Rutter & Brodie 1988). The viability of these transitions depends on the grain size produced during deformation, which in monophase materials at high temperatures is controlled by dynamic recrystallization (Twiss 1977). Rheological regime maps f o r dry olivine

In the very high temperature experiments of Karato et al. (1982) creep at low strains is controlled by slip on the [a](001) and [c](100) systems. After recrystallization, the strain rate becomes similar to the rate predicted for creep on the 'hard' [c](010) slip system. This is consistent with the idea that all slip systems in olivine must be activated so that general strains can be accommodated in olivine aggregates (Beeman & Kohlstedt 1993; Hirth & Kohlstedt 1995b). Thus, in this case the strain rate decrease associated with recrystallization is related to a transition from creep on weak slip systems to multiple slip on weak and strong slip systems. The rheological regime map constructed in Figure 2a for the Karato et al. (1982) experiments shows that the dynamic recrystallized grain size actually plots in the GSS-power law creep field. For example, at a creep stress of 50 MPa after recrystallization to a grain size of about 100 Ixm, the total strain rate is predicted to be 2.6 x 10 - 4 s - I , with 18% of the strain accommodated by GSI-power law creep and 82% of the strain accommodated by GSSpower law creep. The strain rate reduction caused by the change to multislip dislocation creep is partly offset by the increase in strain rate accommodated by grain boundary sliding in the medium-grained aggregate. The deformation map suggests that under these conditions recrystallization of a coarse aggregate would be associated with an increase of strain rate during creep by a factor of 4.4 or with softening by a factor of 1.5 during constant strain rate deformation.

DYNAMIC RECRYSTALLIZATION AND STRAIN SOFTENING IN OLIVINE AGGREGATES 147 Dry Olivine 1650~ 1000

I

GSS-power law c r e e p ~

100

~r

-2

6G S ~ S _ I ~ "

10

100

1000

Grain size ~m

Dry Olivine 1227~ 1000

Gss-i}i~ creep' n ' 'l ..... ' ,,.,i,.____.._ ___~tJyI

100

~

.......

IaSI-powerlaw

10

b

1 10

100

1000

Grain size ~m Fig. 2. Rheological regime maps for dry olivine for high-temperature deformation in the laboratory (constructed with the flow laws of Hirth & Kohlstedt, 2003). Contours show the log strain rate. (a) T = 1650 ~ The filled circles are the dynamic recrystallized grain size in the recrystallized single crystals of Karato et al. (1982). The recrystallized grain size falls within the GSS-power law creep regime. The dotted lines show the initial strain rate in the single crystals. In this case the transition from creep on easy slip systems to polycrystal creep results in a decrease of strain rate at constant stress. (b) T = 1227 ~ Filled circles are dynamic recrystallized grain size from van de Wal et al. (1993) filled triangles from Zhang et al. (2000). The open triangles labelled with By and Zh show the initial grain size and flow stress of the experiments of Bystricky et al. (2000) and Zhang et al. (2000) with the arrows illustrating the amount of grain size reduction and changes of stress during deformation.

A rheological regime map for high temperature torsion and shear experiments is shown in Figure 2b. Data for recrystallized grain size in dry olivine (van der Wal et al. 1993; Zhang et al. 2000) are plotted on the deformation map as well as the grain size variations found by Bystricky et al. (2000) and Zhang et al. (2000) in their high-strain experiments. The recrystallized grain size plots close to the transition from power law creep to linear creep. This is roughly consistent with the hypothesis of de Bresser et al. (1998) that dynamic recrystallized grain size is controlled by a balance between grain-size reduction in the dislocation creep regime and grain growth in the GSS-linear creep regime. Mont6si & Hirth (2003) have shown that the 20% strain softening observed by Bystricky et al. (2000) can be accounted for by an increased component of GSS-linear creep. However, the rheological regime map (Fig. 2b) suggests that a transition from GSI-creep to GSS-linear creep does not occur at high temperatures. The deformation mechanism maps clearly show that the moderate softening observed in experiments can be explained by a transition to GSS-power law creep. Lee et al. (2002) have found key evidence for significant grain boundary sliding in fine-grained recrystallized bands in samples deformed at 1200 ~ to high shear strains. In fact, the amount of softening predicted from the map, a 50% reduction in flow stress at constant strain rate (Fig. 2b), overestimates the observed weakening of 20% found by Bystricky et al. (2000). Thus, it is not necessary to invoke GSS-linear creep to explain the strain softening in high-strain olivine experiments. The weakening can be accounted for by grain size reduction in the GSS-power law creep regime. It is important to note that the high-strain strength found by Bystricky et al. (2000) is comparable to finegrained aggregates deformed to lower strains. This suggests that the softening is indeed related to grain size reduction and not to 'geometric softening', that is, the development of a softer lattice preferred orientation. In recent lower temperature experiments (9501100 ~ Holyoke & Tullis (2003) report a transition from climb-accommodated dislocation creep at low stress to recrystallization accommodated dislocation creep at high stress. In both regimes, significant strain softening was associated with recrystallization. The strain softening at low temperature and low stress may be explained by the occurrence of GSS-power law creep as found at high temperatures ( T > 1200 ~ Holyoke & Tullis (2003) suggest that the strain softening at low temperatures and high stress is

148

M.R. DRURY

related to the production of bands of soft, dislocation-free, recrystallized grains. At present there are insufficient rheological data to include a rheological regime associated with recrystallization accommodated dislocation creep in deformation maps. To a first approximation, recrystallization accommodated dislocation creep may occur mainly in the exponential creep regime.

Olivine (250 ppm H/Si) 1300~

0.3GPa

104 Gss

[,,

........

cree;I

1000

X~]

........

,

.......

GSI-power law creep

~

IK2]'~N,~x,

,

ioo

-5

Rheological regime maps f o r wet olivine It has been known for many years that water has a strong effect on the strength of olivine (Chopra & Paterson 1981) but it is only recently that good constraints on the quantitative effect of water on rheology have been derived (Mei & Kohlstedt 2000a, b; Karato & Jung 2003; Hirth & Kohlstedt 2003). Rheological regime maps are constructed for olivine at low (250 ppm) and high (1400 ppm) water content corresponding to the experiments of Chopra & Paterson (1981) (Fig. 3a) and Jung & Karato (2001a, b) (Fig. 3b). In both cases, under experimental conditions the recrystallized grain size falls in the GSI-power law creep regime in the neighbourhood of the transition to GSS-linear creep. The maps for wet olivine in Figure 3 are based on the Hirth & Kohlstedt (2003) flow laws. Another map is constructed in Figure 4 for the low water content experiments, this timeoUSing the GSI-power law for Anita Bay and Aheim dunite (Chopra & Paterson 1981) and the GSSlinear creep flow law of Hirth & Kohlstedt (2003). The dynamic recrystallized grain size in this map plots close to the transition between GSI-power law creep and GSS-linear creep (de Bresser et al. 2001). Various data points for the strength of different samples as a function of grain size are also plotted in Figure 4. The strength of fine-grained samples (40-100 txm) is clearly lower than coarse-grained (900 ~m) polycrystals and [011]c orientation single crystals. This apparent grain size dependence of strength in the dislocation creep implicates a role for grain boundary sliding in the strength of wet polycrystals (Chopra & Paterson 1981; Chopra 1986). Hirth & Kohlstedt (1996, 2003), however, have suggested that the olivine in Aheim dunite has a lower water content, thus explaining the higher strength compared to Anita Bay dunite. Mei & Kohlstedt (2000b) concluded that the strength of wet olivine in the dislocation creep regime was controlled by the easy [a](010) slip system. If so, then it is difficult to understand how strain incompatibility between adjacent grains can be accommodated by single slip

10

-

l

t

,

,,I,,,I

~

l 111+11

10

a

,

I,

,,X,,,,l

100

,

,

....

1000

104

Grain size ~m Olivine (1400 ppm H/Si) 1250~ 10 4 ~

........

~

........

.•

~

........

t

2GPa .......

GSI-power law creep

1000

t ]

ow

~

100

10

1

b

10

100

1000

104

Grain size pm

Fig. 3. Rheological regime maps for wet olivine for high-temperature deformation in the laboratory (constructed with the flow laws of Hirth & Kohlstedt 2003). (a) Low water content and 1300 ~ Filled circles are recrystallized grain size for wet olivine (van der Wal et al. 1993). The recrystallized grain size falls in the GSI-power law creep field but close to the field boundary with GSS-linear creep. Sample JK21 in the GSS-power law creep regime has a type-B olivine CPO (Jung & Karato 2001b). (b) High water content, 1250 ~ 2 GPa. Filled circles are recrystallized grain size from Jung & Karato (2001a).

without the operation of other mechanisms such as grain boundary diffusion and sliding. If GBS occurs then there should be some grain size dependence of the flow strength (Gifkins 1976). Evidence for the operation of GBS in wet

DYNAMIC RECRYSTALLIZATION AND STRAIN SOFTENING IN OLIVINE AGGREGATES Olivine (250ppm H/Si) 1300~ 0.3GPa 10 4

.....

I

. . . . . . . .

AB

1000

\

I

. . . . . . . .

10.6

/

10.5

-3 (AB&Ah) ,[011]c

9 /

..........................

.... ?..,k,2

,

. . . . . . .

GSI-power law creep

\

~x)~ lOO

I

o

.........

0.0001 0.001 Grain size him

0.01

Fig. 4. Rheological regime maps for wet olivine for high-temperature deformation in the laboratory (constructed with the flow laws from Chopra & Paterson 1981, and Karato et al. 1986). The field boundary between GSI-power law creep and GSS-linear creep and strain rate contours are shown for the Aheim dunite GSI-flow law and the low-stress Anita Bay dunite flow law of Chopra & Paterson (1981). Filled circles are recrystallized grain size (van der Wal et al. 1993). With both flow laws the recrystallized grain size falls along the field boundary between GSS and GSI creep. The unfilled circles are mechanical data from Chopra & Paterson (1981), Karato et al. (1986) and Mackwell et al. (1985).

dunites has been reported by Fitz Gerald & Chopra (1984). Furthermore, the occurrence of grain boundary alignments in the wet olivine samples of Karato et al. (1986, fig. 10) provide indirect evidence for the operation of GBS (Drury & Humphreys 1988) in wet samples deforming by dislocation creep. If the strain rate from dry GSS-power law creep is compared with the wet flow laws of Hirth & Kohlstedt (2003) (Fig. 3a), then dry GSS-power law creep is faster than wet GSI-power law creep and GSS-linear creep at high stress and small grain size when T = 1300 ~ Experiments on single crystals show that water increases the creep rate on the [a](010) system (Mackwell et al. 1985) and water increases the rate of grain boundary migration and diffusion (Karato 1989). So both grain boundary sliding and its accommodation by slip on [a](010) should be enhanced by water; therefore, a dislocation accommodated GBS creep regime should also occur in wet olivine. Further studies on fine to coarse olivine polycrystals with well-controlled water content are needed to investigate this problem.

149

Jung & Karato (2001a) have recently shown that different types of crystallographic preferred orientation (CPO) develop in high-strain wet olivine compared to dry olivine. In dry olivine, the type-A CPO is defined by a strong [a] maximum subparallel to the shear direction, whereas in wet olivine type-B CPO and type-C CPOs have a [c] axis maximum subparallel to the shear direction. The occurrence of the different CPOs in wet olivine are shown in the deformation mechanism maps for wet olivine in Figure 3. It is interesting that the type-B CPO is found in a sample JK21, which was deformed under conditions where dry GSS-power law creep is faster than water-enhanced GSI-power law creep and GSS-linear creep. The microstructures in samples with type-B CPO are consistent with significant deformation by grain boundary sliding (Jung & Karato 2001a). It is important to understand the micro-physical processes involved in strain softening so that the amount of weakening under natural conditions can be estimated. The scaling law between dynamic recrystallized grain size and deformation conditions along with the flow laws for the different deformation mechanisms are the key information that must be known for reliable extrapolation from laboratory experiments to the lithosphere. Our knowledge on flow laws for different theological regimes in olivine is now quite detailed, but much less is known about the controls of dynamic recrystallized grain size.

Dynamic recrystallized grain size in wet and dry olivine It is well known that the dynamic recrystallized grain size is strongly dependent on the flow stress (Twiss 1977) and several experimental studies have investigated scaling laws between recrystallized grain size and deformation conditions in olivine (Ross et al. 1980; Karato et al. 1980; van der Wal et al. 1993; Zhang et al. 2000; Jung & Karato 2001b). D -- gl o"-p

(6)

where D is the recrystallized grain size, and K1 and p are constants. Several theoretical models have been proposed for recrystallized grain size (Derby & Ashby 1987; de Bresser et al. 1998; Shimuzu 1998) and these all predict that the grain size is temperature dependent with D = K2 0 " - q e x p ( A Q / R T )

(7)

where K2 and q are constants and AQ is the difference in temperature dependence between

150

M.R. DRURY

the grain size reduction process and the grain growth process. Using a metal alloy as a rock analogue, de Bresser et al. (1998) detected the temperature dependence of the recrystallized grain size in a careful set of experiments at constant stress with variable temperature. It is important to determine if the recrystallized grain size in olivine is also temperature dependent so that the experimental scaling laws can be properly extrapolated to natural conditions. The existing data set for recrystallized grain size in olivine shows no obvious temperature dependence (Karato et al. 1980; van der Wal et al. 1993); however, the data set is biased in the sense that high-stress experiments were conducted at low temperatures and vice versa. Furthermore, recent studies have shown that the dynamic recrystallized grain size varies with water content (Jung & Karato 2001b). The recrystallized grain size data for olivine are re-examined here to specifically test for any temperature dependence.

T e m p e r a t u r e d e p e n d e n c e o f recrystallized grain size

Recent studies have shown that recrystallized grain size in olivine is strongly dependent on stress and also shows some dependence on water content (Jung & Karato 2001b). Zhang et al. (2000) have also shown that the grain size developed in samples deforming by GSSpower law creep is smaller than the grain size found in coarse olivine rocks deformed at the same stress (van der Wal et al. 1993). Before these variations are discussed, the methods of grain size measurement in each study must be considered. Van der Wal et al. (1993) measured linear intercepts in recrystallized areas of wet olivine samples and the mean linear intercept was multiplied by a factor of 1.75 to convert the mean linear intercept to an average grain diameter. The factor of 1.75 applies to grains with a polygonal shape (Pickering 1976). For the dry samples, reported by van der Wal et al. (1993), with limited recrystallization, grain diameters were measured and the mean diameter was multiplied by 1.2 to account for the sectioning of spherical grains (van der Wal 1993). Zhang et al. (2000) measured linear intercepts and obtained an average grain diameter using a factor of 1.2 times the mean linear intercept. Thus, some of the variation of grain sizes found between the different studies is caused by the use of different correction factors. In the deformation mechanism maps presented in this paper the grain size data from Zhang et al. (2000)

have been multiplied by a factor of 1.75/1.2 to allow direct comparison with the data from van der Wal et aL (1993). To investigate the possibility that the recrystallized grain size in olivine is temperature dependent, the wet and dry grain size data from van der Wal et al. (1993) and Karato et al. (1980) have been reanalysed. The stress dependence of recrystallized grain size can be determined from experiments performed at the same temperature. Data for Aheim dunite at 1200 ~ and Anita Bay dunite at 1200 ~ and recrystallized single crystals at 1650 ~ provide values of the stress exponent of grain size between 1.14 and 1.44. With the isothermal stress exponent q the grain sizes can be normalized to a constant stress. A plot of the IOgeof the normalized grain size (D*) against l I T can be used to evaluate the temperature dependence in Equation (7) (Fig. 5). The variation of D* in wet Aheim dunite between 1190 and 1300 ~ is consistent with a small temperature dependence for recrystallized grain size with AQ = 26 kJ/mol. As predicted from the grain size models at constant stress the recrystallized grain size decreases with temperature. In contrast, in wet Anita Bay dunite, between 1100 and 1300 ~ AQ = - 2 7 . 4 k J / m o l ; that is, at constant stress a larger recrystallized grain size occurs at a higher temperature. The temperature dependence of recrystallized grain size in dry olivine can be evaluated by comparing the data at 1650 ~ from Karato et al. (1980) with the 1300 ~ data from van der Wal et al. (1993). The dry data give a value of AQ = - 2 6 . 3 kJ/mol, similar to the estimate for wet Anita Bay dunite. So a small temperature dependence of the recrystallized grain size in olivine is possible but the available data indicate that AQ < 30 kJ/mol. Taking all of the dry and wet data together, and taking into account uncertainties, a conservative value of AQ = 0 + 50 kJ/mol is consistent with the current data for dry olivine and wet olivine with low water content <300 ppm H/Si. The activation energy term for grain size in wet Anita Bay dunite and dry olivine is actually negative at around - 3 0 kJ/mol. A negative activation energy term is possible if the temperature dependence of grain size reduction is less than that of grain growth. Likewise the recrystallized grain size will be temperature independent if the activation energies for grain size reduction and grain growth are equal (de Bresser et al. 2002). The weak temperature dependence of recrystallized grain size in olivine may be explained if grain size reduction is controlled by slip on the easy [a](010) system (385-400 kJ/mol)

D Y N A M I C R E C R Y S T A L L I Z A T I O N A N D S T R A I N S O F I ' E N I N G IN O L I V I N E A G G R E G A T E S Wet Aheirn dunite Qd=27 kJ/mol -9

o

-10

45.4 gm

.......................................................................................................................................................................................

N

..=

9 1

O

-ll

I

~1,

mi ~i

-

.............................!.............................................................. -..............................--"..............................i ............................ 16.7 tam

o -12 ............................................................................................ i..........................................................................................6.141am

-13 0.00063

I

I

0.00064

0.00065

0.00066

I

0.00067

I

0.00068

0.00069

1/T

Anita Bay dunite Qd=-27 kJ/mol

*

-10

45.4 lam

-11

16.7 lam

-12

6.141am

Z" e~ o ,--1

b

-13 0.00062

0.00064

0.00066

0.00068

0.0007

0.00072

0.00074

1/T

-9

D r y o l i v i n e Q d = -26 k J / m o l , i

1

-10

45.4 ~tm

16.7 lam

O

9

,.d -12 -

C

-13 0.0005

6.14~tm

i 0.00052 0.00054 0.00056 0.00058

l I 0.0006 0.00062 0.00064

1/T Fig. 5. The loge of normalized recrystallized grain size plotted against reciprocal temperature for (a) wet ,a,heim dunite, (b) wet Anita Bay dunite and (e) dry dunites and recrystallized single crystals. A very small temperature dependence of recrystallized grain size is indicated in each case.

151

152

M.R. DRURY

and growth is controlled by grain boundary diffusion (315-380 kJ/mol). The recrystallization models (Derby & Ashby 1987; Shimizu 1998; de Bresser et al. 1998) would then predict AQ = 1-40 kJ/mol. In summary, reanalysis of the recrystallized grain size data for olivine with low water content over the temperature range between 1100 and 1650 ~ suggests that the temperature dependence of recrystallized grain size is small with AQ = 0 + 50 kJ/mol. Based on this result the recrystallized grain size will be taken as temperature independent so that the high-temperature scaling law for grain size can be directly extrapolated to lower temperatures in the lithosphere.

Extrapolation from the laboratory to the lithosphere If the flow laws for different deformation mechanisms and the scaling law for dynamic recrystallization are known, then deformation mechanism maps can be used to explore the high-strain deformation mechanisms and theology of the lithosphere (Karato et al. 1986; Rutter & Brodie 1988; Drury & Fitz Gerald 1998; Hirth 2002). Oceanic lithosphere formed at mid-ocean spreading ridges is highly depleted by melt extraction and should therefore be dry (Hirth & Kohlstedt 1996). The water content of continental lithosphere is expected to be more variable. Continental cratonic lithosphere formed by accretion of asthenosphere diapirs (de Smet et al. 1998) should be dry owing to extensive melting (Karato 2003), but as continental lithosphere ages it can become progressively hydrated by infiltrating fluids. It is likely that hydrous regions such as the mantle wedge above subduction zones and hydrated lithospheric shear zones are strongly water weakened (Karato & Wu 1993; Kohlstedt et al. 1995; Hirth & Kohlstedt 2003). A potential problem in extrapolation from the laboratory to the lithosphere arises from the possibility that dynamic recrystallization mechanisms may change with temperature and stress. In low stacking fault energy (SFE) metals there is a lower stress limit to dynamic recrystallization (Twiss & Sellars 1978) and experimental data on recrystallization in olivine are limited to high stress so it is possible that dynamic recrystallization may not be significant in the lithosphere if olivine behaves in the same way as low SFE metals. The lower stress limit for dynamic recrystallization in metals applies only to discontinuous recrystallization involving

the development and growth of dislocationfree grains. This mechanism is usually called bulging recrystallization in the geological literature (Drury et al. 1985; Hirth & Tullis 1992; See De Meer et al. 2002, p. 5-6, for a comparison of recrystallization mechanisms in metals and minerals). While there is evidence for a stress limit for bulging recrystallization in minerals (Hirth & Tullis 1992; Holyoke & Tullis 2003), other recrystallization mechanisms such as rotation recrystallization and continuous migration recrystallization occur at lower stress (Drury et al. 1985; Hirth & Tullis 1992). Indeed, microstructures in naturally deformed peridotites suggest that dynamic recrystallization is important in olivine even at the low stress levels in the asthenosphere (Dijkstra et al. 2002).

R h e o l o g i c a l regime m a p s f o r dry lithosphere

Rheological regime maps for dry olivine are relevant for the oceanic lithosphere (Hirth 2002) and highly depleted portions of the continental lithosphere. Hirth (2002) has constructed a series of theological regime maps for dry olivine based on the set of self-consistent flow laws of Hirth & Kohlstedt (2003). As shown in Figure 2, at high temperatures GSS-power law creep is restricted to high stresses, so this rheological regime is not predicted to be significant in the deep lithosphere and asthenosphere. In contrast, rheological regime maps for 600-800 ~ (Fig. 6) show that GSS-power law creep is the dominant mechanism over a broad range of conditions. At 800 ~ (Fig. 6a) the recrystallized grain size plots along the field boundary between GSS-power law creep and GSS-linear creep. The map shows that significant softening can be produced by grain size reduction. For example, at a constant strain rate of 1 x 10 -1~ s -~ a factor 5 reduction in flow stress will be associated with recrystallization to 10 ~m from an initial grain size of 1000 ~,m. Note that this is a much greater amount of softening than found under experimental conditions. In this case there is actually a transition from exponential creep to GSS-power law creep with decreasing grain size. This transition should be particularly favourable for the development of localized shear zones, because material in the shear zone will soften while the matrix outside the band strain hardens (Evans & Goezte 1979). The prediction of increasing strain softening with decreasing temperature is supported by the recent results of Holyoke & Tullis (2003) who found a stress

DYNAMIC RECRYSTALLIZATION AND STRAIN SOFTENING IN OLIVINE AGGREGATES Dry Olivine 800~ 104

Exponential creep

-5

1000

~ L

creep l

power law

r~ 10

1 10

100

1000

104

Grain size ~tm Dry Olivine 600~ 10 4

-5

Exponential creep

1000

100

PL

ra~

10

1 10

100

1000

104

Grain size pm Fig. 6. Rheological regime maps for dry lithosphere, constructed with the flow laws of Hirth & Kohlstedt (2003). (a) 800 ~ The extrapolated recrystallized grain size data for dry olivine plot close to the field boundary between GSS-power law creep and GSS-linear creep. In consequence, recrystallization from a coarse initial grain size (illustrated by grey arrow) will result in a factor 5 reduction in flow stress at constant strain rate. 600 ~ (b) even at very low temperatures the extrapolated recrystallized grain size data plot close to the field boundary between GSS-power law creep and GSS-linear creep.

reduction of 70% during high-strain direct shear tests at 1050 ~ in the climb-controlled dislocation creep regime. The rheological regime maps for dry lithosphere do not include a regime associated with

153

recrystallization accommodated dislocation creep (Tullis & Yund 1985; Hirth & Tullis 1992). This regime is likely to occur at higher strains in the exponential creep regime and structural softening caused by recrystallization would result in the expansion of the exponential creep regime at the expense of GSS-power law creep and GSI-power law creep. In consequence, recrystallization accommodated dislocation creep could be an important softening and localization mechanism in the shallow mantle. The GSS-power law creep flow law is not as well constrained as other flow laws for dry olivine so it is instructive to examine the effect of variations in the flow law parameters on extrapolation to natural conditions. Estimates for the activation energy for GSS-power law in dry olivine range from 3 8 5 k J / m o l for T = 1100-1300 ~ (Drury & Fitz Gerald 1998) to 400 kJ/mol at T < 1250 ~ and 600 kJ/mol at T > 1250 ~ (Hirth & Kohlstedt 2003). Mei & Kohlstedt (2000a) report an activation energy of 510 kJ/mol for fine-grained dry olivine samples between 1200 and 1300 ~ If the activation energy for GSS-power law creep is lower than that for GSI-power law creep then the amount of softening produced by a transition from dislocation creep to GSS-power law creep will be higher under natural conditions than found in the laboratory. Further, if the activation energy for GSS-power law creep ( 3 8 5 4 0 0 k J / m o l ) is indeed similar to that of GSS-linear creep then the transition between the two mechanisms will occur under similar conditions in nature as found in the laboratory. Theoretical flow laws for GBS creep predict a stress dependence of n = 2 - 3 , and a grain size exponent m = 1 - 2 (Gifkins 1976). Current data for olivine suggest n = 3.5 and m --- 2 (Hirth & Kohlstedt 2003). The influence of small changes in n and m on the field boundaries for a rheological regime map are shown in Figure 7a. With n = 3.5 and m = 2 the GSSpower law creep regime is restricted to highstress low-temperature conditions in the shallow lithosphere. Changing n from 3 to 2 results in the extension of the GSS-power law creep regime to lower stress levels and if n = 2 and m = 2 then the GSS-power law regime occurs as an intermediate regime separating the GSI power law creep and GSS-linear creep fields at high and low stress. So, based on the current constraints for GSS-power law creep this regime should be restricted to relatively low temperatures in dry lithosphere; however, GSS-power law creep could be important at higher temperatures if the stress exponent is less than currently estimated.

154

M.R. DRURY (a)

Dry Olivine 1300~ field boundaries.

1000

100

r~

lO

10

100

1000

Grain size ram Olivine (250ppm H/Si) 800~

(b) 104

'

'

' ' ' " ' 1

,

' ' ' " " 1

. . . .

'"'1

'

'

'

....

1000

g,

b

100

10

1

10.6

10.5

0.0001

0.001

0.01

Grain size jam Fig. 7. Rheological regime maps illustrating the large uncertainty in the position of rheological regime field boundaries in dry and wet olivine. (a) Map for dry olivine at 1227 ~ showing fields for GSI creep, GSS creep and two alternative fields for GSS-power law creep depending on the value of the stress exponent (n).

(b) Map for wet olivine at typical lithospheric temperature. Filled circles are extrapolated recrystallized grain size data for wet olivine. The field boundary between GSS and GS! creep is shown for different GSI flow laws (,~h, ABhi, ABlo from Chopra & Paterson 1981, KPF from Karato et al. 1986).

Rheological regime maps for hydrated lithosphere If the available flow laws and grain size scaling laws for wet olivine are extrapolated to the

lower temperatures in hydrated lithosphere then rheological regime maps are obtained (Fig. 7b) that suggest that recrystallization will take the grain size into the dominant GSS-linear creep regime (Rutter & Brodie 1988). Before discussing the viability of this type of transition, the sensitivity of the extrapolation to lower temperatures will be examined. Figure 7b shows the variation of the transition between GSI-power law creep and GSS-linear creep extrapolated to 800 ~ based on different GSI-power law creep flow laws. It is clear that variations in the flow law parameters produce a huge variation in the relative position of the recrystallized grain size data and the rheological field boundary. In fact, taking the low-stress dislocation creep law of Chopra & Paterson (1981) for Anita Bay dunite obtained from stress relaxation experiments the recrystallized grain size data plot fight along the field boundary. At high temperatures in the GSS-linear creep regime a stress exponent n = 1 and a grain size exponent m = 3 are consistent with Coble creep (Hirth & Kohlstedt 2003) or diffusioncontrolled grain boundary sliding. The only low-temperature data on GSS creep is from McDonnell et al. (1999, 2000) who deformed forsterite at lithospheric temperatures of 8501000~ They found GSS deformation with n = 1.7-2.4 and m = 3, and they concluded that deformation occurred by water-enhanced grain boundary sliding accommodated by diffusion and/or dislocation activity. If GSS creep in the lithosphere also occurs by a similar mechanism then the slope of the field boundary between dislocation creep and GSS creep will be steeper than found at high temperatures. This illustrates the large uncertainties involved in extrapolating high-temperature field boundaries to lithospheric temperatures.

Viability of a recrystallization-induced transition to GSS-linear creep The idea that recrystallization could produce a transition to GSS-linear creep (White 1976; Twiss 1976) was quickly questioned by Etheridge & Wilkie (1979), who proposed that the recrystallized grain size stress scaling law must always lie within the dislocation creep because dislocation creep is an essential process in the recrystallization processes. Karato et al. (1986) produced the first experimental data for GSS creep in olivine and they found recrystallization did not produce a transition to diffusion creep at high temperatures. It was Rutter & Brodie (1988) who showed that

DYNAMIC RECRYSTALLIZATION AND STRAIN SOFTENING IN OLIVINE AGGREGATES 155 if the wet olivine data are extrapolated to lowtemperature lithospheric conditions then the recrystallized grain size scaling law apparently extends into the GSS-linear creep regime. Handy (1989) pointed out that such a fine grain size would not be stable against surface energy driven grain growth and GSS creep experiments on olivine clearly show this time-dependent hardening (Karato et al. 1986). de Bresser et al. (1998) have proposed that the dynamic recrystallized grain size is controlled by a balance between grain size reduction in the dislocation creep regime and grain growth in the diffusion creep regime. In consequence, the stress recrystallized grain size scaling law always plots in the neighbourhood of the mechanism transition from dislocation creep to diffusion creep (de Bresser et al. 1998, 2001) and a transition to dominant diffusion creep should never occur. In other models of dynamic recrystallization a stable grain size is maintained by a balance between grain growth and grain size reduction processes with grain size reduction controlled by the rate of dislocation creep (Derby & Ashby 1987; Shimizu 1998). These models do not consider the role of grain boundary deformation mechanisms but they do suggest that the recrystallization scaling law is independent and not coincident with the mechanism boundary as proposed by de Bresser et al. (1998). Thus, if this type of model is valid, the relative position of the rheological field boundary and grain size scaling law will vary with both stress and temperature (see also Montrsi & Hirth 2003). In fact, Lee et aL (2002) have found evidence that the contribution of grain boundary sliding deformation varies with temperature in fully recrystallized olivine. They found weak CPOs and microstructures indicating grain neighbour switching in olivine samples deformed at 1200 ~ but at higher temperature 1300 ~ recrystallized samples have a very strong CPO and grain switching microstructures were absent. It follows that there may be some conditions where the recrystallization scaling law extends partly into the GSS-linear creep regime and in this situation dynamic recrystallization may produce a transient transition to dominant GSSlinear creep. Any transient transition to dominant diffusion creep would be limited by grain growth (Handy 1989; Drury et aL 1991), which will produce hardening and may result in a transition back to dominant GSI or GSS power law creep, de Bresser et al. (2001) suggest that the subgrain size-stress relationship will form the lower grain size limit to possible grain size reduction in the GSS deformation field. Thus, it can be argued that a transient transition to

dominant GSS-linear creep induced by recrystallization is possible under some circumstances but current data suggest that this is unlikely in dry olivine and very uncertain in wet olivine. Mont~si & Hirth (2003) have shown that a transition to GSS-linear creep is possible during post-seismic creep events in fine-grained shear zones. The occurrence of a recrystallization induced transition to GSS-linear creep is less likely if GSS-power law creep or recrystallization accommodated dislocation creep (Tullis & Yund 1985) are important in wet olivine. The high-strain shear experiments on wet olivine (Jung & Karato 2001a, b) do indeed strongly suggest that recrystallization results in a transition to deformation involving cooperative grain boundary sliding, grain boundary migration and dislocation creep. If GSS-power law creep is important in wet olivine then recrystallization will most probably result in a transition from GSI-power law to GSS-power law creep as found for dry olivine (Fig. 6). Further studies on the role of grain boundary sliding and recrystallization in the deformation of wet olivine would help to resolve this issue.

Conclusions (1) The moderate strain softening observed in high-strain, high-temperature (12001300~ olivine experiments can be explained by grain size reduction in the GSS-power law creep regime. (2) A regime of GSS-power law creep probably occurs in wet olivine under experimental conditions but further studies on coarse olivine with well-controlled water content are needed. (3) An analysis of recrystallized grain sizestress data for wet and dry olivine indicates that the temperature dependence of recrystallized grain size is small and it is reasonable to assume that grain size is temperature-independent in olivine aggregates with low water content. (4) Extrapolation of current flow laws and the temperature-independent recrystallization scaling law to conditions in the lithosphere shows that GSS-power law creep will produce larger degrees of strain softening in dry lithosphere than found under experimental conditions. The degree of softening predicted is sensitive to the parameters of the GSS-power law creep flow law, which are not well constrained. (5) Extrapolation of wet olivine laboratory data to conditions in hydrated lithosphere

156

M.R. DRURY

suggests that a recrystallization-induced transient transition to GSS-linear creep may occur under limited conditions. It is more likely, however, that recrystallization in wet olivine will produce a similar transition as found in dry olivine, from GSI-power law creep to GSS-power law creep.

J. H. P. de Bresser and C. J. Spiers are thanked for discussions. G. Hirth kindly provided a pre-print of his paper on olivine flow laws. M. S. Paterson and G. Ranalli are thanked for helpful reviews. This research was supported by an NWO-PIONIER subsidy.

References BAI, Q., MACKWELL,S. J. 8z KOHLSTEDT,D. L. 1991. High temperature creep of olivine single crystals 1. Mechanical results for buffered samples. Journal of Geophysical Research, 96, 2441-2463. BEEMAN, M. L. & KOHLSTEDT, D. L. 1993. Deformation of fine-grained aggregates of olivine plus melt at high temperatures and pressures. Journal of Geophysical Research, 98, 6443-6452. BYSTRICKY, M., KUNZE, K., BURLINI, L. & BURG, J.-P. 2000. High shear strain of olivine aggregates: rheological and seismic consequences. Science, 290, 1564-1567. CHOPRA, P. N. 1986. The plasticity of some finegrained aggregates of olivine at high pressure and temperature. In: HOBBS, B. E. & HEARD, H. C. (eds) Mineral and Rock Deformation. Geophysical Monographs, 36, 25-34. CHOPRA, P. N. 8z PATERSON, M. S. 1981. The experimental deformation of dunite. Tectonophysics, 78, 453-473. CHOPRA, P. N. 8z PATERSON, M. S. 1984. The role of water in the deformation of dunite. Journal of Geophysical Research, 89, 7861-7876. DE BRESSER, J. H. P., PEACH, C. J., REIJS, J. P. J. & SPIERS, C. J. 1998. On dynamic recrystallization during solid state flow: Effects of stress and temperature. Geophysical Research Letters, 25, 3457-3460. DE BRESSER,J. H. P., TER HEEGE, J. H. & SPIERS, C. J. 2001. Grain size reduction by dynamic recrystallization: can it result in major theological weakening? International Journal of Earth Sciences, 90, 28 -45. DE BRESSER, J. H. P., EVANS, B. & RENNER, J. 2002. On estimating the strength of calcite rocks under natural conditions. In: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. & PENNOCK, G. M. (eds) Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 309-329. DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P & PENNOCK, G. M. 2002. Current issues and new developments in deformation mechanisms,

rheology and tectonics. In: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. & PENNOCK, G. M. (eds) Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 1-27. DE SMET, J. U., VAN DEN BERG, A. P. & VLAAR, N. J. 1998. Stability and growth of continental shields in mantle convection models including recurrent melt production. Tectonophysics, 296, 15-29. DERBY, B. & ASHBY, M. F. 1987. On dynamic recrystallisation. Scripta Metallurgica, 21, 879-884. DIJKSTRA, A. H., DRURY, M. R. & FRIJHOFF, R. M. 2002. Microstructures and lattice fabrics in the Hilti mantle section (Oman ophiolite): Evidence for shear localization and melt weakening in the crust-mantle transition zone. Journal of Geophysical Research, 107 (Bll), 2270, doi:10.1029/ 2001JB000458. DRURY, M. R. & HUMPHREYS,F. J. 1988. Microstructural shear criteria associated with grain boundary sliding during ductile deformation. Journal of Structural Geology, 10, 83-89. DRURY, M. R. & FITZ GERALD, J. D. 1998. Mantle rheology: Insights from laboratory studies of deformation and phase transition. In: JACKSON, I. N. S. (ed) The Earth's Mantle - Composition, Structure and Evolution. Cambridge University Press, Cambridge, UK, 503-559. DRURY, M. R., HUMPHREYS, F. J. t~ WHITE, S. H. 1985. Large strain deformation studies using polycrystalline magnesium as a rock analogue: Part II dynamic recrystallization mechanisms at high temperatures. Physics of the Earth and Planetary Interiors, 40, 208-222. DRURY, M. R., VISSERS, R. L. M., HOOGERDUIJN STATING, E. n. • VAN DER WAL, D. 1991. Shear localisation in upper mantle peridotites. Pure and Applied Geophysics, 137, 439-460. ETHERIDGE, M. A. 8z WILKIE, J. C. 1979. Grain size reduction, grain boundary sliding and the flow strength of mylonites. Tectonophysics, 58, 159-178. EVANS, B. & GOEZTE, C. 1979. The temperature variation of hardness of olivine and its implication for polycrystalline yield stress. Journal of Geophysical Research, 84, 5505-5524. FITZ GERALD,J. D. t~ CHOPRA, P. N. 1984. Distribution of strain in some geological materials experimentally deformed at high temperature and pressure. In: TRESSLER, R. E. & BRADT, R. C. (eds) Deformation of Ceramic Materials, 2, 321-328. FROST, H. J. & ASHBY, M. F. 1982. Deformation Mechanism Maps. Pergamon Press, Oxford. GIFKINS, R. C. 1976. Grain boundary sliding and its accommodation during creep and superplasticity. Metallurgical Transactions Series A, 7, 12251232. GOETZE, C. 1978. The mechanisms of creep in olivine. Philosophical Transactions of the Royal Society London, Series A, 288, 99-119. HANDY, M. R. 1989. Deformation regimes and the theological evolution of fault zones in the lithosphere: the effects of pressure, temperature, grainsize and time. Tectonophysics, 163, 119-152.

DYNAMIC RECRYSTALLIZATION AND STRAIN SOFFENING IN OLIVINE AGGREGATES HIRTH, G. 2002. Laboratory constraints on the theology of the upper mantle. In: KARATO, S.-I. & WENK, H.-R. (eds) Plastic Deformation of

Minerals and Rocks. Reviews in Mineralogy & Geochemistry, 51, 94-120. HIRTH, G. & TULLIS, J. 1992. Dislocation creep regimes in quartz aggregates. Journal of Structural Geology, 14 (2), 145-159. HmTH, G. & KOHLSTEDT, D. L. 1995a. Experimental constraints on the dynamics of the partially molten upper mantle one: Deformation in the diffusion creep regime. Journal of Geophysical Research, 100, 1981-2001. HmTH, G. & KOHLSTEDT, D. L. 1995b. Experimental constraints on the dynamics of the partially molten upper mantle 2. Deformation in the dislocation creep regime. Journal of Geophysical Research, 100, 15441-15449. HIRTH, G. & KOHLSTEDT, D. L. 1996. Water in the oceanic upper mantle: implications for rheology, melt extraction and the evolution of the lithosphere. Earth and Planetary Science Letters, 144, 93-108. HIRTH, G. & KOHLSTEDT,D. L. 2003. Rheology of the upper mantle and the mantle wedge: A view from the experimentalists. In: EiLER, J. (ed) Inside the Subduction Factory. Geophysical Monographs, 138, 83-105. HITCHINGS, R. S., PATERSON, M. S. & BITMEAD, J. 1989. Effects of iron and magnetite addtions in olivine-pyroxene rheology. Physics of the Earth and Planetary Interiors, 55, 277-291. HOLYOKE, C. W. & TULLIS, J. 2003. Low temperature bulging recrystallization in olivine aggregates: a mechanism for strain weakening and localization.

EOS Transactions of the American Geophysical Union, 84 (46) (Fall meeting supplement), Abstract $22A-0426. JIN, J., KARATO, S.-I. & OBATA, M. 1998. Mechanisms of shear localization in the continental lithosphere: inference from the deformation microstructures of peridotites from the Ivrea zone, northwestern Italy. Journal of Structural Geology, 20 (2-3), 195 -209. JUNG, H. & KARATO, S.-I. 2001a. Water induced fabric transitions in olivine. Science, 293, 1460-1463. JUNG, H. & KARATO, S.-I. 200lb. Effects of water on dynamically recrystallized grain-size of olivine. Journal of Structural Geology, 23, 1337-1344. KARATO, S.-I. 1989. Grain growth kinetics in olivine aggregates. Tectonophysics, 168, 255-273. KARATO, S.-I. 2003. The Dynamic Structure of the Deep Earth. Princeton University Press, Princeton. KARATO, S.-I. & WU, P. 1993. Rheology of the upper mantle: a synthesis. Science, 260, 771-778. KARATO, S.-I. & JUNG, H. 2003. Effects of pressure on high temperature dislocation creep in olivine. Philosophical Magazine, 83, 401-414. KARATO, S., TORIUMI, M. & FUJII, T. 1980. Dynamic recrystallization of olivine single crystals during high-temperature creep. Geophysical Research Letters, 7, 649-652. KARATO, S.-I., TORIUMI, M. & FUJI, T. 1982. Dynamic recrystallization and high temperature rheotogy of

157

olivine. In: AKIMOTO, S. & MANGHNANI, M. H. (eds) High Pressure Research in Geophysics. Advances in Earth and Planetary Science, 12, 171-189. KARATO, S.-[., PATERSON, M. S. & Fitz GERALD, J. D. 1986. Rheology of synthetic olivine grain aggregates: influence of grain size and water. Journal of Geophysical Research, 91, 8151 -8176. KOHLSTEDT, D. L., EVANS, B. & MACKWELL, S. J. 1995. Strength of the lithosphere: constraints imposed by laboratory experiments. Journal of Geophysical Research, 100, 17587- ! 7602. LEE, K.-H., JIANG, Z. & KARATO, S.-I. 2002. A scanning electron microscope study of the effects of dynamic recrystallization on lattice preferred orientation in olivine. Tectonophysics, 351, 331-341. MACKWELL, S. J., KOHLSTEDT, D. L. & PATERSON, M. S. 1985. The role of water in the deformation of olivine single crystals. Journal of Geophysical Research, 90, 11319-11333. MCDONNELL, R. D., PEACH, C. J. & SPIERS, C. J. 1999. Flow behavior of fine-grained synthetic dunite in the presence of 0.5wt% H20. Journal of Geophysical Research, 104, 17823-17845. MCDONNELL, R. D., PEACH, C. J., VAN ROERMUND, H. L. M. & SHERS, C. J. 2000. The effect of varying enstatite content on the deformation behavior of fine-grained synthetic peridotite under wet conditions. Journal of" Geophysical Research, 105, 13535-13553. MEI, S. & KOHLSTEDT,D. L. 2000a. Influence of water on plastic deformation of olivine aggregates 1. Diffusion creep regime. Journal of Geophysical Research, 105, 21457-21469. MEI, S. & KOHLSTEDT,D. L. 2000b. Influence of water on plastic deformation of olivine aggregates 2. Dislocation creep regime. Journal of Geophysical Research, 105, 21471-2148 i. ME1, S., BA1, W., HIRAGA, T. & KOHLSTEDT, D. L. 2002. Influence of melt on the creep behaviour of olivine-basalt aggregates under hydrous conditions. Earth and Planetary Science Letters, 201, 491-507. MONTI~SI, L. G. J. & HIRTH, G. 2003. Grain size evolution and the rheology of ductile shear zones: from laboratory experiments to postseismic creep. Earth and Planetary Science Letters, 211, 97-110. PATERSON, M. S. 1987. Problems in the extrapolation of laboratory rheological data. Tectonophysics, 133, 33-43. PATERSON', M. S. 2001. A granular flow theory for the deformation of partially molten rock. Tectonophysics, 335, 51-61. [hCKERING, F. B. 1976. The Basis of Quantitative Metallography. Institute of Metallurigcal Technicians, London. PIER1, M., BURL1NI, L., KUNZE, K., STRETTON, I. & OLGAARD, D. L. 2001. Rheological and microstructural evolution of Carrara marble with high shear strain: results from high temperature torsion tests. Journal of Structural Geology, 23, 1393-1413.

158

M.R. DRURY

REGENAUER-LIEB,K. & KOHL, T. 2003. Water solubility and diffusivity in olivine: its role in planetary tectonics. Mineralogical Magazine (Journal of Mineral Sciences), 67, 697-715. RENNER, J. & EVANS, B. 2002. Do calcite rocks obey the power law equation? In: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. & PENNOCK, G. M. (eds) Deformation Mechanisms,

Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 293-307. Ross, J. V., Ave LALLEMANT, H. G. & CARVER, N. 1980. Stress dependence of recrystallized-grain and subgrain size in olivine. Tectonophysics, 70, 39-61. RUTTER, E. & BRODIE, K. 1988. The role of grainsize reduction in the rheological stratification of the lithosphere. Geologische Rundschau, 77, 295-308. SniMuzu, I. 1998. Stress and temperature dependence of recrystallized grain size: A subgrain misorientation model. Geophysical Research Letters, 25, 4237 -4240. TULLIS, J. & YUND, R. A. 1985. Dynamic recrystallization of felspar: a mechanism for ductile shear zone formation. Geology, 13, 238-241. TwIss, R. J. 1976. Structural superplastic creep and linear viscosity in the Earth's mantle. Earth and Planetary Science Letters, 33, 86-100. TwIss, R. J. 1977. Theory and applicability of a recrystallized grain size paleopiezometer. Pure and Applied Geophysics, 115, 227-244.

TwIss, R. J. & SELLARS, C. M. 1978. Limits of applicability of the recrystallized grain size geopiezometer. Geophysical Research Letters, 5 (5), 337-340. VAN DER WAL, D. 1993. Deformation processes in mantle peridotites. Geologica Ultraiectina, 102. PhD thesis, Utrecht University. VAN DER WAL, D., CHOPRA, P. S., DRURY, M. R. & Fitz GERALD, J. D. 1993. Recrystallised grainsize stress relationships in experimentally deformed olivine-rocks. Geophysical Research Letters, 20, 1479-1482. WHITE, S. H. 1976. The effects of strain on the microstructure, fabrics and deformation mechanisms in quartzite. Philosophical Transactions of the Royal Society of London, Series A, 283, 69-86. WHITE, S. H. 1979. Large strain deformation: a report on a Tectonic Studies Group discussion meeting held at Imperial College, London, on 14 November 1979. Journal of Structural Geology, 1, 333-339. WHITE, S. H., DRURY, M. R., ION, S. E. and HUMPHREYS, F. J. 1985. Large strain deformation studies using polycrystalline magnesium as a rock analogue: Part I: grainsize palaeopiezometry in mylonite zones. Physics of the Earth and Planetary Interiors, 40, 201-207. ZHANG, S. • KARATO, S.-I. 1995. Lattice preferred orientation of olivine deformed in simple shear. Nature, 375, 774-777. ZHANG, S., KARATO, S.-I., FITZ GERALD, J. D., FAUL, U. U. & ZHOU, Y. 2000. Simple shear deformation of olivine aggregates. Tectonophysics, 316, 133-152.

Continental subduction and exhumation: an example from the Ulten Unit, Tonale Nappe, Eastern Austroalpine G I O R G I O R A N A L L I l, S I L V A N A M A R T I N 2 & R E Z E N E M A H A T S E N T E t

~Department of Earth Sciences and Ottawa-Carleton Geoscience Centre, Ottawa K1S 5B6, Canada (e-mail: granalli @ ccs. carleton, ca) 2Dipartimento Scienze Chimiche Fisiche & Matematiche, Universitgt dell'Insubria, 1-22100 Como, Italy Abstract: Some exhumed complexes in collisional belts consist of continental basement containing slivers of mafic and ultramafic material showing evidence of UHP metamorphism (P c. 3 GPa). Their PTt history can be interpreted in terms of subduction of continental material to depths > 100 km and subsequent exhumation. This type of tectonic history is illustrated by the Late Palaeozoic evolution of the Ulten Unit, Tonale Nappe, Eastern Austroalpine. The upper crustal felsic component (c. 80% by volume) incorporated mafic material at the trench, and peridotitic material at deeper levels in the subduction zone. The peridotites show evidence of a P-increasing, T-decreasing path before incorporation in the felsic material, compatible with flow in the mantle wedge above the subducting slab. After emplacement of the peridotites, which occurred at or near peak metamorphic conditions (P > 2.7 GPa, T > 850 ~ the complex underwent a twostage pre-Alpine exhumation path: a first, fast stage (c. 0.1-1 cm a-l), lasting c. 30 Ma and bringing rocks from depths > 100 km to approximately 25 km; and a second, slow stage (c. 0.01-0.1 cm a-l), lasting c. 100 Ma and bringing rocks to depths <20 km. The subduction of felsic material to the required depths can be modelled by analysing the time-evolution of negative buoyancy, which confirms that relatively light continental upper crust can be subducted to depths >200 km if attached to a mature oceanic slab that does not break-off during the early stages of continental subduction. The first exhumation stage can be accounted for by buoyancy-driven tectonic extrusion of continental slices along the subduction channel during continuing subduction. A force balance analysis shows that such a mechanism is compatible with the rheology of felsic and intermediate rocks at high temperature. The second exhumation stage is compatible with isostatic rebound and tectonic denudation following slab break-off. The conclusion that fast exhumation occurs during continuing subduction and before slab break-off is in accordance with the observed rates, which show fast movement of the rising slices with respect to the surrounding material. Slab break-off, on the other hand, generates a long-wavelength gentle upwarping of the overlying region, which is more compatible with later and slower exhumation rates.

In this paper we present a model of subduction and exhumation of continental crust, containing slivers of mafic and ultramafic material, which may be applicable to the Late Palaeozoic history of the Ulten Unit, Tonale Nappe, Eastern Austroalpine, and similar high-grade exhumed basement complexes. The regional tectonic setting of the Ulten Unit is dominated by the Cretaceous to Tertiary oblique convergence of the African (Adria) and Eurasian plates (cf. for example, Martin et al. 1998; Dal Piaz 1999). The Ulten Unit has been interpreted as a fragment of the Palaeozoic 'Variscan' belt,

only lightly overprinted by the Alpine orogenesis (Godard et al. 1996). This is in agreement with larger-scale interpretations of the Eastern Austroalpine system (Dal Piaz 1993; Neubauer & v o n Raumer 1993; von Raumer & Neubauer 1993). The regional pre-Alpine tectonics is generally thought to have involved three stages: oceanic closure before the Devonian (t > 430 Ma); continental collision during the Devonian and Early Carboniferous ('Variscan orogeny', 400 < t < 330Ma); and Permian to Jurassic lithospheric attenuation and tiffing (opening of the Tethys, 300 < t < 150 Ma).

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and

Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 159-174. 0305-8719/05/515.00

9 The Geological Society of London 2005.

160

G. RANALLI E T AL.

We first review the evidence relevant to the P T t evolution of the Ulten Unit from oceanic

closure in pre-Variscan times until lithospheric attenuation in the Permo-Triassic. Then we show that this evolution can be explained by a model involving subduction of felsic continental basement, emplacement of mantle wedge peridotites in the subducted continental basement, and a two-stage exhumation history, in which a relatively short, fast stage, governed by buoyancydriven tectonic extrusion of felsic slices during continuing subduction, is followed by a later, slow stage, possibly related to slab break-off and consequent upwarping and erosion. Several models have been proposed for the exhumation of high-pressure (HP) and ultrahigh-pressure (UHP) metamorphic rocks (cf. the review by Ring et al. 1999). The fast exhumation stage of UHP (>2.5 GPa) units is usually envisaged in terms of tectonic extrusion of relatively continuous thin slices (Chemenda et al. 1995, 1996; Ernst et al. 1997; Ernst 2001) or upward viscous flow along the subduction channel (Shreve & Cloos 1986; Cloos & Shreve 1988; Mancktelow 1995). The latter has been modelled quantitatively by Burov et al. (2001) and Gerya et al. (2002). The two processes are not mutually exclusive, and both are compatible with the geological evidence of synsubduction, syncollisional exhumation of UHP rocks.

Geological framework Petrology

The high-grade basement of the Ulten Unit is exposed in the Nonsberg-Ultental area in the NE part of the Tonale Nappe, between the Giudicarie, Tonale, and Pejo faults (Figs 1 & 2). It consists of strongly foliated garnet § kyanite _+ sillimanite banded gneisses and leucocratic migmatites, containing small (decimetric to metric) mafic lenses of retrogressed eclogites and larger (metric to decametric) lenses of garnet- and spinel-peridotites (Fig. 2). This rock complex (Ulten tectonic mrlange) consists of c. 70-90% by volume of felsic rocks, with subordinate (c. 10-15%) ultramafic bodies and rare (c. 5%) mafic lenses. We use the term 'mrlange' in the same sense as Miller et al. (1980) for the Tauern window, and Godard (1983) for the Vendre eclogites, to denote a complex containing tectonically incorporated fragments and slices, in the present case during subduction (hence 'subduction mrlange'). It shows well-preserved pre-Alpine parageneses and fabrics recording high P and T conditions. Detailed discussions of the petrology can be found in work by Godard et al. (1996), Martin et al. (1998), and Del Moro et al. (1999).

Some characteristic features of the mrlange rocks are shown in Figure 3. The migmatitic gneisses exhibit mm-to-cm spaced bands consisting of alternating biotite + garnet § kyanite + rutile-rich melanocratic components and quartz § plagioclase + K-feldspar-rich leucocratic components (Fig. 3a). They are eclogitic schists, intensely overprinted by high-grade amphibolite facies assemblages up to migmatization, with local K-feldspar-rich granitic portions. Ductile deformation and recrystallization under high-P conditions has resulted in a strongly foliated mylonitic fabrics (Fig. 3b). Occasionally, leucocratic pockets cross-cutting the foliation occur (Godard et al. 1996; Martin et al. 1998). The mafic lenses are mainly amphibolites containing relics ofeclogite parageneses (Fig. 3c). In zones of intense migmatization, they recrystallized to coarse-grained amphibolites, consisting of amphibole + garnet + plagioclase § clinopyroxene and minor phases. Where migmatization was less intense, relics of an eclogite paragenesis (clinopyroxene § garnet § quartz + plagioclase § rutile + amphibole + fluorapatite § graphite) can be observed. The surrounding migmatites contain large amphibole crystals and derive from a metasomatic exchange between mafic rocks and migmatitic leucosome (Martin et al. 1998). The peridotitic bodies are usually located at the transition between underlying strongly foliated and weakly migmatized gneisses and overlying coarse-grained migmatites composed of predominant quartz and plagioclase (trondhjemites). They range from weakly deformed, coarsegrained porphyroclastic spinel-lherzolites to highly deformed, fine-grained garnet-lherzolites, garnet-dunites, amphibole-spinel-dunites, and partly serpentinized rocks (Fig. 3d, e, f). Their fabric is usually concordant with the foliation of the surrounding gneisses. Spinel-bearing garnetfree lherzolites with protogranular coarse texture are regarded as the protolith of the fine-grained garnet-beating peridotites. Garnet grew at the expense of spinel, which is sometimes still preserved as relics within the garnet, during prograde metamorphism (Godard et al. 1996). The retrograde decompressional evolution of peridotites is documented by development of symplectitic orthopyroxene + clinopyroxene _+ amphibole +_ spinel coronas around garnet, with occasional complete replacement of garnet by amphibole § spinel aggregates (Godard & Martin 2000). PTt-path

The P T t - p a t h of the Ulten Unit is shown in Figure 4. Here we summarize only the mainly Palaeozoic, pre-Alpine history (t > 200Ma)

CONTINENTAL SUBDUCTION AND EXHUMATION

161

Fig. 1. Metamorphic map of the Central and Eastern Alps. The area shown in Figure 2 is identified by the rectangle.

Fig. 2. General map (lower fight-hand comer) and detailed map (location shown by the rectangle in the general map) of the Tonale Nappe and related units. (1) Ortler Nappe: gneisses and micaschists. (2) Tonale Nappe: (a) sillimanite paragneisses; (b) garnet-kyanite paragneisses; (c) leucocratic migmatites; (d) ultramafic bodies. (3) Southern Alps: (a) Permian volcanics; (b) Permo-Jurassic sedimentary cover; (c) Cretaceous flysch and Tertiary magmatic rocks.

162

G. RANALLI E T AL. a

c

d

0

f

Fig. 3. Rocks of the Ulten m61ange. (a) Biotite (bt) + garnet (gt) + kyanite (ky) metanocratic level in banded gneiss; (b) inylonitic foliation in banded gneiss with quartz ribbons and alkali feldspar (feld) porphyroclast; (c) amphibolite with garnet (gt), clinozoisite (cz), and rutile (ru); (d) peridotite with orthopyroxene (opx) and clinopyroxene (cpx) porphyroclasts exsolving garnet (gt); (e) prograde Cr-rich spinel (cr-spl) and olivine (ol) included in a garnet porphyroblast (gt) in recrystallized dunite; (t') amphibole porphyroblast (amph) derived from after-garnet kelyphites and including relict spinel (spl). that is the object of our study (for detailed discussions of the evidence, cf. Obata & Morten 1987; Nimis & Morten 2000; Tumiati 2002; Tumiati et al. 2003). The protogranular coarse-grained spinel peridotites equilibrated at P c. 1.5 GPa, T c. 1250 ~ (Nimis & Morten 2000), at an uncertain but in any case old age (t > 500 Ma; Ntaflos et al. 1993; Th6ni 1999). These PT-conditions suggest a high geothermal gradient, probably reflecting conditions in the asthenosphere just below a thin, possibly back-arc lithosphere (z c. 50 km), in the mantle wedge above a subducting slab (Nimis & Morten 2000). After equilibration, the peridotites followed a P-increasing, T-decreasing path, to reach peak conditions in the garnet stability field (P > 2.7 GPa, T > 850 ~ at t c. 330 Ma (Tumiati et aL 2003). Equilibrium conditions for the peridotites were estimated by a combination of

F e - M g olivine-garnet and garnet-orthopyroxene exchange thermometers with the Al-orthopyroxene barometer (O'Neill & Wood 1979; Harley 1984; Brey & K6hler 1990). At or near the peak conditions, peridotite fragments and lenses were emplaced in a subducted continental felsic basement, probably of Caledonian age (t >__400 Ma), already including mafic fragments, and undergoing partial migmatization. The process of emplacement cannot be defined on the basis of the present evidence, but in all likelihood was related to a gravitational sinking mechanism (Brueckner 1998), assisted by shear concentration near the top of the subducting slab. The lack of UHP relics in the felsic component is attributable to retrogression, as all components of the m~lange yield garnetwhole rock and garnet-clinopyroxene S m - N d ages indicating an isotopic homogenization

CONTINENTAL SUBDUCTION AND EXHUMATION

163

P (GPa) t = 330 Ma P_> 2.7 G P a T>_ 8 5 0 " C

3--

[]

I

~ , ~

""

unknown age P ~ 1.5 G P a T ~ 1250~

=

P ~ 0.7 G P a []

~:. . . .

lg

[]

.....

t < 100 Ma

I 200

t ~ 205 Ma T ~ 5 0 0 " C

p ~ 0 ~ Gpa

T ('c)

T~ 350*C

I 4oo

I 600

I 800

I 1ooo

I 12oo

14 o )

Fig. 4. PTt-path of the Ulten m~lange. Dash-dot line (1) denotes the path of the mantle peridotites before their incorporation in the subducted continental basement. The full line denotes the pre-Alpine exhumation path of the mtlange, consisting of a fast stage (from 2 to 3), and a slow stage (from 3 to 4).

event at t c. 330 Ma (Tumiati et al. 2003). This constrains the peak metamorphism of the peridotites and the partial melting of the felsic component to the same age near the end of the Variscan orogenesis. The solidus temperature of wet felsic material at P c. 2.5-3.0 GPa is Tm c. 800-900 ~ (cf. Bousquet et al. 1997), while wet ultramafic material is below the solidus at the same pressures (Tin > 1000 ~ cf. Wyllie 1988). The common age also implies that the mtlange shared a common history since the incorporation of the peridotite fragments in the subducted continental basement. The first stage of exhumation took the Ulten Unit from peak conditions to midcrustal depths in about 30 Ma (P = 0.7 GPa, T = 500 ~ - closure temperature of white mica - at t = 300 Ma), resulting in an average vertical exhumation rate of c. 0.25 cm a-1 and an average cooling rate of c. 12 ~ Ma -1. Assuming that exhumation took place along the subduction channel with an average dip angle of 30 ~ this yields c. 0.5 cm a -1 for the average rate along the direction of motion, and c. 3 ~ km-~ for the corresponding temperature gradient, well within the ranges estimated for gradients near the top of subducting slabs (Peacock 1996; Ranalli et al. 2000). However, these are only time-averaged rates; there are no controls over their time-dependence during this exhumation stage. After 300 Ma, both exhumation and cooling rate slowed down by more than one order of magnitude. The average vertical exhumation

rate in the Permo-Triassic (300-205 Ma) was <0.01 c m a -1. Subduction by then had likely ended, possibly as a consequence of slab breakoff (Davies & von Blanckenburg 1995). The PTt history of the Ulten Unit during this stage is probably the combined result of upwarping and erosion due to isostatic adjustment after break-off and lithospheric attenuation and rifting related to the opening of the Tethys. During the Alpine orogenesis, according to most reconstructions the Ulten Unit was part of the overriding plate (Adria), and consequently was no longer involved in subduction (cf. for example, Dal Piaz 1993).

G e o d y n a m i c scenario

On the basis of the evidence discussed above, the following scenario (schematically depicted in Fig. 5) seems to accurately describe the preAlpine tectonic evolution of the Ulten m~lange (cf. also Nimis & Morten 2000; Tumiati 2002; Tumiati et al. 2003; Martin 2003; Ranalli 2003): 9 Oceanic closure had already occurred before incorporation of the peridotites in the continental basement, and subduction of continental material had lasted a sufficiently long time for the continental basement to have reached minimum depths of c. 100 km. 9 Spinel-peridotites, equilibrated in the mantle wedge above the subducting slab, were dragged by convection close to the upper

G. RANALLI ETAL.

164

A

of subduction, and subsequent lithosphere attenuation and opening of the Tethys in Permian-Jurassic times.

time range : 500 - 330 Ma

km 0 9

50

. . . . . . . . . . . . . . . . . . . .

//

I ii

....

-

....................

Geodynamic model

7 5 -~

loo

~.Z~.
"xl .

/-

time range : 330 - 300 M a kmO~

/...:?;~,a) ............... 50 75 100

We concentrate on three first-order aspects of the above scenario, to show that it is mechanically and rheologically plausible. We demonstrate the feasibility of subduction of continental crust to depths of the order of hundreds of km; we show that the positive buoyancy of felsic material generates forces sufficient to overcome the strength of rocks at the relevant P,Tconditions; and we argue that, at least in the case under study, slab break-off, while a distinct possibility, has probably occurred rather late in the process.

/ /

Subduction of continental lithosphere

C km 0 -

50

time range : 300 - 205 Ma .................................... )~---~7 / /" ..............................

75

.~>// -

~>~ ....... -

/

100

/ Fig. 5. Geodynamic scenario: mantle wedge peridotites (1) are emplaced in subducted continental basement (2), which is rapidly exhumed by buoyancy-driven tectonic extrusion while subduction continues (3), and then more slowly, possibly as a consequence of slab break-off (4) and possible partial melting (crosses).

surface of the slab along a P-increasing, T-decreasing path, and were emplaced in the subducted continental crust near the top of the slab, possibly by a gravitational sinking mechanism, at or near peak conditions at t c. 330 Ma; peak conditions in the m61ange were achieved at this time or shortly after incorporation of the peridotites; partial migmatization of the felsic component of the m61ange occurred at about the same time. Exhumation was fast in the first 30 Ma (motion rate c. 0.5 cm a-1 along the subduction channel) and occurred by a buoyancydriven tectonic extrusion mechanism during continuing subduction; it then became one order of magnitude slower, probably as a consequence of slab break-off and cessation

The Ulten peridotites were equilibrated in the upper mantle, at a depth z c. 50 km and temperature T c. 1250 ~ (Nimis & Morten 2000). These conditions show that the thermal gradient was high and the overlying lithosphere thin, a tectonic environment probably corresponding to a mantle wedge above a subducting oceanic slab. Their initial PT-path is compatible with flow lines and temperature distribution in a convecting mantle wedge (Honda 1985; van den Beukel 1992; Peacock 1996; Nimis & Morten 2000). Models of slab-induced flow in the mantle wedge show that flow lines are subhorizontal and directed towards the slab in the uppermost part of the wedge (e.g. Honda 1985; Giunchi et al. 1996). Isotherms are subhorizontal sufficiently far from the slab, before bending downwards in the proximity of the slab. Therefore, during their initial (horizontal) approach to the slab, the peridotites followed an isobaric, nearly isothermal path, followed by a compressional, cooling path as they began their descent toward the upper surface of the slab where they were eventually emplaced. If oceanic subduction was already active at the time of equilibration of the peridotites, the slab tip must have reached the mantle transition zone by the time the peridotites were incorporated in the slab, even for very slow subduction rates ( > 0 . 5 - 1 c m a 1). The emplacement of mantle wedge peridotites into continental basement implies that the continental material had reached depths _>100 kin. Even with a conservative estimate of the average subduction dip angle, this requires the subduction of continental material for a downdip length > 150-200 km.

CONTINENTAL SUBDUCTION AND EXHUMATION Continental crust is upwardly buoyant. Even neglecting the sluggish kinetics of phase transitions, detailed density calculations show that continental material is less dense than mantle material up to and including UHP facies for felsic and intermediate composition (Bousquet et al. 1997). This implies that the subduction of continental crust gradually reduces the overall downward buoyancy of the subducting slab. Nevertheless, continental material can be subducted to depths of 100-300 kin, as shown by geological evidence (cf. for example, Chopin 1984, 2003; Ernst et al. 1995; Ernst 2001), analogue models (Chemenda et al. 1995, 1996; Regard et al. 2003), and numerical models (Ranalli et al. 2000). We estimate the maximum amount of subductable continental material by assuming that subduction ceases when the overall buoyancy of the slab becomes zero. The time-dependent evolution of temperature within the slab since the beginning of subduction is calculated by solving the coupled mass, momentum and energy conservation equations in a 1800 km thick and 2900 km wide 'box' containing the slab and surrounding mantle (Fig. 6; Mahatsente & Ranalli 2004 for details). An imposed horizontal velocity drives the subducting plate towards the trench, and the surrounding mantle material reacts by viscous flow to the sinking of the slab. To minimize the effects of the vertical boundaries of the box on the thermal structure of the slab, they are placed approximately 1450 km away on either side of the trench. The temperature distribution in the slab and surrounding mantle is the result of combined diffusion and advection. The thermal analysis is posed as a transient problem with time-dependent . . . . . . . . V=O . . . . . . . . . . . . . . .

Continenlal

lithnsnh~re.

v=..,,t,w~

~ d ' ~ li:l#loJ~ P h ~ ' ~

_

.

~

v=0 l e e o o

u m e o o o e o u

9 ooo

9 o o o u o u

o o e o o o o o e o ~ o e o ~

q

2929 km

Fig, 6, Model box used for the computation of temperature distribution and negative buoyancy. Black circles represent no-slip boundary conditions. The overriding plate does not move; the subducting plate has an imposed time-dependent velocity V(t). The subducting plate is initially oceanic; continental material is assumed to arrive at the trench when the tip of the slab has reached a prescribed depth.

165

temperature and temperature and pressuredependent thermomechanical parameters. Table 1 shows the densities, their P,T-derivatives, thermal conductivities, specific heats, and the initial elastic parameters in the slab and surrounding mantle. Viscous deformation is modelled with the theology of olivine, garnet, and perovskite in the upper mantle, transition zone, and lower mantle, respectively (for parameters and further discussion, see Ranalli et al. 2000; Mahatsente & Ranalli 2004). The equations are solved with the explicit finite element code ABAQUS 6.3 (Hibbitt et al. 2001). The negative buoyancy as a function of time is estimated from the thermal structure and composition of the slab, by integrating the density anomalies within the slab. A full account of the procedure can be found in the work of Mahatsente & Ranalli (2004). Figures 7 and 8 show the evolution of negative buoyancy since the initiation of subduction. The subducting oceanic lithosphere (dip angle 45 ~ is assumed to be thermally old, and to consist of a 7 km thick crust overlying a 93 km thick lithospheric mantle; the continental lithosphere consists of a 35 km thick crust overlying a 65 km thick mantle lid. The width of the passive subducting margin (the distance over which the crust changes from oceanic to continental thickness) is taken as 150 km. After the arrival of continental crust, the positive (upward) buoyancy of felsic or intermediate material (assuming it does not enter the eclogite stability field) decreases the downward negative buoyancy of the slab. The time span required to bring the total buoyancy to zero depends on lithospheric age, convergence rate, and dip angle. In the case of a constant subduction velocity (Fig. 7), the subduction of continental crust, assumed to begin when the tip of the slab has reached the bottom of the mantle transition zone, reduces the buoyancy to zero in c. 15-20 Ma for convergence rates of the order of a few cm a -1. If the convergence rate is assumed to decrease exponentially after the arrival of continental material at the trench and then the slab is allowed thermally to equilibrate with the surrounding mantle (Fig. 8), the time span over which the buoyancy is reduced to zero does not change by more than a few Ma. In both cases, if subduction continues until the overall buoyancy is reduced to zero, continental material can be subducted to depths of 200300 kin, as also concluded by Ranalli et al. (2000) on the basis of a similar model. This is a lower bound estimate if the subducted crust is of mafic composition. Taking the vertical axis positive downwards and assuming that subduction continues until

166

G. RANALLI ETAL.

Table 1. Thermoelastic parameters used in the model (p, density; K, thermal conductivity; C, specific heat; E, Young's modulus; v, Poisson's ratio) Po • 103 Ko Co x 103 Eo x 1011 (kgm -3) (Win - I K -1) (Jkg - I K -1) (Pa) Oceanic crust Continental crust Mantle: (z < 410 kin) (410 < z < 660) (z > 660 kin)

Vo

Op/OT• 1012 Op/OP • 10 -8 (kgm - 3 K - ) (kgm-3pa -l)

2.92 2.9

2.7 2.5

0.9 0.9

1.724 1.724

0.278 0.278

- 5.5 - 6.0

3.8 4.0

3.36 3.58 4.12

5.9

0.8 0.8 0.8

2.344 2.715 4.173

0.294 0.295 0.298

- 8.0 -6.8 -11.5

2.4 2 1.6

The subscript zero refers to initial conditions. Densities of each phase vary with temperature and pressure as shown in the last two columns. (FromMahatsente& Ranalli 2004.) the total buoyancy of the slab vanishes (i.e. excluding earlier slab break-off and convective instabilities), the maximum downdip length (along the slab) of subducted continental material is given approximately by (Ranalli et al. 2000; Ranalli 2003) Lc -- -Fs/(ApcgHc)

(1)

where Lc and He are downdip length and thickness of subducted continental crust below the depth of the Moho (there is no positive buoyancy until subducted crust has reached sub-Moho depths), Ape is the average density difference between subducted continental crust and surrounding mantle, g is acceleration of gravity, and Fs = Apmg Hs Ls is the downward buoyancy generated by the slab (both oceanic and continental) of downdip length Ls below the depth of the lithosphere/asthenosphere boundary, thickness Hs, and average density excess Apm with respect to the surrounding mantle. Downdip length of subduction of continental crust (for a dip angle 6 = 3 0 ~) is given in Figure 9 as a function of overall downdip length of the slab for Apm = 50 kg m -3 (which is an average value reflecting the positiondependent density difference estimated by the numerical model), Hs/Hc = 3.0, and two endmember hypotheses as to the density contrast (at depths of 80-100kin) between the felsic continental material and the surrounding mantle: light (metastable) continental crust (Ape = - 4 0 0 k g m -3) and equilibrated continental crust (Ape -- - 2 5 0 kg m - 3; for estimates of densities and P,T-derivatives see Bousquet et al. 1997; Ranalli et aL 2000). A change in subduction angle does not significantly affect the results. These are only rough estimates, using average values for the quantities involved and neglecting the time-dependence of subduction velocity. Nevertheless, unless the subducted continental crust is unusually thick (an unlikely

situation at a passive margin), the depths required by the peak conditions in the Ulten m61ange are well within the model predictions. The mechanism of emplacement of the peridotites in the subducted continental basement was most likely a sinking intrusion as proposed by Brueckner (1998). At P c. 2 - 3 GPa, T c. 7 0 0 900 ~ densities of felsic material are in the 2 9 0 0 - 3 1 0 0 k g m -3 range, while mantle densities are about 3 3 0 0 - 3 4 0 0 k g m -3 (Bousquet et al. 1997; Ranalli et al. 2000). Peridotite lenses can therefore be emplaced in the underlying continental material by gravitational sinking. The process is probably helped by shear-induced mixing. Both peridotites and felsic rocks deform in a ductile fashion at the relevant temperatures (800-900 ~ but the effective viscosity of peridotites is 1 - 3 orders of magnitude larger than that of felsic materials (Ranalli 1995, 2000). This difference accounts for the peridotite lenses keeping their individuality during the entrainment and the subsequent deformation. Fast exhumation: buoyancy-driven tectonic extrusion

In a period of about 30 Ma (330-300 Ma), the Ulten m~lange was exhumed at a vertical average rate c. 0.25 c m a -1. We envisage this process as occurring by buoyancy-driven detachment and extrusion of slices of subducted continental crust along the subduction channel during continuing subduction, as proposed by van den Beukel (1992), Ernst & Peacock (1996), Ernst et aI. (1997), and Ernst (2001). Analogue models of this mechanism have been provided by Chemenda et al. (1995, 1996). Exhumation can occur also by viscous channel flow (e.g. Mancktelow 1995; Burov et al. 2001; Gerya et aL 2002). However, in many cases tectonic slices are exhumed as units at different (discrete) times, and are separated by a normal-sense

CONTINENTAL SUBDUCTION AND EXHUMATION

167

6.5 ......................................................................................................... ' 6

(A)

5.5

E

5

Z

V= 5 cm a ~

-~o 4.5 4 3.5

~_

3

2.5

2 1,5

0.5 0 0

5

10

15

20

25

30

35

40

4.5 4 ] (B)

E

z

3.5

o

,

3-{

u

V = 3.5 cm a -1

i

o

2

,,

~ 1.51

~

i

/

"" 0 . 5 ,

~,

0 3.5

--vE z

5

10

15

20

25

.................................................................................................................................................

30

45

40

..................................................................................

(c)

3 2.5

~o -.~

35

:

2

~1.5 //

/

0.5

0

zJ

...., V = 2 . 5 cm a -1

/J

~,.

f/ i

-.

5

t0

15

20

25 30 35 Time (Ma)

40

45

50

"

55

60

Fig. 7. Time evolution of negative buoyancy since subduction initiation for three different constant subduction rates. Full lines: purely oceanic slab; dotted lines: continental material arrives at the trench 20 Ma after the beginning of subduction.

shear zone from the overlying material and by a reverse-sense shear zone from the underlying material (Ernst e t al. 1995, 1997 for a review). This seems to favour a mechanism of tectonic extrusion.

The continental crust is less dense than the surrounding mantle material below the Moho, and is therefore subject to positive buoyancy. This is true also of the Ulten m~lange, because of its relatively small content of mafic and ultramafic

G. RANALLI ETAL.

168

5

~4.5 E z

4

~o3.5 r" 3 ~2.5 O

u..

2

=~

=o 1 m 0.5 0 0

5

10

15

20

25

30

35

40

3.5 'E Z O

( B )

3

y-------.

V= 3.5 cm a -1

2.5 /4... . . . . . . .

v

..

\\

2

O

,,o 1,5

;'.

,/

,

\

",~ \

O

~,o.s 0

5

10

15

20

25

30

35

40

2.5 ";

;: ( C )

:

V = 2 -5cma-1

r~1.5 u.

1

~'0.5

0

5

10

15

20

25

30

35

40

Time ( Ma )

Fig, 8. Time evolution of negative buoyancy under time-varying subduction velocity (increasing in the first 5 Ma, then constant for 15 Ma at the peak values shown, then decreasing, followed by thermal equilibration). Full lines: purely oceanic slab; dotted lines: continental material arrives at the trench 20 Ma after the beginning of subduction. lenses. Exhumation occurs when the positive buoyancy overcomes the resistance to detachment and motion. Here we give an order-ofmagnitude estimate of the force balance, and compare it with the inferences that can be

drawn from the rheological properties materials at high P,T. The results are shown Figure 10. The updip (parallel to the slab) c o m p o n e n t the positive buoyancy of a crustal slice

of in of of

CONTINENTAL SUBDUCTION AND EXHUMATION

169

140

120

Apc = - 400 kg m - 3 &Pc = - 250 kg m-3

.-~

100

~ 80 ,x

.Y

~.~

60

~ I ~ i - ---~-~

40 20 0 90

110

130

150

170

190

210

Ls(km) Fig. 9. Downdiplength of continental crust subductable below the Moho (Lc) as a function of total downdip length of the slab below the lithosphere/asthenosphere boundary (Ls), for two different average density differences between subducted crust and surrounding mantle.

thickness D and downdip length L~c below the depth of the Moho is given to the first order by Fa: = 0,

z
F& = ApcgDL,~ sin 6 = ApcgD(z - He),

z >Hc

to detachment, which is given approximately by the integral of the critical shear stress along the top and bottom boundaries of the crustal slice. Assuming these to be roughly parallel to the subduction zone, this resistance is

(2) For upward motion (along dip) to occur, the positive buoyancy must overcome the resistance

Rc = ( 2 r / s i n 6)z,

z> 0

(3)

where ~-is the average shear strength along the perimeter of the slice (assumed to be twice the

1.2E+13

1.2E+13 -,Apc = - 400 kg m - 3 D = 30 km -L -*-- Apc=-250kgm -3'D 30kmjFSc

1E+13

.... z = 2 0 M P a ; 5 = 3 0 ....... I: =12

~ 't ~

1E+13

/

8E+12

8E+12

=.,

...-.

"7

E Z

E

6E+12

6E+12

4E+12

4E+12

2E+12

2E+12

O

0

20

~r

.

40

.

.

.

60

.

80

100

120

z

0 140

Z (km) Fig. 10. Upward buoyancy of continental material (left vertical axis, full lines) and resistance to detachment of continental slices (right vertical axis, dotted lines) as functions of maximum depth reached by the continental material, for different density contrasts and shear strengths.

170

G. RANALLI ET AL.

total downdip length) reaching a depth z from the surface. A necessary and sufficient condition for the continental upward buoyancy force to overcome the resistance to detachment is that the absolute value of F~c increase with depth more rapidly than that of Ro, namely ]dFe,c/dz[ > [dRc/dzl

(4)

which places an upper bound on the average shear strength along the perimeter of the crustal slice. For D = 3 0 k m and 6 = 3 0 ~ these bounds are approximately 30 and 20 MPa for low-density and high-density crust, respectively. Additionally, the condition that the crossover of buoyancy and resistive forces has to be reached at a depth of about 100 km further reduces these upper limits (between 20 and 13 MPa, respectively; see Fig. 10). These limits decrease by a factor equal to D/Hc if the detached slice is considerably thinner than the crust, as is usually the case (typically, 5 - 1 0 km thick; Ernst 2001). However, in the above argument it has been assumed that detachment occurs simultaneously along the whole downdip length of the slice, which is certainly an overestimate. A reduction by a factor f - - D / H c of the downdip length of the detached slice reduces the total frictional resistance by the same amount as the positive buoyancy. Consequently, the upper bound of 10-20 MPa on the shear stress is a reasonable estimate. It is reasonable to assume that the rheology of the exhumed material is controlled by the volumetrically predominant continental rocks. Both plagioclase- and quartz-rich rocks are soft at high temperature, with creep strengths < 1 0 M P a at T c. 850~ for strain rates _<10 -12 s -1 (Ranalli 1995, 2000). Direct extrapolation of experimental results, therefore, leads to estimates of strengths compatible with the limits imposed by the previous force balance argument. However, neither pressure nor metamorphic transformation effects on strength are known with any accuracy, although recent experimental results on the rheology of garnet and garnet-rich rocks suggest that even rocks in the eclogite facies may be relatively soft at high temperature (Wang & Ji 1999, 2000; Ji et al. 2003). Dynamically recrystallized grain sizes in the Ulten garnet peridotites are in the range 20-200 ~m (Tumiati 2002), corresponding to flow stresses of 1 0 - 1 0 0 M P a in wet olivine (Drury & Fitz Gerald 1998). Shear stresses at the interplate boundary along subduction zones are usually estimated to increase with depth to

a maximum of about 50 MPa at z c. 30-50 kin, then to decrease gradually with depth (van den Beukel & Wortel 1988; Hoogerduijn Strating & Vissers 1991; van den Beukel 1992; Peacock 1996). They can be reduced considerably by high pore fluid pressures (Hyndman & Wang 1993; Oleskevich et al. 1999). Therefore, the above upper bound estimates on shear zone strength necessary for detachment and exhumation are compatible with experimentally determined ductile strengths and tectonic stresses. Another force acting to drive exhumation could be provided by the pressure gradient acting on the crustal slice along the downdip direction if the slice is wedge-shaped (thickness decreasing with increasing depth), similarly to the mechanism of tectonic extrusion proposed by Mancktelow (1995) for sedimentary prisms, and by Thompson et al. (1997) for orogenic roots. Buoyancy-driven tectonic extrusion takes place during continuing subduction of the slab, and slab break-off is not a necessary condition for its occurrence. As mentioned previously, the exhumed slices tectonically overlie a reverse ductile shear zone, and are in turn overlain by a normal ductile shear zone. These features have been identified in Tertiary exhumed units in the Himalayas (e.g. Chemenda et al. 2000; Treolar et al. 2003). In the Tonale Nappe, the Ulten Unit is overlain by the older Cima Mezzana eclogitic slice (Martin 2003).

Slow exhumation: slab break-off and lithosphere extension The fast stage of exhumation of the Ulten m61ange was over by t c. 300 Ma. For the next 100 Ma, both exhumation and cooling rate were much slower. Although precise constraints for a quantitative model are lacking, we think that this stage could correspond to slab break-off in the sense of Davies & von Blanckenburg (1995). According to their estimates, an oceanic slab subducting with a velocity _>1 cm a-1 and dip angle > 20 ~ requires the subduction of continental material to depths > 100 km before the change in buoyancy becomes sufficient to overcome its strength. Break-off can occur at shallower depths (>35kin) if the continental lithosphere arriving at the trench is relatively warm (van de Zedde & Wortel 2001). However, the depth of continental subduction in the Ulten case shows that break-off occurred, if at all, in the oceanic part of the slab at z > 100 kin. Models of slab break-off show that the consequent uplift affects a wide region, with maximum

CONTINENTAL SUBDUCTION AND EXHUMATION surface uplift of the order of a few km (Buiter et al. 2002), and that uplift rates (comparable with exhumation rates if erosion keeps approximate pace with uplift) are low ( < l mm a -1 as an order of magnitude; Giunchi et al. 1996). If exhuming crustal slices had reached approximate gravitational equilibrium in the midcrust before break-off, the subsequent exhumation rate would therefore be one order of magnitude lower than before, as observed in the Ulten m61ange. The slow cooling rate during this stage may be attributed to the depth difference between the slices and the level of the break-off. For a material with thermal diffusivity K c. 10 -6 m 2 s-1, such as the lithosphere, the characteristic diffusion time of a thermal perturbation over a distance of several tens of km is of the order of 100 Ma. Using a simple one-dimensional model, Davies and von Blanckenburg (1995) estimate that the temperature increase over 50 Ma at depth z = 20 km would be of the order of 100 ~ if temperature is raised instantaneously to asthenospheric values at z = 80 km. Therefore, both uplift rate and cooling rate of the Ulten m61ange during slow exhumation are compatible with a relatively deep slab break-off. Evidence for the late timing of slab break-off may be related to the production of trondhjemitic melts, which have formed dykes and veins that cut the partially migmatized continental basement. The geochemistry of these trondhjemites excludes melting in situ, and exhibits mixed lower crustal and mantle affinity (Martin et al. 1998). Their age of formation is uncertain, but available results point to c. 300 Ma or soon afterwards (Del Moro et al. 1999), which could date slab break-off and consequent melting of rocks that came in contact with asthenospheric material. Removal of a cool lithospheric root, either by break-off or by convective instability, results in deviatoric tensional stresses in the overlying lithosphere (e.g. Fleitout & Froidevaux 1982). Consequently, slab break-off may have generated a sequence of events involving not only cessation of subduction and slowing down of exhumation, but also lithospheric extension and the widespread Permo-Triassic magmatism of the Eastern Alps.

Concluding remarks In our view, the Ulten m61ange formed by entrainment of mantle wedge peridotites in subducted felsic continental basement. The emplacement of the peridotites occurred at, or shortly before, the attainment of peak conditions. Continental material, therefore, had been subducted to

171

a depth z c. 100 km by 330 Ma. For subduction dip angles and velocities in the ranges 2 0 - 4 0 ~ and 1-5 cm a -~, respectively, this implies that continental subduction had gone on for 1030 Ma before formation of the mrlange. A kinematic model of oceanic and continental subduction proves that this is feasible, provided that the total downdip length of the slab (including the oceanic part) is approximately a factor of two or more larger (depending mainly on the density and thickness ratios) than the downdip length of subducted continental crust. After 330 Ma, the Ulten m~lange underwent a common pre-Alpine exhumation history. For the first 30 Ma, exhumation was fast (c. 0.5 cm a-a along the subduction channel). Similar, or even faster, exhumation rates have been inferred elsewhere, for instance in the Dora Maira Massif, Western Alps (Michard et al. 1993; Rubatto & Hermann 2001), the Himalayas (Chemenda et al. 2000; Treolar et al. 2003), and the Sulu region of China (Ji et al. 2003). The evidence can be accounted for by assuming upward movement of crustal slices along the subduction channel, driven by the upward buoyancy of crustal material. A rough force balance calculation confirms that this evolution is compatible with the theological properties of the materials involved. There is no need to invoke slab break-off at this stage, and there is no evidence that subduction had stopped. The fast exhumation stage brought the Ulten m~lange to depths of c. 25 km. Exhumation and cooling rates decreased by one order of magnitude during the final stages of its pre-Alpine history, coinciding with Permo-Triassic lithospheric extension, rifting, and magmatism. Such two-stage exhumation histories are common in Variscan and Alpine orogens (e.g. Spalla et al. 1996; Duch~ne et al. 1997; Lardeaux et al. 2001). Slab break-off during this stage is compatible with the evidence, and requires a relatively deep level of detachment. In conclusion, we think that the pre-Alpine evolution of the Ulten rocks provides a clear record of subduction of continental crustal material to depths of at least 100 kin. It also suggests strongly that fast exhumation preceded slab detachment, and occurred by buoyancydriven extrusion (as proposed by Chemenda et al. 1995, 1996; Ernst et al. 1997; Ernst 2001). The following slow stage was related to slab break-off (as documented by the occurrence of trondhjemitic melts; Del Moro et al. 1999) and consequent isostatic uplift and lithospheric extension. This sequence of events is compatible both with the evidence and with the rheological properties of the materials.

172

G. RANALLI ETAL.

The research of the senior author is financed by a discovery grant from NSERC (Natural Sciences and Engineering Research Council of Canada), which has also provided support for R. Mahatsente's Postdoctoral Fellowship. We thank G. V. Dal Piaz, G. Godard and P. Nimis for many discussions on Alpine geology over the years. Thorough reviews by T. V. Gerya and F. Gueydan helped with the revisions.

References BOUSQUET, R., GOFFE, B., HENRY, P., LE PICHON, X. & CHOPIN, C. 1997. Kinematic, thermal and petrological model of the Central Alps: Lepontine metamorphism in the upper crust and eclogitisation of the lower crust. Tectonophysics, 273, 105-127. BREY, G. P. & KOHLER, T. 1990. Geothermobarometry in four-phase lherzolites, II. New thermobarometers, and practical assessment of existing thermobarometers. Journal of Petrology, 31, 1352-1378, BRUEC~:NER, H. K. 1998. Sinking intrusion model for the emplacement of garnet-bearing peridotites into continent collision orogens. Geology, 26, 631-634. BUITER, S. J. H., GOVERS, R. & WORTEL, M. J. R. 2002. Two-dimensional simulations of surface deformation caused by slab detachment. Tectonophysics, 354, 195-210. BUROV, E., JOLIVET, L., LE POURHIET, L. & POLIAKOV, A. 2001. A thermomechanical model of exhumation of high pressure (HP) and ultrahigh pressure (UHP) metamorphic rocks in Alpine-type collision belts. Tectonophysics, 342, 113-136. CHEMENDA, A. I., MATTAUER,M., MALAVIEILLE,J. & BOKt;N, A. N. 1995. A mechanism for syncollisional rock exhumation and associated normal faulting: results from physical modelling. Earth and Planetary Science Letters, 132, 225 -232. CHEMENDA, A. I., MATTAUER, M. & BOKUN, A. N. 1996. Continental subduction and a mechanism for the exhumation of high-pressure metamorphic rocks: new modelling and field data from Oman. Earth and Planetary Science Letters, 143, 173-182. CHEMENDA, A. I., BURG, J.-P. & MATTAUER, M. 2000. Evolutionary model of the Himalaya-Tibet system: geopoem based on new modeling, geological and geophysical data. Earth and Planetary Science Letters, 174, 397-409. CHOPIN, C. 1984. Coesite and pure pyrope in highgrade blueschists of the Western Alps: a first record and some consequences. Contributions to Mineralogy and Petrology, 86, 107-118. CHOPIN, C. 2003. Ultrahigh pressure metamorphism: tracing continental crust into the mantle. Earth and Planetary Science Letters, 212, 1-14. CLOOS, M. & SHREVE,R. L. 1988. Subduction-channel model of prism accretion, melange formation, sediment subduction, and subduction erosion at convergent plate margins. Pure Applied Geophysics, 128, 455-499.

DAL PIAZ, G. V. 1993. Evolution of Austroalpine and Upper Penninic basement in the northwestern Alps from Variscan convergence to post-Variscan extension. In: VON RAUMER,J. F. & NEUBAUER,F. (eds) The Pre-Mesozoic Geology in the Alps. Springer, Berlin, 327-343. DAL PIAZ, G. V. 1999. The Austroalpine-Piedmont nappe stack and the puzzle of Alpine Tethys. Memorie Scienze Geologiche, Universit?t di Padova, 51, 155-176. DAVIES, J. H. & v o n BLANCKENBURG,F. 1995. Slab breakoff: a model of lithosphere detachment and its test in the magmatism and deformation of collisional orogens. Earth and Planetary Science Letters, 129, 85-102. DEL MORO, A., MARTIN, S. & PROSSER, G. 1999. Migmatites of the Ulten zone (NE Italy), a record of melt transfer in the deep crust. Journal of Petrology, 40, 1803-1826. DRURY, M. R. & FITZ GERALD, J. D. 1998. Mantle rheology: insights from laboratory studies of deformation and phase transitions. In: JACKSON, I. (ed) The Earth's Mantle - Composition, Structure, and Evolution. University Press, Cambridge, 503 -559. DUCHI~NE, S., LARDEAUX, J. M. & ALBARI~DE, F. 1997. Exhumation of eclogites: insights from depth-time path analysis. Tectonophysics, 280, 125-140. ERNST, W. G. 2001. Subduction, ultrahigh-pressure metamorphism, and regurgitation of buoyant crustal slices - implications for arcs and continental growth. Physics of Earth and Planetary Interiors, 127, 253-275. ERNST, W. G. & PEACOCK, S. M. 1996. A thermotectonic model for preservation of ultrahigh-pressure phases in metamorphosed continental crust. In: BEBOUT, G. E., SCHOLL, D. W., KIRBY, S. H. & PLATT, J. P. (eds) Subduction - Top to Bottom. American Geophysical Union, Geophysical Monograph, 96, 171-178. ERNST, W. G., LtOU, J. G. & COLEMAN, R. G. 1995. Comparative petrotectonic study of five Eurasian ultrahigh-pressure metamorphic complexes. International Geology Review, 37, 191 - 21 I. ERNST, W. G., MARUYAMA, S. & WALLIS, S. R. 1997. Buoyancy-driven, rapid exhumation of ultrahighpressure metamorphosed continental crust. Proceedings of the U. S. National Academy of Sciences, 94, 9532-9537. FLEITOUT, L. & FROIDEVAUX,C. 1982. Tectonics and topography for a lithosphere containing density heterogeneities. Tectonics, 1, 21-56. GERYA, T. V., STt3CKHERT, B. & PERCHUK, A. L. 2002. Exhumation of high-pressure metamorphic rocks in a subduction channel: a numerical simulation. Tectonics, 21, 1056, doi:10.1029/ 2002TC001406. GIUNCHI, C., SABADINI, R., BOSCHI, E. & GASPERINI, P.,1996. Dynamic models of subduction: geophysical and geological evidence in the Tyrrhenian Sea. Geophysical Journal International, 126, 555-578.

CONTINENTAL SUBDUCTION AND EXHUMATION GODARD, G. 1983. Dispersion tectonique des 6clogites de Vend6e lors d'une collision continentcontinent. Bulletin de Mindralogie, 106, 719-722. GODARD, G. & MARTIN, S. 2000. Petrogenesis of kelyphites in garnet peridotites: a case study from the Ulten zone, Italian Alps. Journal of Geodynamics, 30, 117-145. GODARD, G., MARTIN, S., PROSSER, G., KIENAST,J. R. & MORTEN, L. 1996. Variscan migmatites, eclogites and garnet peridotites of the Ulten zone, Eastern Austroalpine system. Tectonophysics, 259, 313-341. HARLEY, S. L. 1984. An experimental study of the partitioning of Fe and Mg between garnet and orthopyroxene. Contributions to Mineralogy and Petrology, 86, 359-373. HIBBITT, KARLSSON& SORENSENInc. 2001. ABAQUS User's Manual: Version 6.3. HKS, Pawtucket, Rhode Island. HONDA, S. 1985. Thermal structure beneath Tohoku, northeast Japan - a case study for understanding the detailed thermal structure of the subduction zone. Tectonophysics, 112, 69-102. HOOGERDUIJN STRATING, E. H. & VISSERS, R. L. M. 1991. Dehydration-induced fracturing of eclogitefacies peridotites: implications for the mechanical behaviour of subducting oceanic lithosphere. Tectonophysics, 200, 187-198. HYNDMAN, R. D. & WANG, K. 1993. Thermal constraints on the zone of major thrust earthquake failure: the Cascadia subduction zone. Journal of Geophysical Research, 98, 2039-2060. JI, S., SARUWATARI, K., MA1NPRICE, D., WIRTH, R., XU, Z. & X1A, B. 2003. Microstructure, petrofabrics and seismic properties of ultra high-pressure eclogites from Sulu region, China: implications for rheology of subducted continental crust and origin of mantle reflections. Tectonophysics, 370, 49-76. LARDEAUX, J. M., LEDRU, P., DANIEL, I. & DUCn~NE, S. 2001. The Variscan French Massif Central - a new addition to the ultra-high pressure metamorphic club: exhumation processes and geodynamic consequences. Tectonophysics, 332, 143-167. MAHATSENTE, R. & RANALLI, G. 2004. Time evolution of negative buoyancy of an oceanic slab subducting with varying velocity. Journal of Geodynamics, 38, 117-129. MANCKTELOW, N. S. 1995. Nonlithostatic pressure during sediment subduction and the development and exhumation of high pressure metamorphic rocks. Journal of Geophysical Research, 100, 571-583. MARTIN, S. 2003. Tectonic setting and pre-Alpine evolution of the Tonale Nappe, Eastern Austroalpine. Memorie Scienze Geologiche, Universitit di Padova, 54, 167-170. MARTIN, S., GODARD, G., PROSSER, G., SCHIAVO, A., BERNOULLI, D. & RANALLI, G. 1998. Evolution of the deep crust at the junction Austroalpine/ Southalpine: the Tonale nappe. Memorie Scienze Geologiche, Universitb di Padova, 50, 3-50.

173

MICHARD, A., CHOPIN, C. & HENRY, C. 1993. Compression versus extension in the exhumation of the Dora-Maira coesite-bearing unit, Western Alps, Italy. Tectonophysics, 221, 173-193. MILLER, C., SATIR, M. & FRANK, W. 1980. Highpressure metamorphism in the Tauern window. Mitteilungen der Osterreichischen Geologischen Gesellschaft, 71/72, 89-97. NEUBAUER, F. & VON RAUMER,J. F. 1993. The Alpine basement - linkage between the Variscides and East-Mediterranean mountain belts. In: YON RAUMER, J. F. & NEUBAUER, F. (eds) The PreMesozoic Geology in the Alps. Springer, Berlin, 641-661. NIMIS, P. & MORTEN, L. 2000. P-T evolution of 'crustal' garnet peridotites and included pyroxenites from the Nonsberg area (Upper Autroalpine), NE Italy: from the wedge to the slab. Journal of Geodynamics, 30, 93-115. NTAFLOS, T., THONI, M. & YIN, Q. 1993. Geochemie der Ultentaler ultramafitite. Mitteilungen der Osterreichischen Geologischen Gesellschaft, 138, 169-178. OBATA, M. & MORTEN, L. 1987. Transformation of spinel lherzolite to garnet lherzolite in ultramafic lenses of the Austridic crystalline complex, northern Italy. Journal of Petrology, 28, 599-623. OLESKEVICH, D. A., HYNDMAN, R. D. & WANG, K. 1999. The updip and downdip limits to great subduction earthquakes: thermal and structural models of Cascadia, south Alaska, SW Japan, and Chile. Journal of Geophysical Research, 104, 14965-14991. O'NEILL, H. S. C. & WOOD, B. J. 1979. An experimental study of F e - M g partitioning between garnet and olivine and its calibration as a geothermometer. Contributions to Mineralogy and Petrology, 70, 59-70. PEACOCK, S. M. 1996. Thermal and petrologic structure of subduction zones. In: BEBOUT, G. E., SCHOLL, D. W., KIRBY, S. H. & PLATT, J. P. (eds) Subduction - Top to Bottom. American Geophysical Union, Geophysical Monograph, 96, 119-133. RANALLI, G. 1995. Rheology of the Earth, 2nd edn. Chapman & Hall, London. RANALLI, G. 2000. Rheology and subduction of continental lithosphere. In: RANALLI,G., RICCI, C. A. & TROMMSDORF, V. (eds) Crust-Mantle Interactions. Proceedings Siena International School of Earth and Planetary Science, 21-40. RANALLI, G. 2003. A model of Palaeozoic subduction and exhumation of continental crust: the Ulten Unit, Tonale Nappe, Eastern Austroalpine. Memorie Scienze Geologiche, Universith di Padova, 54, 205-208. RANALLI, G., PELLEGR1NI, R. & D'OFFIZI, S. 2000. Time dependence of negative buoyancy and the subduction of continental lithosphere. Journal of Geodynamics, 30, 539-555. REGARD, V., FACCENNA, C., MARTINOD, J., BELLIER, O. & THOMAS, J.-C. 2003. From subduction to collision: control of deep processes on the evolution of convergent plate boundary.

174

G. RANALLI ET AL.

Journal of Geophysical Research, 108, 2208, doi: 10.1029/2002JB001943. RING, U., BRANDON, M. T., WILLETT, S. D. & LISTER, G. S. 1999. Exhumation processes, ln: RING, U., BRANDON, M. T., LISTER, G. S. & WILLETT, S. D. (eds) Exhumation Processes: Normal Faulting, Ductile Flow and Erosion. Geological Society, London, Special Publications, 154, 1-27. RUBATTO, D. & HERMANN, J. 2001. Exhumation as fast as subduction? Geology, 29, 3-6. SHREVE, R. L. & CLOOS, M. 1986. Dynamics of sediment subduction, melange formation, and prism accretion. Journal of Geophysical Research, 91, 10229-10245. SPALLA, M. I., LARDEAUX, J. M., DAL PIAZ, G. V., Gosso, G. & MESSIGA, B. 1996. Tectonic significance of Alpine eclogites. Journal of Geodynamics, 21, 257-285. THOMPSON, A. B., SCHULMANN,K. & JEZEK, J. 1997. Thermal evolution and exhumation in obliquely convergent (transpressive) orogens. Tectonophysics, 280, 171-184. THONI, M. 1999. A review of geochronological data from the Eastern Alps. Scweizerischen Mineralogie und Petrographie Mitteilungen, 79, 209-230. TREOLAR, P. J., O'BRIEN, P. J., PARRISH, R. R. & KHAN, M. A. 2003. Exhumation of early Tertiary, coesite-bearing eclogites from the Pakistan Himalaya. Journal of the Geological Society, London, 160, 367-376. TUMIATI, S. 2002. Le peridotiti a granato di Cima Vedetta Alta (Val d'Ultimo, Alto Adige). Thesis, University of Padova.

TUMIATI, S., THONI, M., NIMIS, P. & MARTIN, S. 2003. Mantle-crust interactions during Variscan subduction in the Eastern Alps (Nonsberg-Ulten zone): geochronology and new petrological constraints. Earth and Planetary Science Letters, 210, 509-526. VAN DEN BEUKEL, J. 1992. Some thermomechanical aspects of the subduction of continental lithosphere. Tectonics, 11, 316-329. VAN DEN BEUKEL, J. & WORTEL, M. J. R. 1988. Thermomechanical modeling of arc-trench regions. Tectonophysics, 154, 177-193. VAN DE ZEDDE, D. M. A. & WORTEL, M. J. R. 2001. Shallow slab detachment as a transient source of heat at midlithospheric depths. Tectonics, 20, 868-882. YON RAUMER, J. F. & NEUBAUER, F. 1993. Late Precambrian and Palaeozoic evolution of the Alpine basement - an overview. In: YON RAUMER, J. F. & NEUBAUER, F. (eds) The PreMesozoic Geology in the Alps. Springer, Berlin, 625 -639. WANG, Z. & Jt, S. 1999. Deformation of silicate garnets: brittle-ductile transition and its geological implications. Canadian Mineralogist, 37, 525-541. WANG, Z. & JI, S. 2000. Diffusion creep of finegrained garnetite: implications for the flow strength of subducting slabs. Geophysical Research Letters, 27, 2333-2336. WYLLIE, P. J. 1988. Magma genesis, plate tectonics, and chemical differentiation of the Earth. Reviews of Geophysics, 26, 370-404.

Kinematics of syneclogite deformation in the Bergen Arcs, Norway: implications for exhumation mechanisms H U G U E S R A I M B O U R G 1, L A U R E N T J O L I V E T l, L O I C L A B R O U S S E l, YVES LEROY 2 & DOV AVIGAD 3

1Laboratoire de Tectonique, UMR 7072, Universitd Pierre et Marie Curie, T 2 6 - 0 El, case 129, 4 place Jussieu, 75252 Paris cedex 05, France (e-mail: hugues, raimbourg @ lgs.jussieu.fr) 2Laboratoire de Mdcanique des Solides, Ecole Polytechnique, Palaiseau, France 3Institute of Earth Sciences, Hebrew University of Jerusalem, Jerusalem, Israel Abstract: The northwestern part of Holsncy island, in the Bergen Arcs, Norway, consists of

a granulite-facies protolith partially transformed at depth in eclogite (700 ~ > 19 kbars) and amphibolite (650 ~ 8-10 kbars) facies during the Caledonian orogenesis. Eclogitized zones are mainly planar objects (fractures with parallel reaction bands and cm-to100 m-scale shear zones). Eclogitic zones are distributed in two sets of orientations and the associated deformation can be described as 'bookshelf tectonics'. The major shear zones strike around N120 and dip to the North, and show consistent top-to-the-NE shear sense throughout the area. In the large-scale kinematic frame of Caledonian NW-dipping slab, eclogitic shear zones are interpreted as the way to detach crustal units from the subducting slab and to prevent their further sinking. As the retrograde amphibolitic deformation pattern is similar to the eclogitic one, the detached crustal units started their way up along these eclogitic shear zones. Radiometric ages of eclogitic and amphibolitic metamorphism and their comparison with the chronology of Caledonian orogenesis show that the deformation recorded on HolsnCy occurred in a convergent context. The mechanism we propose can thus account for the first steps of exhumation during collision.

Relics of high pressure (HP) and ultra high pressure (UHP) parageneses record the passage at depth of crustal units (Chopin 1984; Smith 1984; Wain 1997; Chopin & Schertl 2000; Ernst & Liou 2000; Jolivet et al. 2003). Such witnesses of a deep history, found in most current or past collision zones, can provide constraints on the deep structure and time evolution of the collision. In the Norwegian Caledonides, as well as in the Alps or in the Himalayas, plate collision was preceded by the closure of an ocean by subduction. In this context, oceanic and then continental crust can be dragged down to large depths. Downpull by the underlying lithospheric mantle in a subducting slab is indeed a widely accepted mechanism for burial of crust. On the contrary, the processes involved in the exhumation of buried tectonic units are not well understood, and several models have been designed in order to account for the return path to the surface (see review by Platt 1993). Exhumation can result from a drastic change in either vertical or horizontal boundary conditions. Andersen

et al. (1991) propose that the exhumation of HP units in the Caledonides is achieved through the eduction of an orogenic wedge, caused by the detachment of the lithospheric mantle root followed by a large-scale isostatic rebound. Alternatively, deeply buried rocks can be brought back to the surface in a reverse movement, as a result of plate kinematics change from convergence to divergence (Fossen 1992). A comparison of the ages of HolsnCy eclogites and amphibolites given by Boundy et al. (1992) and Ktihn (2002), with the ages of the eclogites in the nearby Western Gneiss Region (WGR) (review by Torsvik & Eide 1998; Carswell et al. 2003; Tucker et al. in press), indicates that Holsncy eclogites were exhuming at the same time as WGR was being buried ( 4 1 0 400 Ma). The first steps of HolsnCy eclogites exhumation occurred while the collision was still active, therefore it can only be accounted for by collisional models. Most collisional models, such as corner or wedge flow models (Jischke 1975; England &

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 175-192. 0305-8719/05/$15.00

(c) The Geological Society of London 2005.

176

H. RAIMBOURG ET AL.

Holland 1979; Cloos 1982; Shreve & Cloos 1986; see review by Mancktelow 1995), as well as analogue models (Chemenda et al. 1995), do not take into account rheological changes due to metamorphic reactions. The two-dimensional model of Burov et al. (2001) does consider metamorphic reactions caused by rock burial, but the reactions depend only on the P - T conditions and are instantaneous; that is, 100% of crustal material is transformed to eclogite as soon as the boundary of the P - T eclogitic field is crossed. The simple one-dimensional isostatic model of Dewey et al. (1993) takes into account firstorder reaction kinetics. As demonstrated by Austrheim (1987), eclogitization reactions need some fluid input to occur. As a result, the fluid supply determines which fraction of the granulitic unit can be transformed into eclogite, and the fluid transport limits the transformation rate. Eclogitic metamorphic reactions depend not only on variations in P - T conditions but also on fluid availability. Albeit not well accounted for in many models, eclogite-facies reactions are likely to have important consequences on the history of rocks affected: the complete eclogitization of the HolsnCy granulite leads to a density increase of 7%, as well as a viscosity decrease visible in the field (Austrheim 1987). The aim of this study is to analyse the deformation recorded within eclogites on Holsncy and its relations with the tectonic history of the unit during the Caledonian orogenesis. Eclogitefacies deformation is essentially non-coaxial, eclogitic structures showing a 'bookshelf' geometry with a top-to-the-NE sense of shear on the main set of shear zones. We interpret this deformation as the first step towards exhumation. We propose a geometrical model where the eclogitization reactions enable crustal slivers to detach from the downgoing slab along eclogitic shear zones and to start their way up.

Geological setting The Scandinavian Caledonides result from the closure of the Iapetus ocean and the NW-dipping continental subduction of Baltica under Laurentia, in a time spanning from late Ordovician to early Devonian periods. A series of allochtonous nappes, originating from the Iapetus ocean and from the western margin of Baltica were thrust eastward onto the autochtonous shield of Baltica (Krogh 1977; Roberts & Gee 1985). The Bergen Arcs consist of a series of arcuate thrust sheets centred around Bergen, Norway (Fig. 1). The Bergen Arcs nappe pile is in tectonic contact to the east with the parautochtonous WGR through the Bergen Arc shear zone

(BASZ, see Fig. 1). The BASZ involves parts of three tectonic units: WGR, the Kvalsida Gneiss and the Major Bergen Arc zone (Wennberg 1996). Another fault zone, the Main Caledonian Thrust (MCT), constitutes the western border of the Bergen Arcs and separates the structurally lower Oygarden Gneiss Complex from the Minor Arc and Ulriken Gneiss Complex (Fig. 1). The Lind,s nappe lies in the highest tectonic position in the Bergen Arcs nappe pile. This crustal nappe is supposed to be originally a lower crust unit and contains abundant mafic material, from gabbroic anorthosite to pure anorthosite, as well as various rocks of the charnockite suite like mangerite (Kolderup & Kolderup 1940). The northwestern part of HolsnCy island (Fig. 2a), which is the object of our study, belongs to the Lind,s nappe, and is composed of mangerites and anorthositic granulites (Austrheim & Griffin 1985), intensively reworked under eclogite and amphibolite-facies conditions. It is separated from the Meland Nappe (southeastern HolsnCy) by the Rossland Shear Zone (Fig. 2a), which is up to 2.5 km wide and shows a complex metamorphic history from amphibolitefacies to greenschist-facies conditions (Birtel et al. 1998; Schmid et al. 1998). The Meland Nappe consists of granulites with amphibolite to greenschist-facies mylonites, without any evidence of eclogite-facies metamorphism (Birtel et al. 1998).

Field relations on northwestern Holsn0y The geology of the area and its metamorphic history have been studied in detail by Austrheim and co-workers (Austrheim 1987, 1990a, b, 1994; Austrheim & Griffin 1985; Austrheim & MCrk 1988; Austrheim & Engvik 1995; Boundy et al. 1992). During the Caledonian history, the granulitic protolith has recorded successive parageneses on the exhumation path, under eclogite, amphibolite and finally greenschistfacies P - T conditions. Granulite-facies protolith

Northwestern Holsn0y granulitic anorthosites present two different fabrics. In the well-foliated parts, the granulitic foliation is defined by alternating plagioclase-rich and pyroxene-garnetrich layers. In the more coronitic parts, a plagioclase matrix encloses ellipsoidal granulitic coronas, composed of a core of orthopyroxene surrounded by layers of clinopyroxene and garnet. The strong shape fabric of the coronas

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

177

Fig. 1. Geological map of the Bergen Arcs (modified from Ragnhildsveit & Helliksen 1997). General map of Norwegian Caledonides modified from Roberts & Gee (1985). defines a granulitic lineation. Both foliation and lineation were acquired during the 900Ma Grenvillian granulite-facies deformation (Cohen et al. 1988; Bingen et al. 2001). Eclogite-facies metamorphism

The anorthositic complex was partially reworked under eclogite-facies conditions (Fig. 2b). The predominant eclogitic assemblage in anorthosite is omphacite, garnet, kyanite, zoisite, and minor phengite, + rutile, +_quartz, +__amphibole (Boundy et al. 1992). Gabbroic eclogites consist predominantly of omphacite, garnet, and minor phengite, rutile, quartz, _+carbonates. Input of fluid was necessary for eclogitic reactions to occur since both assemblages are hydrous. Therefore a shortage in available fluid resulted in the metastable preservation of part of the

granulite in the eclogitic P-T field. A heterogeneous distribution of the fluid supply enabled different grades of the transformation to be 'frozen in' (Austrheim 1987; Austrheim & MCrk 1988; see Fig. 3). Some parts of the granulitic unit are macroscopically not affected by eclogitization reactions. Elsewhere, the granulite is cut by mm-wide fractures containing hydrous eclogitic minerals and quartz. On both sides of these fractures, a dm-wide dark band corresponds to partially eclogitized granulite. A further progression in the transformation is illustrated by minor eclogitic shear zones. These minor shear zones, 10 cm to 2 m wide and a few tens of metres long, display a well-defined foliation and cut through a coherent skeleton of granulite. Where the density of the shear zones increases, they form an anastomosing network surrounding rounded, 1 to 10 m large blocks of

178

H. RAIMBOURG E T A L .

Fig. 2. (a) Geological map of Holsn~y by Birtel et al. (1998), Boundy et al. (1997a), KiJhn (2002), Ragnhildsveit & Helliksen (1997). Location of Figures 2b, 4 and 9. Coordinates correspond to UTM grid, zone 32:22 ---- 67 22 000, 81 = 2 81 000. (b) Northwestern Holsn~y (Austrheim et al. 1996; Boundy et al. 1997b) with heterogeneous distribution of eclogitic overprint on granulite. Eclogite, Breccia and Granulite correspond to degrees of eclogitic metamorphism >80%, 40-80%, and <40%, respectively. Location of Figures 5a-b, 6, 7b. Hundskjeften, Skurtveit, Lower Eldsfjellet and Upper Eldsfjellet are the four eclogitic shear zones studied.

granulite. These blocks are no longer connected and their foliations show large relative rotations. This structural type is referred to hereafter as 'eclogite breccia' (Austrheim & Mcrk 1988). Finally, roughly 100m wide eclogite shear zones that contains few granulite boudins can be followed along strike for hundreds of metres. The localization of strain in the eclogitized fraction demonstrates that the eclogite is rheologically weaker than the granulite.

Amphibolite-facies

metamorphism

Amphibolite-facies metamorphism locally affected preserved granulite-facies areas as well as eclogitized areas. Typical amphibolitic assemblages comprise amphibole, phengite, margarite, paragonite, plagioclase, zoisite and chlorite (Ktihn 2002). Structurally, the amphibolite-facies metamorphic rocks occur as: (1) patches in contact with eclogitized zones, (2) halos around quartzfeldspar veins that cross-cut eclogitized zones

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

179

Fig. 3. Schematic drawings of the different stages of eclogite-facies metamorphism associated with shear zone formation, after Boundy et al. (1992). (a) Eclogite-facies reaction zone along hydrous mineral-filled fractures cutting through granulite, without macroscopic deformation. (b) Minor eclogitic shear zone, with internal foliation and no trace of a former central fracture. The deflection of the granulitic foliation in the transition zone indicates a dextral shear sense. (c) Eclogite breccia made of rotated granulite boudins in a network of anastomosing eclogitic shear zones. Some blocks have an asymmetrical shape. (d) Eclogitic major shear zone mainly made of a well foliated and lineated eclogitic matrix and sparse granulite boudins. Photographs correspond to diagrams (a) and (c). Note the different scales.

and (3) variable size (cm to m wide) shear zones (Ktihn 2002). Textural relations, such as amphibolitic symplectites formed after omphacite, show that the eclogite-facies pre-dates the amphibolite-facies metamorphism. This is further supported by structural evidence such as crosscutting relations between amphibolitic veins and eclogite, and the dragging of eclogitic foliations into mnphibolitic shear zones where both coexist.

Greenschist-facies

On northwestern Holsn0y, the greenschist-facies retrogression is recorded only by the reactivation of eclogite-facies and amphibolite-facies fabrics (Schmid et al. 1998). Greenschist-facies metamorphism is much more conspicuous on the Meland Nappe, with the formation of shear zones of various sizes, up to 1 0 0 m thick, which cross-cut the higher grade sbear zones or

180

H. RAIMBOURG ET AL.

simply reactivate them. A late greenschist stage is also recorded by the formation of tight to open folds, especially along the southeastern coast of HolsnCy. In metagabbros and amphibolites, the typical greenschist mineralogy is actinolite and plagioclase, _+epidote, __.biotite, _+sphene, while metasyenites and metagranites are typically composed of biotite, plagioclase, _+quartz, _+epidote, _+muscovite, + chlorite, _+calcite, _ actinolite.

P-T-t evolution of LindAs Nappe The Lind,s nappe has been intensively affected by granulite-facies metamorphism associated to the Grenvillian orogeny around 900 Ma (Cohen et al. 1988; Bingen et al. 2001). During the Caledonian orogenesis, the LindAs nappe was buried and then exhumed, experiencing eclogite, amphibolite and greenschistfacies metamorphism. The granulitic protholith and eclogitized and amphibolitized zones are particularly well preserved on northwestern Holsnr Eclogitic parageneses in eclogitic shear zones of northwestern HolsnCy yield P - T conditions of 700-800 ~ and 16-19 kbars (Austrheim & Griffin 1985), 650-750 ~ and 16-19kbars (Jamtveit et al. 1990), 670 _+ 50 ~ and > 14.6 kbars (Boundy et al. 1992), and 700 ~ and 17 kbars (Mattey et al. 1994). Boundy et al. (1996) estimated P - T conditions of 690 ~ and 8-12 kbars for high-grade amphibolite in the northern and eastern areas in the Lind,s Nappe, whereas estimates by Ktihn (2002) on amphibolites on Northwest Holsnr using P - T pseudosections and amphibole thermometry, yield 600 ~ and 8 kbars. The chronology of the H P - L T metamorphism is still debated. Two scenarios can be proposed, corresponding to a metamorphism of Lind,s Nappe early in the Caledonian history, or coeval with the main collisional phase (Scandian phase, dated c. 425 Ma; Torsvik et al. 1996): 9 Ages around 460 Ma for the eclogite metamorphism (U/Pb on sphene and epidote by Boundy et al. 1997a and U/Pb on zircon by Bingen et al. 2001), followed by rapid cooling to 500 ~ at 455-445 Ma (Ar/Ar on hornblende, Boundy et al. 1996), and to 375 ~ at 430 Ma (Ar/Ar on muscovite, Boundy et al. 1996). 9 Ages around 420 Ma for the eclogite metamorphism (419 _+ 4 Ma with U/Pb on zircon on Holsncy, Bingen et al. 1998) consistent with a highly unconstrained age of 421 _+

68 Ma with Rb/Sr mineral isochrone (Cohen et al. 1988). Amphibolite-facies metamor-

phism dating yielded 4 0 9 _ 8 Ma (Rb/Sr isochron age, Austrheim 1990a) and 418 + 9 Ma for the Fonnes trondjhemic dyke intruding Proterozoic rocks under P - T conditions of 8 - 1 0 kbars and 650-700 ~ at Austrheim locality, in the northeast of Lindfis nappe (Kiihn 2002). This short time span between amphibolite and eclogitefacies metamorphism may result either from rapid exhumation or from the fact that, at the same time within Lind,s Nappe, HolsnCy was more deeply buried than Austrheim.

Kinematics of eclogitic deformation Eclogitic lineations

Eclogitic stretching lineations are present in the field at two different scales, as described by Rey et al. (1999) and Boundy et al. (1992). 9 At the mm scale, the orientation of rod-shaped omphacite defines a mineral lineation. 9 At the c m - d m scale, the ellipsoidal granulitic coronas, where eclogitized and deformed, form elongate dark mineral aggregates. At the microscopic scale, crystallographic preferred orientations (CPO) of omphacite show a strong correlation with structural directions (Boundy et al. 1992; Labrousse 2001; Bascou et al. 2002). Strain r e g i m e f r o m p r e v i o u s studies

Boundy et al. (1992) concluded that the strain regime was non-coaxial. This was based on CPO patterns of omphacites from four samples in the Lower Eldsfjellet Shear zone, but the analysis was extremely local and did not yield any sense of shear. A top-to-the-NE sense of shear was inferred from the offset of one metagabbro dyke, from offsets across minor eclogitic shear zones (Boundy et al. 1992) and from asymmetric mineral clusters and crystallization tails (Rey et al. 1999). These observations are nevertheless too few to draw any conclusion at the scale of granulite unit. We therefore extensively collected kinematic indicators in major shear zones, as well as in less transformed areas, to decipher whether consistent kinematics could be found at large scale.

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

Eclogitic deformation in large shear zones Geometry of major shear zones and lineation directions. The granulitic unit is cut by a set of 10 to 100 m large shear zones that can be followed along strike for several hundreds of metres (Fig. 4). These shear zones trend on average between N90 and N120, except east of Skurtveit and in Lower Eldsfjellet, where they swing to strike around N60. All shear zones have a northward dip between 10 and 40 ~ Our collected measurements of lineations and foliations in the major eclogitic shear zones are in agreement

181

with the data of Rey et al. (1999). On average, the lineations in Hundskjeften strike N90, N70 in Upper Eldsfjellet and N40 in Lower Eldsfjellet. The Skurtveit shear zone shows two distinct sets of lineations. The scarcity of data collected on this shear zone prevents us from trying to explain the coexistence of these two sets. In all shear zones, lineations have a northeastward plunge between 10 and 30 ~

Kinematic indicators in major shear zones. Formerly spherical objects, when deformed by

80

85

25

1 km

t

i North 22

::: .:j Eclogite !. . . . . . ; "Breccia" ....................... Granulite tr

} _,,_ 30

Eclogitic Lineation Eclogitic Foliation dip angle

Fig. 4. Eclogitic foliations and lineations in major shear zones. Arrows: sense of shear. Hund., Skurt., L.Elds. and U.Elds. refer to Hundskjeften, Skurtveit, Lower Eldsfjellet and Upper Eldsfjellet, respectively.

182

H. R A I M B O U R G ET AL.

Fig. 5. (a) Sense of shear deduced from asymmetrical shapes of granulite boudins, sigmoidal eclogite foliation around boudins and deflected granulite foliation along boudins borders. Locations on Figure 2b. (b) Sense of shear deduced from S-C structures and from asymmetrically deformed coronas or mafic aggregates. These criteria show consistent top-to-the-NE sense of shear. See locations on Figure 2b.

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

183

Fig. 5. Continued.

simple shear, get an asymmetrical shape (Passchier & Trouw 1998). At the same time the foliation in the matrix wrapping around the objects adopts a sigmoidal morphology. This morphological description, which is scaleindependent, was applied to two different classes of objects on northwestern Holsncy: eclogitized omphacite-rich former coronas embedded in a kyanite-zoisite-rich matrix ( 1 - 1 0 c m in size), and untransformed granulite boudins embedded in eclogite matrix (50 cm to 10 m in size) in eclogitic shear zones (Fig. 5a and b, respectively). The sense of shear can also be inferred from the deformation of the internal granulitic foliation in granulitic boudins. In the relatively narrow transition zone between preserved granulites in the core of the boudin and eclogitic shear zones, the granulitic foliation, where oblique to the eclogitic foliation, is bent to become parallel to the eclogitic foliation, giving shear sense along the border of granulitic boudins (Fig. 5b). The study of large-scale asymmetric features such as granulitic boudins is particularly relevant for kinematics analysis: the progressive formation of eclogitic shear zones tends to isolate large blocks, from one to several tens of metres large, of undeformed and resistant granulite in an eclogitic matrix. The pertinent

characteristic 'grain size' is thus very large and the search for regionally significant kinematic indicators requires observations at a large scale. Both within cm-scale coronas and m-scale boudins, most collected criteria yield top-tothe-NE sense of shear throughout the granulitic unit. This sense of shear is compatible with the overall sigmoidal shape of the eclogite foliation throughout the island. Although asymmetrical objects can be found in every shear zone, their density is highly variable. They are sparse in the Upper and Lower Eldsfjellet shear zones and much more ubiquitous in the Hundskjeften shear zone and to a smaller extent in the Skurtveit shear zone. In order to demonstrate this density we carried out a detailed mapping of a 100 x 100 m area in the Hundskjeften shear zone containing many metric sized boudins (Fig. 6). The asymmetrical shape of most boudins, the sigmoidal eclogitic foliation around them and the deflection of granulitic foliation along boudin borders demonstrate a consistent dextral sense of shear at the map scale. The average eclogitic foliation in this area is N120 40N and the lineation trends N80. The apparent dextral sense of shear results from top-to-the-ENE movement.

184

H. RAIMBOURG ET AL.

Fig. 6. Map of asymmetrical boudins in the Hundskjeften 'breccia' zone showing consistent top-to-the-NE sense of shear. See location on Figure 2b.

Eclogitic deformation in little transformed zones In moderately transformed zones, eclogitization is present only in 1 to 10 cm wide bands, either on both sides of a m m - w i d e fracture or as narrow shear bands without central fracture. A large proportion of eclogitic fractures - that is, ram-wide fractures with 1 to 10 c m wide eclogitic reaction bands on both sides - does not show any brittle nor ductile deformation. These 'static' fractures seem to be randomly oriented. In the rest of the eclogitic bands, the kinematics can be deduced from the offset/ deflection of granulitic foliation/lineation across the eclogitized zone. The orientations of these bands, as well as the deformation that affects them, show a very consistent structural pattern and can be separated in two main sets (Fig. 7a). The first set of bands strikes between N90 and N120, dips northeastward, and displays n o r m a l dextral m o v e m e n t associated with NE-striking lineations. The second set strikes between N30 and N80, dips northwestward, and displays n o r m a l - s i n i s t r a l m o v e m e n t associated with NW-striking lineations. Representative outcrops (Fig. 7b) show either sinistral bands (B), dextral bands (C), or both (A and B'). Where both sets of bands are present, the dextral bands are wider than the sinistral ones. The geometry and kinematics of narrow eclogitic bands suggest a bookshelf mechanism: the rhomboedric blocks of granulite are bounded by a dominant set of normal dextral shear bands and a smaller conjugate set of normal sinistral shear bands (Fig. 8).

(a)

-~--~ ........

N

Shear zones/Fractures with dextral movement Shear zones/Fractures with sinistral movement

Fig. 7. (a) Stereographic plot of the two sets of minor eclogitic shear zones/fractures. The major set (solid lines) corresponds to normal-dextral movement. The minor set (dotted lines) corresponds to normal-sinistral movement. (b) Representative outcrops showing the distribution of eclogitic fractures and narrow shear zones, with two sets of orientations (minor set striking N30-60, with apparent sinistral movement, major set striking N90-120, with apparent dextral movement). Relative movement is deduced either from deflected granulite foliation or from offsets of pyroxene-garnet lenses. Sketch (B) shows shear zones with apparent sinistral movement, sketch (C) shows eclogite-facies shear zones/fractures with apparent dextral movement. Sketches (A) and (B') show the two sets of shear zones/ fractures. See locations on Figure 2b (same location for B and B').

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

185

Fig. 7. Continued.

Moreover, the dextral set of eclogitic bands is comparable in orientation and shear direction with the major shear zones observed in the more eclogitized areas. The bookshelf-type deformation in moderately eclogitized areas is thus

Kinematics of amphibolitic deformation

Fig. 8. Bookshelf geometry with top-to-the-NE major shear zones and conjugate subordinate shear zones related to counterclockwise block rotation. Minor shear zones strike between N30 and N60, with normalsinistral sense of shear. Major shear zones strike between N90 and N120, with normal-dextral sense of shear.

Within the granulitic unit, amphibolite-facies retrogression is less conspicious in the field than the eclogite-facies metamorphism. Amphibolitic shear zones are typically 10 cm to 1 m wide and cannot be followed along strike for more than a few metres. These shear zones, which cross-cut the granulite foliation, strike in average N W - S E with a large scatter (Fig. 9). The deflection of the granulitic foliation near the shear zones gives an average top-to-the-SE sense of shear. The Rossland amphibolitic shear zone, which separates the Meland Nappe from the granulitic unit, is at least a few hundred metres wide and can be followed for hundreds of metres in an E N E - W S W direction. This shear zone contains a few asymmetric granulitic boudins indicating a dextral sense of shear, similar to the sense of shear observed in the minor shear zones in northwestern HolsnCy. Shear criteria are nevertheless relatively rare, due to an extensive greenschistfacies overprint. The dextral sense of shear is roughly compatible with the early amphibolitic top-to-the-SE phase Sd-1 described by Schmid et al. (1998) in the Rossland Shear Zone.

fully compatible with the top-to-the-NE shear sense on major shear zones.

186

H. RAIMBOURG ET AL.

Fig. 9. Map of amphibolitic deformation on northwestern Holsn0y, Rossland Shear Zone and north of Meland Nappe.

Discussion Geodynamic interpretation of the northwestern HolsnOy kinematicsgeometrical model Our field study of eclogitic shear zones reveals top-to-the-NE shear deformation. The deformation in the amphibolitic field is on average top-to-the-SE (first amphibolitic deformation (Sd-1) in Schmid et al. 1998), with a nevertheless much larger scatter than the eclogitic data. Despite these variations, there is a continuum of deformation from eclogitic to amphibolitic conditions, which leads us to interpret the eclogitic deformation as the first recorded stage of exhumation for the granulite unit.

A collisional eclogitic deformation. The history of the Caledonian orogenesis is traditionally decomposed in a collisional stage in the late Ordovician-Silurian, followed by an extensional stage starting from the Devonian. Fossen (1992) states that extension started around 400Ma with kinematics reversal along the basal decollement zone of the Caledonian nappes. As noted earlier (see 'geological setting'), there are controversies on the age of eclogite metamorphism, either 460 or 420 Ma. Whatever the age, the shortest time span of 20 Ma between all eclogite ages and the start of extensional tectonics (syn- or post-collisional) shows that the eclogitic deformation studied on HolsnCy is coeval with the collisional phase. The presence of eclogite-facies metamorphism as young as 405 + 5 Ma (Root et al. 2001; Root et al. in press) or ultra-high-pressure metamorphism

dated at 401.6 _+ 1.6 Ma (Carswell et al. 2003) or at 402 _+ 2 Ma (Tucker et al. in press) in the WGR also illustrates that while Holsn0y started to exhume, the whole orogen was experiencing convergence. HolsnOy eclogitic deformation within the Caledonides and Bergen Arcs kinematic framework. The general structure of the Caledonian orogen is thought to result from the NW-dipping subduction of Baltica under Laurentia (Krogh 1977). This subduction geometry is inferred from nappe-thrust kinematics and from variations in the degree of metamorphism in the WGR (Griffin et al. 1985). The nappes of western Norway were thrusted onto the Baltic shield toward the SE. Furthermore, the difference between the maximum burial (more than 100 km) for coesite-bearing units in the NW of the Western Gneiss Region (Smith 1984; Wain 1997; Carswell et al. 2001) and the much shallower burial (pressure peak of 13kbars, c. 40 kin) in the SE of the WGR (eclogites in the vicinity of the Solund basin; Chauvet et al. 1992; Hacker et al. 2003 is roughly in agreement with the SE-trending metamorphic gradient defined by Griffin et al. (1985). This direction is inferred to be parallel to the thrusting direction. On a more regional scale, other units of the Bergen Arcs were deformed during the Caledonian orogenesis. Within the highly sheared Major Bergen Arc, a garnet-amphibole micaschist recorded a top-to-the-SE sense of shear, corresponding to Caledonian contractional deformation (Wennberg 1996). The contractional phase in the Ulriken Gneisses and its psammitic cover is recorded in sheath-like folds and S-C

KINEMATICS OF ECLOGITE-FACIES DEFORMATION structures, both indicating thrusting toward the east (Fossen 1988). In the Oygarden Gneisses, structures indicating top-to-the-E sense of shear developed under upper greenschist-facies during the Caledonian collision (Fossen & Rykkelid 1990). Caledonian contractional structures in southwestern Norway are thus characterized by sense of shear and transport directions top-to-the-SE in the SE of the Caledonian nappes, and top-tothe-E in the Bergen Arcs (Fossen 1992). This sense of shear is roughly in agreement with the top-to-the-NE movement shown by eclogitic shear zones on northwestern Holsnoy. LateCaledonian extensional shear zones such as BASZ (Wennberg 1996) and the formation of the arcuate structure of the Bergen Arcs may have caused the slight misorientation.

Post-eclogitic rotation of northwestern HolsnOy To place the eclogitic deformation in northwestern HolsnCy within the overall Caledonian collision geometry, we also need to assess the

187

rotation that affected the granulite unit during its exhumation. In their tectonic model of Caledonian orogenesis, based on the south Norwegian Caledonides, Andersen et al. (1991) proposed that the orogenic extensional collapse was initiated by the detachment of the subducted lithospheric mantle. This resulted in isostatic rebound and eduction of deeply buried crust. Geometrically, during exhumation, the educted crust undergoes top-to-the-SE rotation (clockwise when looking NE in a N W - S E section of the subduction zone, see Fig. 10a). This general top-to-the-SE rotation, related at large scale to the NW-dipping subduction may, in detail, be expressed as top-to-the-E rotation in the Bergen Arcs, where tectonic features are slightly misorientated with respect to general geometry of the orogen (see above the distribution of shear senses, as described by Fossen (1992)). In addition, some rotations may stem from lateorogenic extension. The mode II extension proposed by Fossen (1992, 2000), is made through large and deeply-rooted normal shear zones/ faults, among which is the Bergen Arc Shear Zone. Normal movements on these listric faults cause rotations: large normal displacements

Fig. 10. (a) Model for 'top-to-the-SE rotation': the present geometry is rotated clockwise about a horizontal NE trending axis, when looking to the NE. (b) Restoration of Holsn~y in its original position within the Caledonian slab. The rotation shown on the stereographic plot corresponds to a top-to-the-W 60* rotation about a NS axis. As a result, eclogite-facies shear zones, oriented in their present position (direction: N120, dip: 20NE) with normal-dextral movement, are rotated into new position (direction: N20, dip: 50W) as reverse-dextral shear zones, corresponding to their former position in the Caledonian subducting slab.

188

H. RAIMBOURG ET AL.

along the WSW-dipping BASZ, on the eastern border of the Bergen Arc, result in top-to-ENE rotation of the nappe pile. Therefore, both processes proposed above lead roughly to topto-the-E rotation of the Lindfis nappe after eclogitic metamorphism. To restore the eclogitic shear zones to their original position at depth, within the Caledonian slab, we rotate them counterclockwise with respect to a N - S horizontal rotation axis (Fig. 10b). After about 20 ~ rotation, shear zones dip toward the W N W instead of the NE, and their normal sense of displacement becomes reverse (Fig. 10b shows the situation after an arbitrary rotation of 60~ Both the restored geometry (c. W N W dipping) of the shear zones and their reverse sense of shear are thus much more compatible with the general frame of the NW-dipping subduction.

Interpretation of eclogitic deformation.

As a result of this geometric reconstruction, we propose the following scheme for the history of crustal units affected by H P - L T metamorphism (Fig. 11). The crust, whose lower part is relatively rigid and coherent, is dragged down by the dense lithospheric mantle as the upper part of the subducting slab. During burial, P-T changes make minerals metastable and,

depending on the supply of fluids, a fraction of this crust is transformed into weaker and denser eclogite. Deformation is localized in eclogitized parts that constitute shear zones setting apart units of undeformed granulite. The crust is cut out in nappes, which can start their way up to the surface by thrusting toward the east and piling up onto each other. Such an interpretation also implies that a normal shear zone separates the rigid buttress (Laurentia) from the exhuming units within the continental accretionary wedge (see arrows along the upper boundary of the subduction wedge in Fig. 11). To our knowledge, there is no evidence for such a zone in the Bergen Arcs. An easy but not completely satisfactory explanation for this absence is that there is not a single but several units stacked in the subduction wedge. The units affected by the normal sense of shear may not have been exhumed up to the surface, unlike units affected by reverse sense deformation. Another reason for the lack of a large normal shear zone may be that the extent of reverse sense deformation within the subduction wedge is larger than the extent of normal sense deformation: reverse movements are caused by exhumation and by the displacement of the subducting slab relative to units in the subduction wedge, whereas normal movements are caused by exhumation only.

Fig. 11. Interpretative sketch of processes occurring in the buried crust. The zoom shows large eclogitic shear zones wrapping around decoupled crustal units, while eclogitic reactions propagate into untransformed crust through fractures and shear zones distributed in two main sets of orientation.

KINEMATICS OF ECLOGITE-FACIES DEFORMATION CoIlisional exhumation models Strain localization. Collisional exhumation models can be described in terms of two endmembers, penetrative corner/wedge flow models (Jischke 1975; England & Holland 1979; Cloos 1982; Shreve & Cloos 1986; Mancktelow 1995) and localizing rigid slab models (Chemenda et al. 1995). The two types of model differ first by the spatial distribution of strain (see review by Platt 1986). In wedge/corner flow models strain varies smoothly within exhumed units, whereas in Chemenda's models strain is extremely localized into two shear zones, one normal at the top and the other one reverse at the bottom, enabling a unit of undeformed crust to reach the surface after its burial. The distribution of eclogitic strain on HolsnCy is highly heterogeneous at the kilometre scale. Undeformed kmscale crustal units (i.e. untransformed granulite) are separated by highly strained 100m-scale bands (i.e. large eclogitic shear zones). This heterogeneity must be accounted for in modelling kin-scale structures and phenomena, such as subduction channel or wedge dynamics. In contrast, an average rheology that smoothes such strain-localization effects is relevant to model geodynamics of the whole orogeny. A delicate balance between shear stresses and buoyancy forces. The two end-member models also deal differently with the way shear forces are handled in exhumation processes. In Chemenda's models, exhumation is only caused by the positive buoyancy of crustal units dragged into the denser mantle. The presence of a weak decoupling layer between rigid crustal units and underlying mantle, in the sinking slab, decides whether the ascent of buoyant crust is possible or not. In corner/flow models, shear forces alone can cause upward motion: if the subduction channel that contains the subducting crust and its sedimentary cover is narrowing, the downward circulation created by basal shear stresses is balanced under certain conditions by upward flow. Despite this difference, the concept of subduction channel also assumes partial decoupling of the crustal units and the lithospheric mantle within the subducting slab. Decoupling being partially controlled by the mechanical behaviour of the buried crust, a question is to what extent a change in crustal rheology affects the processes in subduction and collision zones. Looking at complete P-T loops of a crustal unit, from near the surface to the mantle and back to the surface, a decrease in viscosity has a two-fold consequence. At depth, it increases the

189

decoupling of the crust from the downgoing lithospheric mantle and therefore it enhances its capacity to exhume. On the other hand, early exhumation of buoyant crust may prevent a large fraction of the crust from reaching and equilibrating at high pressures. The balance between both effects could only be analysed through the study of metamorphic terranes at the scale of the whole orogen, which is by far beyond the scope of this paper. In addition, on HolsnCy we do not deal with permanent rheological changes but with a change in viscosity that happens only when the crust reaches high pressures. In such circumstances, the consequences of eclogitization are quite straightforward: the rigid granulite is dragged down until it gets partially eclogitized; the eclogitized zones, weaker than the granulite, act as decoupling surfaces, which enable granulitic slivers to start their way up. Therefore, in terms of rheology, eclogitization unambiguously favours exhumation and is the first necessary step toward exhumation. An important question still to be answered about this rheological weakening is whether it is induced by hydration and grain size reduction only or also by the mineralogical change. The second major physical consequence of eclogitization is the density change, which can be up to 15% (Austrheim 1987). As a result, a completely eclogitized lower crust becomes denser than the surrounding mantle (Le Pichon et al. 1992, 1997; Dewey et al. 1993), annihilating the buoyancy force, which in all models is a major driving force to exhumation, if not the only one. In terms of exhumation, eclogitization has thus two opposite consequences. Eclogitization is an evolutionary process, in time - that is, the set of reactions progresses continuously between transformed and untransformed states - and in space - that is, some parts of the granulitic unit are still pristine while others are completely reacted. The balance between viscosity decrease and density increase, respectively promoting and impeding exhumation, depends on the way metamorphic reactions progress and propagate. The eclogitic shear zones cutting through preserved granulite, as observed on Holsn0y as well as in Flatraket (Krabbendam et al. 2000; Wain et al. 2001), enable mechanical decoupling without large density increase when averaged over the whole granulitic unit. One may imagine another situation, where very large fluid advection would induce much more pervasise eclogitization. There, the associated density increase would lead to the further sinking of eclogitized crust. Of course, such a situation is purely

190

H. RAIMBOURG ET AL.

hypothetical, since what we observe is only what has finally been exhumed; therefore, only a comprehensive mechanical modelling of eclogitization propagation could help to understand deep processes affecting rocks that finally come back up, as well those that allegedly disappear in the mantle.

Conclusions The northern part of Holsncy Island is a granulitic unit that experienced H P - L T metamorphism during the Caledonian orogenesis. This metamorphism heterogenously affected a granulite-facies protolith terrane, resulting in the juxtaposition of well-preserved granulitic zones and eclogitic zones. The eclogitic zones range from 10 cm wide reaction zones along fractures, involving only fluid diffusion from the fracture and little deformation, to eclogitized zones that are strongly deformed, forming shear zones of size ranging from 1 cm to 100 m. The comparison of P-T-t paths of the rocks on Holsncy with paths of rocks in the surrounding areas or in the W G R indicates that exhumation of the northern Holsn0y unit occurred while some close units were still being buried to large depths. A detailed study of eclogite-facies deformation in large shear zones (in particular that of asymmetrically deformed objects such as granulitic boudins or coronas) shows a consistent pattern of normal-dextral shear zones with topto-the-E sense of shear throughout northwestern Holsncy. In little transformed zones, a second set of minor shear zones, showing n o r m a l sinistral sense of shear, mimics a bookshelf mechanism. A small amount of amphibolite-facies retrogression occurs in thin shear zones crosscutting both granulite and eclogitic shear zones. The deformation in these amphibolitic shear zones reproduces roughly the eclogitic kinematic pattern, showing that eclogitic deformation recorded on Holsnoy corresponds to the first stages of exhumation. The Caledonian orogenesis resulted from the subduction of Baltica under Laurentia, in a N W - S E convergent context (Mckerrow et al. 1991; Torsvik et al. 1996). When replaced in a geodynamical context of NW-dipping subduction, the eclogitic deformation pattern can be interpreted as deep thrusting and piling-up of several crustal slivers. We thus propose the following scheme for the history of crustal units affected by H P - L T metamorphism. The crust, whose lower part was relatively rigid, was dragged down by the dense lithospheric mantle as the upper part of the

subducting slab. During burial, P-T changes made the protolith granulite-facies minerals metastable and, depending on the supply of fluids, a fraction of the crust was transformed into weaker and denser eclogite. Deformation was localized in eclogitized shear zones setting apart units of undeformed granulite. The crust was cut out into nappes that started their way up to the surface by buoyant ascent, by thrusting onto each other.

References ANDERSEN, T. B., JAMTVEIT, B., DEWEY, J. F. & SWENSSON, E. 1991. Subduction and eduction of continental crust: major mechanisms during continent-continent collision and orogenic extensional collapse, a model based on the south Norwegian Caledonides. Terra Nova, 3, 303-310. AUSTRHEIM, H. 1987. Eclogitization of lower crustal granulites by fluid migration through shear zones. Earth and Planetary Science Letters, 81, 221-232. AUSTRHEIM,H. 1990a. Fluid induced processes in the lower crust as evidenced by Caledonian eclogitization of Precambrian granulites, Bergen Arcs, Western Norway. PhD thesis, University of Oslo. AUSTRHEIM, H. 1990b. The granulite-eclogite facies transition: A comparison of experimental work and a natural occurrence in the Bergen Arcs, western Norway. Lithos, 25, 163-169. AVSTRHEIM, H. 1994. Eclogitization of the deep crust in continent collision zones. Comptes-rendus de l'Acadimie des Sciences de Paris, 319, 761-774. AUSTRHEIM, H. & GRIFFIN, W. L. 1985. Shear deformation and eclogite formation within granulitefacies anorthosites of the Bergen Arcs, Western Norway. Chemical Geology, 50, 267-281. AVSTRUEIM, H. & MORK, M. B. E. 1988. The lower continental crust of the Caledonian mountain chain: evidence from former deep crustal sections in western Norway. Norges Geologiske UndersOkelse, Special publication, 3, 102-113. AUSTRHEIM, H. & ENGVIK, A. K. 1995. Fluid transport, deformation and metamorphism at depth in a collision zone. In: JAMTVEIT, B. & YARDLEY, B. W. D. (eds) Fluid Flow and Transport in Rocks: Mechanisms and Effects. Chapman and Hall, London, 123-137. AUSTRHEIM, H., ERAMBERT, M. & ENGVlK, A. K. 1996. Processing of crust in the root of the Caledonian continental collision zone: the role of eclogitization. Tectonophysics, 273, 129-153. BASCOU, J., TOMMASI, A. & MAINPRICE, D. 2002. Plastic deformation and development of clinopyroxene lattice preferred orientations in eclogites. Geology, 24, 1357-1368. BINGEN, B., DAVIS, G. A. & AUSTRHEIM, H. 1998. Zircon growth during fluid induced Caledonian/ Scandian eclogite-facies metamorphism of the Lindfis Nappe, Caledonides of W Norway. Mineralogical Magazine, 62A, 161 - 162. BINGEN, B., DAVIS, W. J. & AUSTRHEIM, H. 2001. Zircon U-Pb geochronology in the Bergen arc

KINEMATICS OF ECLOGITE-FACIES DEFORMATION

191

eclogites and their Proterozoic protoliths, and implications for the pre-Scandian evolution of the Caledonides in western Norway. Geological Society of America Bulletin, 113, 640-649. BIRTEL, S., ALTENBERGER, U. & PASSCHIER, C. W. 1998. Polyphase structural and metamorphic evolution of the Rossland shear zone at Holsn0y, a ductile nappe boundary in the Middle Allochton of the Norwegian Caledonides. Zentralblatt fiir

CHOPIN, C. 1984. Coesite and pure pyrope in highgrade blueschists of the western Alps: a first record and some consequences. Contributions to Mineralogy and Petrology, 86, 107-118. CHOPIN, C. • SCHERTL, H. P. 2000. The UHP Unit in the Dora-Maira massif, western Alps. In: ERNST, W. G. & LIOU, J. G. (eds) Ultra-High

Mineralogie, Geologie und Paliiontologie, Teil I,

Society of America, Boulder, 133-148. CLOOS, M. 1982. Flow melanges: Numerical modelling and geologic constraints on their origin in the Franciscan complex, California. Geological Society of America Bulletin, 93, 330-345. COHEN, A. S., O'NIONS, R. K., SIEGENTHALER, R. & GRIFFIN, W. L. 1988. Chronology of the pressuretemperature history recorded by a granulite terrain. Contributions to Mineralogy and Petrology, 99, 303-311. DEWEY, J. F., RYAN, P. D. & ANDERSEN, T. B. 1993. Orogenic uplift and collapse, crustal thickness, fabrics and metamorphic phase changes: the role of eclogites. In: PRICHARD,H. M., ALABASTER,T., HARRIS, N. B. W. & NEARY, C. R. (eds) Magmatic Processes and Plate Tectonics. Geological Society of London, Special Publications, 76, 325-343. ENGLAND, P. C. & HOLLAND, T. J. B. 1979. Archimedes and the Tauern eclogites: the role of buoyancy in the preservation of exotic eclogite blocks. Earth and Planetary Science Letters, 44, 287-294. ERNST, W. G. & LIOU, J. G. 2000. Overview of UHP metamorphism and tectonics in well-studied collisional orogens. In: ERNST, W. G. & Llou, J. G. (eds) Ultra-High Pressure Metamorphism and

1-2, 19-44. BOUNDY, T. M., FOUNTAIN, D. M. & AUSTRHEIM, H. 1992. Structural development and petrofabrics of eclogite facies shear zones, Bergen Arcs, Western Norway: implications for deep crustal deformational processes. Journal of Metamorphic Geology, 10, 127-146. BOUNDY, T. M., ESSENE, E. J., HALL, C. M., AUSTRHEIM, H. & HALLIDAY, A. N. 1996. Rapid exhumation of lower crust during continentcontinent collision and late extension: Evidence from 40Ar/39Ar incremental heating of hornblendes and muscovites, Caledonian orogen, western Norway. Geological Society of America Bulletin, 108, 1425-1437. BOUNDY, T. M., MEZOER, K. & ESSENE, E. J. 1997a. Temporal and tectonic evolution of the granuliteeclogite association of the Bergen Arcs, western Norway. Lithos, 39, 159-178. BOUNDY, T. M., HALL, C. M., LI, G., ESSENE, E. J. 8r HALLIDAY, A. N. 1997b. Fine-scale isotopic heterogeneities and fluids in the deep crust; a 40Ar/39Ar laser ablation and TEM study of muscovites from a granulite-eclogite transition zone. Earth and Planetary Science Letters, 148, 223-242. BUROV, E., JOLIVET, L., LE POURHIET, L. & POLIAKOV, A. 2001. A thennomechanical model of exhumation of high pressure (HP) and ultrahigh pressure (UHP) metamorphic rocks in Alpine-type collision belts. Tectonophysics, 342, 113-136. CARSWELL, T., CUTHBERT, S., O'BRIEN, P. 8r TUCKER, R. 2001. Eclogite evolution and the timing of ultra-high pressure metamorphism in western Norway. EUG XI, Strasbourg, France. Journal of Conference Abstracts, 6, abstract 348. CARSWELL, D. A., TUCKER, R. D., O'BRIEN, P. J. 8r KROGH, T. E. 2003. Coesite micro-inclusions and the U/Pb age of zircons from the Hareidland Eclogite in the Western Gneiss Region of Norway. Lithos, 67, 181 - 190. CHAUVET, A., KIENAST, J. R., PINARDON, J. L. & BRUNEL, M. 1992. Petrological constraints and PT path of Devonian collapse tectonics within the Scandian mountain belt (Western Gneiss Region, Norway). Journal of the Geophysical Society, London, 149, 383-400. CHEMENDA,A. I., MATTAUER, M., MALAVIEILLE,J. 8r BOKUN, A. N. 1995. A mechanism for syn-collision rock exhumation and associated normal faulting: results from physical modelling. Earth and Planetary Science Letters, 132, 225-232.

Pressure Metamorphism and Geodynamics in Collision-Type Orogenic Belts. Geological

Geodynamics in Collision-Type Orogenic Belts. Geological Society of America, Boulder, 3-19. FOSSEN, H. 1988. The Ulriken gneiss complex and the Rundemanen Formation: a basement-cover relationship in the Bergen Arcs, West Norway. Norges Geologiske UndersOkelse Bulletin, 412, 67-86. FOSSEN, H. 1992. The role of extensional tectonics in the Caledonides of south Norway. Journal of Structural Geology, 14, 1033-1046. FOSSEN, H. 2000. Extensional tectonics in the Caledonides: Synorogenic or postorogenic? Tectonics, 19, 213-224. FOSSEN, H. & RYKKELID, E. 1990. Shear zone structures in the Oygarden area, West Norway. Tectonophysics, 174, 385-397. GRIFFIN, W. L., AUSTRHEIM, H., BRASTAD, K., BRYHNI, I., KRILL, A. G., KROGH, E. J., MORK, M. B. E., QVALE, H. & TORUDBAKKEN, B. 1985. High-pressure metamorphism in the Scandinavian Caledonides. In: GEE, D. G. & STURT, B. A. (eds) The Caledonide Orogen-Scandinavia and Related Areas (Vol. 2). John Wiley & Sons Ltd, Chichester, UK, 783-801. HACKER, B. R., ANDERSEN, T. B., ROOT, D. B., MEHL, L., MATTINSON, J. M. & WOODEN, J. L. 2003. Exhumation of high-pressure rocks beneath the Solund Basin, Western Gneiss Region of Norway. Journal of Metamorphic Geology, 21, 613-629.

192

H. RAIMBOURG ETAL.

JAMTVEIT, B., BUCHER-NURMINEN,K. & AUSTRHEIM, H. 1990. Fluid controlled eclogitization of granulites in deep crustal shear zones, Bergen arcs, Western Norway. Contributions to Mineralogy and Petrology, 104, 184-193. JISCHKE, M. C. 1975. On the dynamics of descending lithospheric plates and slip zones. Journal of Geophysical Research, 80, 4809-4813. JOLIVET, L., FACCENNA, C., GOFFE, B., BUROV, E. & AGARD, P. 2003. Subduction tectonics and exhumation of high-pressure metamorphic rocks in the Mediterranean orogens. American Journal of Science, 303, 353-409. KOLDERUP, C. F. & KOLDERUP, N. H. 1940. Geology of the Bergen Arc System. Bergens Museum Skrifter, 137 p. KRABBENDAM, M., WAIN, A. & ANDERSEN, T. B. 2000. Pre-Caledonian granulite and gabbro enclaves in the Western Gneiss Region, Norway: Indications of incomplete transition at high pressure. Geological Magazine, 137, 235-255. KROGH, E. J. 1977. Evidence of Precambrian continentcontinent collision in Western Norway. Nature, 267, 17-20. KOHN, A. 2002. The influence of fluid on the granulite to eclogite and amphibolite facies transition: a study in the anorthositic rocks from the Lind~ts Nappe, Bergen Arcs, West Norway. PhD thesis, University of Oslo. LABROUSSE, L. 2001. L'exhumation des roches m~tamorphiques de trks haute pression: le cas des Calddonides de Norvkge. Thrse de l'Universit~ Pierre et Marie Curie, Paris. LE PICHON, X., FOURNIER, M. & JOLIVET, L. 1992. Kinematics, topography, shortening and extrusion in the India-Eurasia collision. Tectonics, 11, 1085-1098. LE PICHON,X., HENRY, P. & GOFFE, B. 1997. Uplift of Tibet: from eclogites to granulites - implications for the Andean Plateau and the Variscan Belt. Tectonophysics, 273, 57-76. MANCKTELOW, N. S. 1995. Nonlithostatic pressure during sediment subduction and the development and exhumation of high pressure metamorphic rocks. Journal of Geophysical Research, 100, 571-583. MATTEY, D., JACKSON, D. H., HARRIS, B. W. & KELLY, S. 1994. Isotopic constraints on fluid infiltration from an eclogite facies shear zone, Holsn0y, Norway. Journal of Metamorphic Geology, 12, 311-325. MCKERROW, W. S., DEWEY, J. F. & SCOTESE, C. R. 1991. The Ordovician and Silurian development of the Iapetus ocean. Special Papers in Palaeontology, 44, 165-178. PASSCHIER, C. W. & TROUW, R. A. J. 1998. Microtectonics. 2nd ed. Springer-Verlag, Berlin, Heidelberg, 289 p. PLATT, J. P. 1986. Dynamics of orogenic wedges and the uplift of high-pressure metamorphic rocks. Geological Society of America Bulletin, 97, 1037-1053. PLATT, J. P. 1993. Exhumation of high-pressure rocks: a review of concept and processes. Terra Nova, 5, 119-133.

RAGNHILDSVEIT,J. & HELLIKSEN,D. 1997. Geologisk kart over Norge, begrunnskart Bergen, 1:250000. Norges Geologiske UndersCkelse. REY, P., FOUNTAIN, D. M. & BOUNDY, T. M. 1999. Kinematic indicator in eclogite-facies shear zones. In: SNOKE, A. W., TULLIS, J. A. and TODD, V. R. (eds) Fault-Related Rocks - A Photographic Atlas. Princeton University Press, 629 pp. RORERTS, D. & GEE, D. G. 1985. An introduction to the structure of the Scandinavian Caledonides. In: GEE, D. G. & STURT, B. A. (eds) The Caledonide Orogen: Scandinavia and Related Areas (Vol. 1). John Wiley, Chichester, 55-68. ROOT, D. B., HACKER, B. R., MATTINSON, J. M. & WOODEN, J. L. 2001. U/Pb zircon geochronology of High- and Ultrahigh-pressure eclogites of the Western Gneiss Region, Norway. l lth Annual V. M. Goldschmidt Conference, Houston, abstract 3883. ROOT, D. B., HACKER, B. R., MATTINSON, J. M. & WOODEN, J. L. 2004 Young age and rapid exhumation of Norwegian Ultrahigh-pressure rocks; an ion microprobe and chemical abrasion study. Earth and Planetary Science Letters, in press. SCHMID, R., ALTENBERGER, U. & OBERHANSLI, R. 1998. Polyphase tectonometamorphic evolution of the northwestern Lindfis Nappe on HolsnCy, Bergen Arcs, Caledonides, S-W Norway. Zentralblatt fiir Mineralogie, Geologie und Paliiontologie, Teil I, 1-2, 1-18. SHREVE, R. L. & CLOOS, M. 1986. Dynamics of sediment subduction, melange formation, and prism accretion. Journal of Geophysical Research, 91, 10.229-10.245. SMITH, D. C. 1984. Coesite in clinopyroxene in the Caledonides and its implications for geodynamics. Nature, 310, 641-644. TORSVIK, T. H. & EIDE, E. A. 1998. NGUGEOCHRO: Database and analysis for Norwegian isotope geochronology. NGU Open File Report 98.003, 54. TORSVIK, T. H., SMETHURST, M. A. et al. 1996. Continental break-up and collision in the Neoproterozoic and Palaeozoic - a tale of Baltica and Laurentia. Earth Science Reviews, 40, 229-258. TUCKER, R. D., ROBINSON, P., GEE, D. G., THORSNES, T., KROGH, T. E., NORDGULEN, 0. & BICKFORD, M. E. Thrusting and extension in the Scandian hinterland, Norway: New U-Pb ages and tectonostratigraphic evidence. American Journal of Science, in press. WAIN, A. 1997. New evidence for coesite in eclogite and gneisses: Defining an ultra-high pressure province in the Western Gneiss region of Norway. Geology, 25, 927-930. WAIN, A., WATERS, D. J. & AUSTRHEIM, H. 2001. Metastability of granulites and processes of eclogitisation in the UHP region of western Norway. Journal of Metamorphic Geology, 19, 607-623. WENNBERG, O. P. 1996. Superimposed fabrics due to reversal of shear sense: an example from the Bergen Arc Shear Zone, western Norway. Journal of Structural Geology, 18, 871-889.

Steep extrusion of late Archaean granulites in the Northern Marginal Zone, Zimbabwe: evidence for secular change in orogenic style T H O M A S G. B L E N K I N S O P 1 & A L E X A N D E R F. M. KISTERS 2

1School of Earth Sciences, James Cook University, Townsville, QLD 4811, Australia (e-mail: Thomas. Blenkinsop @jcu. edu. au) 2Department of Geology, University of Stellenbosch, Private Bag X1, Matieland 7602, South Africa Abstract: The Northern Marginal Zone (NMZ) of the Limpopo Belt in southern Zimbabwe

is an essentially plutonic terrain, now consisting of granulite-gneisses, that was assembled by extensive magmatic under- and intraplating between c. 2.7 and 2.6 Ga during late Archaean crustal shortening. In this study, we present the first evidence for regional-scale normal shear in the NMZ. Normal shear has been identified along an up to 500 m wide, steeply dipping belt of mylonites, the Mtilikwe shear zone. Geochronological and mineralogicaltextural data indicate that normal shear occurred coeval with crustal shortening that was accommodated along gently dipping thrust zones. The bulk strain pattern consisted of horizontal shortening, vertical stretching, and subvertical extrusion of a crustal segment. Horizontal shortening persisted throughout the orogenic evolution in the NMZ and no features typical of gravitational collapse in modern orogens are seen. This could imply that there was never major overthickening of the crust by thrusting, which may be a characteristic feature of Archaean orogeny.

The long-standing debate about the similarity between Archaean and modern tectonics has mostly focused on the classic granite-greenstone terranes of the Archaean cratons (e.g. Kr6ner 1981; de Wit 1998; Hamilton 2004), which exhibit a distinctive tectonic style (Choukroune et al. 1995). However, Archaean gneissic terranes, such as the Limpopo belt in southern Africa, may be more useful in this discussion because comparisons can be readily made with younger orogenic belts (e.g. Marshak 1999). The rationale for possible changes in tectonic style from the Archaean to the present is that the Earth contained more thermal energy due to retention of primordial heat and higher concentration of radioactive elements. This may have elevated crustal geotherms, although this point is debated. The concept of gravitational collapse is increasingly seen as an integral part of orogenesis (Rey et al. 2001). Gravitational collapse may occur in the late- to post-collisional evolution of orogens in response to tectonic overthickening during collision and/or the delamination or convective removal of the lower parts of the thickened lithospheric keel (e.g. Dewey 1988; England & Houseman 1988, 1989). Postconvergence gravitational collapse would be

facilitated by higher geothermal gradients in the crust, and it may also have been enhanced in the Archaean by delamination due to more vigorous mantle circulation (Sandiford 1989a). Fabrics that result from gravitational collapse of midcrustal levels are commonly characterized by flat-lying foliations, recumbent folds and low-angle thrusts that accommodate the vertical shortening and lateral spreading of the crust (e.g. Sandiford 1989b). The component of vertical shortening in the thickened orogenic belt may be enhanced by the catastrophic delamination or convective removal of the lithospheric keel during which the overlying crust rebounds, undergoing further vertical shortening (e.g. Bird 1979; Houseman et al. 1981; Platt & England 1994). Given the higher heat flow in the Archaean, Marshak (1999) suggested that Archaean orogenic belts may be narrower and lower than Phanerozoic belts for a given amount of horizontal shortening, and that they might contain relatively wider belts of plastically deformed rock with shallowly dipping extensional fabrics formed during orogenic collapse. The Northern Marginal Zone (NMZ) in southern Zimbabwe is a late Archaean, granulitegneiss belt that borders the lower-grade metamorphic granite-greenstone terrain of the

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 193-204. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

194

T.G. BLENKINSOP & A. F. M. KISTERS

Zimbabwe Craton to the north (James 1975) (Fig. 1). This narrow belt of high-grade metamorphic rocks records extensive midcrustal plutonism, high-grade metamorphism and associated contractional deformation between c. 2.7 and 2.6 Ga (e.g. Ridley 1992; Mkweli et al. 1995; Rollinson & Blenkinsop 1995; Berger et al. 1995; Kamber et al. 1996). Peakmetamorphic P - T conditions have reached 800MPa and 8 0 0 - 8 5 0 ~ (Rollinson 1989; Kamber & Biino 1995), with an anticlockwise P - T path. Given the similar present-day crustal thicknesses of the NMZ and the Zimbabwe Craton (c. 35-40 km; Stuart & Zengeni 1989; Nguuri et al. 2001), the maximum possible late Archaean crustal thickness may have been 60 km (Rollinson 1989; Kamber et al. 1996; Blenkinsop et al. 2005). The boundary between the NMZ and the Zimbabwe Craton is marked by a thrust zone, the North Limpopo Thrust Zone (Blenkinsop et al. 1995) (Fig. 1), and earlier models of crustal thickening have mainly invoked tectonic stacking during continental collision or terrane accretion and top-tothe-N directed thrusting of rocks of the NMZ onto the Zimbabwe Craton (e.g. Roering et al. 1992; de Wit et al. 1992; Rollinson 1993). More recent works, however, allow for interpreting the crustal thickening of the NMZ to be the result of a magmatic under- and intraplating by

voluminous charnoenderbitic plutons, coinciding with the main phase of craton-wide plutonism and stabilization of the Zimbabwe Craton in the outgoing Archaean (Ridley 1992; Berger et al. 1995; Rollinson & Blenkinsop 1995). The NMZ is an ideal locality to investigate possible secular changes in orogenic style related to elevated geotherms in the Archaean. No evidence has been presented that the thickened and hot crustal section of the NMZ has undergone gravitational collapse. Gneisses in the NMZ are characterized by mainly steeply dipping gneissose structure and downdip lineations, and late Archaean shear zones commonly record top-to-the-N reverse sense of movement, all consistent with the main phase of northerly vergent contractional deformation in the NMZ and the adjacent Zimbabwe Craton at c. 2.6 Ga (Rollinson & Blenkinsop 1995). This raises the question of how the NMZ escaped widespread gravitational collapse, which seems inevitable for the theologically weakened and thickened midcrustal levels. In this contribution we describe for the first time a regional-scale normal sense shear zone (the Mtlikwe shear zone) in the midcrustal rocks of the NMZ. We present evidence that normal shearing occurred under high-grade metamorphic conditions and largely coeval with the main phase of crustal shortening and

Fig. 1. The high-grade metamorphic Limpopo Belt in southern Africa, situated between the Archaean, mainly low-grade granite-greenstone terrains of the Zimbabwe and Kaapvaal Cratons. The location of the study area in the Northern Marginal Zone is shown in the white box: the box in the inset shows the location of the map.

EXTRUSION OF LATE ARCHAEAN GRANULITES thickening. Significantly, the belt of steeply dipping normal-sense mylonites is contained within the regional structural grain of the NMZ; that is, it is parallel to the regional gneissose structure and lineation patterns that have formed during the main contractional phase at c. 2.6 Ga. Normal sense shearing is only observed through the detailed observation of kinematic indicators. This suggests that regionalscale normal shear in the NMZ may be more common than hitherto recognized. The recognition of high-grade metamorphic normal shear zones in the NMZ bears on the exhumation history and mechanisms of the NMZ, the juxtaposition of the high-grade terrain against lower crustal rocks of the Zimbabwe Craton, and possibly on secular variations in tectonic style.

Regional geological setting The ENE trending NMZ forms the northern parts of the late Archaean to Mid-Proterozoic Limpopo Belt, a granulite-gneiss belt that separates the mainly low-grade metamorphic Archaean granite-greenstone terrains of the Zimbabwe Craton in the north from the Kaapvaal Craton in the south (Fig. 1). The NMZ is exposed over a strike length of c. 450 km and a width of c. 60 km in southern Zimbabwe, and comprises three main lithological components, including (1) an older (> c. 2.71 Ga) supracrustal sequence, (2) a voluminous plutonic assemblage (c. 2.712.57 Ga) made up of intrusive enderbites and charnockites, and (3) a later suite of granites and charnockites (c. 2.62-2.57 Ga), referred to as the Razi Suite (Odell 1975; Ridley 1992; Berger et al. 1995; Kamber & Biino 1995). The plutonic charnockites and enderbites, collectively referred to as the charnoenderbite suite (Berger et al. 1995), are by far the most common rocks and the NMZ represents essentially a midcrustal plutonic terrain that was assembled during a protracted period of magmatic under- and intraplating between c. 2.7 and 2.6 Ga (e.g. Ridley 1992; Berger et al. 1995; Rollinson & Blenkinsop 1995; Blenkinsop et al. 2005). Commonly rounded outlines of massive charnoenderbites are interpreted to represent the emplacement of the dry magmas as diapirs into the lower crust (Rollinson & Blenkinsop 1995). The P - T conditions determined by mineral assemblages in supracrustal enclaves from throughout the NMZ show a considerable range from 600 to 850 ~ at pressures of 400-850 MPa (Rollinson 1989; Kamber & Biino 1995). Episodes of repeated hightemperature metamorphism are related by Kamber & Biino (1995) and Kramers et al.

195

(2001) to the prolonged intrusive activity and associated heat advection. P - T results from the study area around Renco mine (Fig. 1) show a range from 600 MPa, 825 ~ near the North Limpopo thrust Zone to 820 MPa, 785 ~ 6 km south of Renco (Rollinson 1989). Various models have been proposed for the origin of the voluminous charnoenderbitic plutonism in the NMZ (Ridley 1992; Berger et al. 1995, Berger & Rollinson 1997). Ridley (1992) suggested a non-plate-tectonic setting for the charnoenderbites, invoking lower-crustal melting and magmatic accretion of the felsic magmas above a zone of mantle downwelling. Geochemical and age similarities between the charnoenderbite suite and the tonalitetrondhjemites of the Zimbabwe Craton suggest that the mid- to lower-crustal plutonic terrain exposed in the NMZ represents a deeper section of the Zimbabwe Craton (Berger et al. 1995; Rollinson & Blenkinsop 1995; Berger & Rollinson 1997). The importance of a very high geothermal gradient and extreme rate of heat generation in the NMZ that appears to be incompatible with any modern-day geological setting has been emphasized by Kramers et al. (2001). Structures in the NMZ are heterogeneously developed and the strain intensity varies from massive, largely undeformed or only weakly foliated charnoenderbites to pervasive protomylonitic and mylonitic fabrics contained in several hundred metre wide, curved high-strain zones that trend subparallel to the regional gneissose fabrics. The regional gneissic banding trends ENE, parallel to the trend of the NMZ, and dips at mainly steep angles (40-80 ~ in a southerly direction (Figs 2 and 3). A downdip mineral stretching lineation is ubiquitous in foliated rocks (Fig. 2). Protomylonites and mylonites show kinematic indicators that point to a topto-the-NNW reverse and thrust sense of movement in some of these shear zones (the medium-grade shear zones in Fig. 3) and the regional structural inventory is consistent with a bulk NNW-SSE subhorizontal shortening strain during the late Archaean (James 1975; Ridley 1992; Roering et al. 1992; Rollinson & Blenkinsop 1995; Mkweli et al. 1995; Blenkinsop et al. 1995). The contact between the high-grade gneisses of the NMZ and the Zimbabwe Craton is a more than 250 km long and several kilometre wide composite thrust zone, collectively referred to as the North Limpopo Thrust Zone (NLTZ, Blenkinsop et al. 1995) (Fig. 1). In detail, the NLTZ consists of an anastomosing network of southeasterly dipping mylonitic and protomylonitic gneisses that enclose low-strain areas in

196

T . G . B L E N K I N S O P & A. F. M. KISTERS

Fig. 2. Contoured stereoplots of poles to foliations (left, 239 data) and lineations (right, 43 data) from throughout the Northern Marginal Zone. Equal area, lower hemisphere projections. Contouring on the surface of the projection sphere; contours in multiples of a uniform distribution, starting at 1 and in intervals of 2 x multiples of a uniform distribution (Rollinson & Blenkinsop 1995).

Fig. 3. Geology of the Mtilikwe shear zone (cross-hatched) and adjacent areas in the Northern Marginal Zone of the Limpopo Belt. Partly based on Chiwara (2003). A - A ' is part of the line of section shown in Figure 10.

EXTRUSION OF LATE ARCHAEAN GRANULITES

197

which the primary magmatic textures of the igneous precursors can still be discerned. Kinematic indicators along the moderately (30-45 ~ SE-dipping NLTZ point to a top-tothe-NW reverse sense of shear, consistent with the southeasterly plunging downdip lineation in mylonites. Frei et al. (1999) have presented geochronological evidence that thrusting was diachronous along the entire length of the NLTZ and progressed from west to east with time between c. 2.65 Ga and 2.57 Ga. Syn- to late-kinematic, 2 . 6 4 - 2 . 6 G a granites of the Razi Suite constrain the timing of thrusting in the west-central portion of the NLTZ (e.g. Mkweli et al. 1995; Berger et al. 1995). An upper age bracket for thrust deformation and uplift of the NMZ along this part of the NLTZ is provided by the intrusion of the Great Dyke that has recently been constrained to 2575_+ 0.7 Ma (Oberthiir et al. 2002). This also indicates that uplift and juxtaposition of the NMZ against the low-grade granite-greenstone terrain of the Zimbabwe Craton has occurred within a very short time span, possibly as little as 10-20 Ma. The Palaeoproterozoic (c. 2 Ga) tectonism that has affected the southern parts of the NMZ (Kamber et al. 1995) is only manifest by locally developed greenschist-facies shear zones, and the structural grain preserved in the NMZ mainly records the late-Archaean tectonism between the Zimbabwe Craton and the Limpopo Belt (Blenkinsop et al. 2005). Detailed mapping of the region around the Mtilikwe shear zone is given in Chiwara (2003), which also documents the very limited extent of the Palaeoproterozoic structures.

Study methods and samples Orientated samples and structural measurements were taken throughout the known 25 km length of the Mtilikwe shear zone. The shear zone is exceptionally well exposed in the east in a river platform of the Mtilikwe river, at Rupati Pools (Figs 3 and 4a), and moderately well exposed along the length of the Gurutsime fiver and its tributaries in the west (Fig. 3). Samples were cut perpendicular to foliation and lineation, and examined by optical microscopy and SEM.

Geology of the Mtilikwe shear zone The Mtilikwe shear zone is an up-to-500 m wide east-northeasterly trending zone of protomylonites and mylonites that can be traced for 25 km along strike (Fig. 3). The belt of mylonitic

Fig. 4. The Mtilikwe shear zone: (a) Rupati Pools outcrop, showing excellent exposure and strong mylonitic foliation dipping steeply to the southeast; (b) downdip mineral lineation on foliation surface defined by quartz and feldspar; (c) Intrafolial isoclinal folds defined by garnet-rich layers.

rocks occurs 15 km to the south of the NLTZ. P b - P b step leach ages from synkinematic garnets from within the mylonites yielded a well-defined isochron indicating an age of 2601 _+ 5 Ma that was interpreted by Blenkinsop & Frei (1996) to represent the age of shearing and accompanying metamorphism. This age is within error to the main phase of thrusting recorded along the east-central parts of the

198

T.G. BLENKINSOP & A. F. M. KISTERS

NLTZ (Mkweli et al. 1995) so that the Mtilikwe shear zone was previously described as part of the regionally developed, anastomosing system of reverse and thrust zones that constitute the composite NLTZ (Blenkinsop & Frei 1996). For most of its extent, the Mtilikwe shear zone is developed in plutonic enderbites and mediumgrained granulites as well as locally developed, strongly gametiferous and sillimanite-bearing gneisses. The latter show a pronounced compositional banding on outcrop scale and are possibly of sedimentary origin (Fig. 4c). The mylonitic foliation in the Mtilikwe shear zone trends ENE and dips steeply to the SSE, parallel to the regional structural grain of the NMZ (Figs 3, 4 and 5). A prominent lineation, made up of quartz and quartz-feldspar rods, has mainly a downdip orientation in the foliation (Fig. 4b). A slight but systematic variation in the plunge of the lineation is noted along strike, changing from southerly plunges in the west to southeasterly plunges in the east of the Mtilikwe shear zone (Fig. 5). Centimetre- to decimetre-scale isoclinal, intrafolial rootless folds in the gneissic layering testify to the transposition of an earlier compositional banding or gneissic layering in probable paragneisses (Fig. 4c). The intrafolial folds plunge parallel to the downdip rodding lineation. Sheath folds were not observed. The

trend of the Mtilikwe shear zone is very closely parallel to the regional gneissic fabric, and lineations in the MSZ are parallel to the SE to S plunge of the lineations in the regional fabric (Fig. 6).

Petrology of the mylonites Protomylonites and mylonites of the Mtilikwe shear zone are composed of high-temperature mineral assemblages including quartz-plagioclase-biotite-ortho and clinopyroxene in enderbitic rocks; quartz-alkali feldspar (microcline and orthoclase)-plagioclase-biotite-ortho- and clinopyroxene in felsic granulites of probably chamockitic origin; and quartz-alkali feldspar (microcline and orthoclase)-plagioclasebiotite-garnet- sillimanite in paragneisses (Fig. 7a). Zircon, apatite, ilmenite and rutile are common accessory minerals in most samples. Texturally, the protomylonites and mylonites of the Mtilikwe shear zone are characterized by extensive dynamic recrystallization of almost all mineral components, which results in the typically finer grain size of the shear zone rocks compared to the surrounding, massive charnoenderbites. Quartz forms several centimetre long ribbons that define the mylonitic

+

Poles to foliation

Foliation

Lineation

Rupati Pools 10 km

Fig. 5. Lower hemisphere,equal area projections of mylonitic fabrics along the Mtilikwe shear zone.

EXTRUSION OF LATE ARCHAEAN GRANULITES

199

Fig. 6. Foliation trajectory map and lineations of gneissic fabrics in the study area, with the Mtilikwe shear zone shown.

foliation (Fig. 7b, c, d). Alkali feldspar and plagioclase commonly occur as augen-shaped mantled porphyroclasts (Fig. 7b, c), but they may also be pervasively recrystallized forming composite feldspar ribbons that alternate with quartz ribbons and thereby imparting a foliation-parallel compositional banding to the mylonites. Garnet appears mainly subrounded and fractured (Fig. 8). The fractures are filled by quartz, feldspar and biotite of the main mineral assemblage indicating that fracturing occurred during the high-temperature and overall ductile deformation (Fig. 8b). Orthopyroxene is only observed in enderbitic protoliths. It commonly displays a prominent undulose extinction and marginal recrystallization into smaller aggregates. The fine-grained, recrystallized orthopyroxene aggregates are locally replaced by biotite. Overall, there is little evidence of retrogression of the rocks. Minor sericite is clearly post-kinematic and can be seen to replace feldspars along cleavage planes. In summary, the extensive dynamic recrystallization of almost all mineral components and the high-grade mineral parageneses preserved in the mylonites suggest that normal shearing along the Mtilikwe shear zone has occurred

close to or slightly post-peak metamorphic granulite-facies conditions.

Kinematic indicators Macroscopic shear sense indicators along the Mtilikwe shear zone are relatively rare and are virtually restricted to relatively coarse-grained rocks such as mylonitized pegmatites. On a microscopic scale, however, shear sense indicators are common, including o'- and ~5-clasts, S-C and S-C' fabrics (Fig. 7). Twenty orientated thin sections were studied, taken along the strike extent of the Mtilikwe shear zone. Kinematic indicators in all sections consistently point to a normal, south-side down sense of shear. There was no evidence for a reverse, top-to-the-NNW sense of shear that commonly characterizes mylonites in the NMZ. A distinctive possible shear sense indicator was observed from the orientation of extension microfractures within the garnet porphyroclasts of the mylonites (Fig. 8). These have a very consistent preferred orientation relative to the mylonitic foliation, inclined at 4 0 - 5 0 ~ to the foliation towards the shortening quadrant of normal sense shearing established independently by other

200

T . G . BLENKINSOP & A. F. M. KISTERS

Fig. 7. Optical and scanning electron microscope photomicrographs of fabrics from the Mtilikwe shear zone. All sections are parallel to the lineation and perpendicular to the foliation, and have been orientated to show the same view, in which dextral shear in the photomicrographs is an extensional shear sense in the field. (a) Sillimanite fabric (dark grey, needle-like inclusions) in large garnet. Biotite (flaky mineral at the tips of the garnet), ilmenite (bright white). Matrix of quartz and K-feldspar. Rupati Pools outcrop. SEM. (b) Delta clast of feldspar in quartz ribbons, indicating dextral shear sense. Rupati Pools outcrop. Cross-polarized light. (c) Sigma clast of microperthite (Per) showing dextral shear sense, surrounded by quartz ribbons. Sample from far west end of Mtilikwe shear zone (UTM coordinates 0293790 7709090). SEM. (d) S-C fabric (lines show S and C orientations) and feldspar sigma clast showing dextral shear sense. Rupati Pools outcrop. Plane polarized light.

Fig. 8. Extension microfractures (arrows) in garnet porphyroclasts, inclined to the foliation to indicate a dextral (extensional) shear sense. Specimens from Rupati Pools outcrop. (a) Parallel alignment occurs between several porphyroclasts. Plane polarized light. (b) Microfracture fillings comprise biotite, quartz and K-feldspar. Quartz and K-feldspar define the mylonitic foliation around the garnet porphyroclast. SEM.

EXTRUSION OF LATE ARCHAEAN GRANULITES

201

Fig. 9. Microfracture orientations in garnet relative to S and C fabrics in garnetiferous mylonites from the Rupati Pools locality. The orientations of 149 microfractures were measured from over 20 garnets in a section perpendicular to the mylonitic foliation and parallel to the lineation. Shear sense was established independently from S-C fabrics, o-- and 6-porphyroclasts. Rose diagrams are moving averages in 10 windows. (a) Microfracture frequency (scale is average number of microfratures/'3. (b) Microfracture length (scale is average length, ~m/~3.

criteria (Fig. 9). A simple interpretation of these microfractures is that they are tensile fractures formed during normal simple shear.

Discussion The interpretation of the structural data presented here depends critically on whether the structures are still in approximately the same orientation in which they were formed. This can be confirmed by the observation that the Great Dyke and its satellites intruded the Northern Marginal Zone in the late Archaean, yet they are not deformed or rotated (Fig. 1; Blenkinsop et al. 2005). The significance of this observation can be extrapolated along strike to the study area because the gneissose structure in the study area has the same dip as generally observed throughout the 450 km length and 60 km width of the NMZ (Fig. 2). Geochronological constraints (Blenkinsop & Frei 1996), high-grade metamorphic mineral parageneses and deformation textures recorded in mylonites of the Mtilikwe shear zone all indicate that normal, top-to-the-S shearing occurred approximately synchronous with the late Archaean, overall contractional deformation in the NMZ at c. 2.6 Ga. Significantly, the planar and linear fabrics of the Mtilikwe shear zone are parallel to and virtually indistinguishable from the steep regional gneissose structures and downdip linear fabrics that characterize the reverse- and thrust-sense shear zones throughout

the NMZ (Figs 2, 5 and 6). The normal shear sense can only be seen at the outcrop scale in the few coarse layers, and is difficult to find in outcrop even in very well-exposed areas such as the platform in the Mtilikwe river. This means that normal-sense shear may be more widespread in the NMZ than hitherto recognized. Two important features need to be considered when discussing the syn- to late-collisional evolution of the NMZ. First, a substantial amount of crustal thickening in the NMZ was achieved by magmatic intra- and underplating that occurred over a protracted period of over 100 Ma, between c. 2.7 and 2.6 Ga (e.g. Berger et al. 1995) and late Archaean geothermal gradients were very high due to the anomalously high radiogenic heat production in the NMZ (Kramers et al. 2001). Secondly, magmatic accretion occurred during N - S crustal shortening (e.g. Ridley 1992; Rollinson & Blenkinsop 1995; Berger et al. 1995). The crustal strength must have been low throughout the evolution of the NMZ and it seems unlikely that this hot and rheologically weak crustal section could have supported large vertical loads due to tectonic thickening. Any vertical tectonic loading of the NMZ during the late Archaean NNW-directed crustal shortening was likely to have been compensated for almost instantaneously. Notably, normal sense shearing along the Mtilikwe shear zone is not a postcollisional feature, but is synchronous with the main phase of crustal thickening. Thus, significant

202

T.G. BLENKINSOP & A. F. M. KISTERS

tectonic overthickening and subsequent thermal equilibration as some of the main prerequisites for regional-scale, extensional collapse are unlikely to have occurred in the NMZ. This observation is consistent with Marshak's (1999) suggestion that Archaean orogens were probably characterized by relatively low topographic relief. While late orogenic events in modern orogens appear to involve extensional collapse and the formation of subhorizontal fabrics in the midlower crust, the latest events in the NMZ were concurrent thrusting on predominantly lowangle structures and normal faulting on steep structures. The bulk strain accommodated by the shear systems and the pervasive, subvertical fabrics and stretching lineations is NNW horizontal shortening and vertical extension. Several possibilities exist to explain this tectonic scenario. A transpressional regime with a high ratio of horizontal shortening to wrench ('pure shear dominated transpression'; Tikoff & Teyssier 1994) could account for the vertical stretch (e.g. Pelletier et al. 2002). However, there is no evidence for a wrench component to the deformation, or indeed for any orogen parallel transport. The dextral transpression seen in the Triangle Shear zone on the southern margin of the NMZ occurred 500 Ma later than the Archaean tectonics discussed here (Kamber et al. 1995). Isostatic readjustment following crustal underplating might explain the observations, but the likely locus of maximum underplating is to the south of the study area, where the highest grade metamorphic conditions were reached. The observed normal sense of shear on the Mtilikwe shear zone is the opposite of that expected for such isostatic readjustment. Horizontal gradients in vertical stretching could account for the concurrent operation of the thrusts and normal faults if the vertical stretch

was concentrated in the area between the Mtilikwe shear zone and the North Limpopo Thrust Zone. Although this is possible, there is no obvious intensification of the fabric in this area to suggest a localization of strain. Subvertical extrusion of a crustal segment during convergence is a satisfactory account for all the kinematic observations (Fig. 10). A mechanical explanation for this behaviour might be sought in the buttressing effect of the Zimbabwe craton during convergence and the specific crustal rheology of the NMZ, combined with the effect of the anisotropy induced by the gneissic fabric, which dips generally steeply to the south in the vicinity of the Mtilikwe shear zone. A clear result from this study is that the late orogenic evolution of the NMZ did not involve any of the structures that are considered typical of gravitational collapse in modern orogens. The normal faulting on the steeply dipping MSZ is not analogous to normal faulting on, for example, the South Tibetan Detachment Fault, which dips at a shallow angle (Burchfiel et al. 1992). The steep fabrics in the NMZ contrast with the subhorizontal attitude expected for midcrustal collapse features, and particularly with Marshak' s (1999) hypothesis that Archaean orogens might contain belts of subhorizontal fabrics. Horizontal shortening apparently persisted throughout the orogenic evolution not only in this part of the NMZ but elsewhere, as seen in the pervasive steeply dipping fabrics. Thus the shortening noted in the study area is not simply a manifestation of gravitational collapse in the adjacent more internal part of the Limpopo Belt. The lack of typical collapse features in the NMZ may be due to a lack of overthickening by thrusting, as described above. Choukroune et al. (1995) suggested that a lack of thrust overthickening, and the great importance of

Fig. 10. Horizontal shortening, subvertical stretching and extrusion of a crustal segment during late Archaean orogeny in the Northern Marginal Zone, Limpopo Belt. Ornaments as in Figure 3. A-A' marks line of section on Figure 3.

EXTRUSION OF LATE ARCHAEAN GRANULITES magmatic processes in crustal thickening, were characteristic of Archaean orogeny, and this contrast with modem orogenies represented a secular change in orogenic style (cf.. Chardon et al. 1998, 2002). Given the magmatically accreted, juvenile nature of the crust and the particularly high geothermal gradient that must have prevailed in the NMZ during the late Archaean (Kramers et al. 2001), the lack of evidence for gravitational collapse is one of the most distinctive aspects of Limpopo Belt geology. The example of the NMZ indicates that specific boundary conditions allowed for the N - S shortening to be accommodated by vertical extrusion of the hot and ductile crustal section without tectonic overthickening, and adds support to the concept that Archaean orogenesis was, in many aspects, different from modern orogeny.

Conclusions Late Archaean tectonics of the Limpopo Belt in the northern part of the Northern Marginal Zone involved horizontal shortening, vertical extension and subvertical extrusion of crust between gently dipping thrusts and a steeply dipping normal shear zone. The extrusion may have been controlled by the buttressing effect of the Zimbabwe craton and the steeply dipping gneissic fabrics of the NMZ. Normal shear sense is observed on careful scrutiny of steeply dipping fabrics that are parallel to the ubiquitous gneissic structure. Previous research may not have detected such normal sense structures, because they utilize fabric that is conventionally interpreted as due to thrusting with horizontal shortening. The late orogenic fabrics and tectonics of the NMZ are fundamentally different from features that characterize gravitational collapse in the late evolution of modern orogens. The lack of gravitational collapse may have been because the crust was not overthickened by thrusting; the latter may distinguish modem from Archaean orogenesis. We gratefully acknowledge the hospitality of Renco Mine and in particular the underground manager at the time, Shepherd Kadzviti. Thorough reviews by D. Chardon, P. Rey and D. Gapais improved this paper greatly.

References BERGER, M. & ROLLINSON, H. R. 1997. Isotopic and geochemical evidence for extensive intracrustal mixing and homogenization during the Archaean. Geochimicha et Cosmochimica Acta, 61, 48094829.

203

BERGER, M., KRAMERS, J. D. & NAEGLER, TH. F. 1995. Geochemistry and geochronology of charno-enderbites in the NMZ of the Limpopo Belt, Southern Africa, and genetic models. Schweizerische Mineralogische und Petrographische Mitteilungen, 75, 17-42. BIRD, P. 1979. Lateral extrusion of the lower continental crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10275-10286. BLENKINSOP,T. G., MKWELI, S. etal. 1995. The North Limpopo Thrust Zone (NLTZ): The northern boundary of the Limpopo belt in Zimbabwe and Botswana. Centennial Geocongress, Extended Abstracts, 1, 174-177. BLENKINSOP, T. G. & FREI, R. 1996. Archaean and Proterozoic mineralization and Tectonics at the Renco Mine (NMZ, Limpopo Belt), Zimbabwe. Economic Geology, 91, 1225-1238. BLENKINSOP, T. G., KRONER, A. & CHIWARA, V. 2005. Single stage, late Archaean exhumation of granulites in the Northern Marginal Zone, Limpopo Belt, Zimbabwe, and relevance to gold mineralization at Renco mine. South African Journal of Geology, in press. BURCHF1EL,B. C., CHEN, Z., HODGES, K. V., LIu, Y., ROYDEN, L. H., DENG, C. & Xu, J. 1992. The South Tibetan detachment system, Himalayan orogen: Extension contemporaneous with and parallel to shortening in a collisional mountain belt. Geological Society of America Special Paper, 269, 1-41. CHARDON, D., CHOUKROUNE, P. & JAYANANDA, M. 1998. Sinking of the Dharwar basin (South India): implications for Archaean tectonics. Precambrian Research, 91, 15-39. CHARDON, D., PEUCAT, J.-J., JAYANANDA, M., CHOUKROUNE, P. & FANNING, C. M. 2002. Archaean granite-greenstone tectonics at Kolar (South India): Interplay of diapirism and bulk inhomogeneous contraction during juvenile are accretion. Tectonics, 21, 710-717. CHIWARA, V. 2003. Late Archaean to early Proterozoic tectonics in the northern marginal zone with special emphasis on gold mineralization. Unpublished MSc thesis, University of Zimbabwe. CHOUKROUNE, P., BOUHALLIER,H. & ARNDT, N. T. 1995. Soft lithosphere during periods of Archaean crustal growth or crustal reworking. In: COWARD, M. P. & RIES, A. C. (eds) Early Precambrian Processes. Geological Society, London, Special Publications, 95, 67-86. DEWEY, J. F. 1988. Extensional collapse of orogens. Tectonics, 7, 1123-1139. DE WIT, M. J. 1998. On Archaean granites, greenstones, cratons and tectonics: does the evidence demand a verdict? Precambrian Research, 91, 181-226. DE WIT, M. J., ROERING,C., HART, R. J., ARMSTRONG, R.A., de RONDE, C. E. J. & GREEN, R. W. E. 1992. Formation of an Archean continent. Nature, 357, 553 -562. ENGLAND, P. C. & HOUSEMAN,G. A. 1988. The mechanics of the Tibetan plateau. Transactions of the Royal Society of London, A, 326, 301-320.

204

T.G. BLENKINSOP & A. F. M. KISTERS

ENGLAND, P. C. & HOUSEMAN, G. A. 1989. Extension during continental convergence, with application to the Tibetan plateau. Journal of Geophysical Research, 94, 17561-17579. FREI, R., BLENKINSOP,T. G. & SCHONBERG, R. 1999. Geochronology of the late Archaean Razi & Chilimanzi Suites of granites in Zimbabwe; implications for the late Archaean, tectonics of the Limpopo Belt and Zimbabwe Craton. South African Journal of Geology, 102, 55-63. HAMILTON, W. B. 2004. An alternative Earth. GSA Today, 13, 4-12. HOUSEMAN, G., MCKENZIE, D. P. & MOLNAR, P. 1981. Convective instability of a thickened boundary layer and its relevance for the thermal evolution of continental convergent belts. Journal of Geophysical Research, 86, 6115- 6132. JAMES, P. R. 1975. A deformation study across the

northern margin of the Limpopo Mobile belt, Rhodesia. PhD thesis, University of Leeds. KAMBER, B. S. & BIINO, G. G. 1995. The evolution of high T - l o w P granulites in the NMZ sensu stricto, Limpopo Belt, Zimbabwe - the case for petrography. Schweizerische Mineralogische und Petrographische Mitteilungen, 75,427-454. KAMBER, B. S., BLENKINSOP, T. G., VILLA, 1. M. & DAHL, P. S. 1995. Proterozoic transpressive deformation in the NMZ, Limpopo Belt, Zimbabwe. Journal of Geology, 103, 493-508. KAMBER, B. S., WIJBRANS, J. R., BIlNO, G. G., VILLA, 1. M. & DAVIES, G. 1996, Archaean granulites of the Limpopo Belt, Zimbabwe: one slow exhumation or two rapid events? Tectonics, 15, 14141430. KRAMERS, J. D., KREISSIG, K. & JONES, M. Q. W. 2001. Crustal heat production and style of metamorphism: a comparison between two Archean high grade provinces in the Limpopo Belt, Southern Africa. Precambrian Research, 112, 149-163. KRONER, A. 1981. Precambrian plate tectonics. In: KRONER, A. (ed) Precambrian Plate Tectonics. Elsevier, Amsterdam, 57-90. MARSHAK, S. 1999. Deformation style way back then: thoughts on the contrasts between Archean/Paleoproterozoic and contemporary orogens. Journal of Structural Geology, 21, 1175-1182. MKWELI, S., KAMBER, B. & BERGER, M. 1995. A westward continuation of the Zimbabwe cratonNMZ tectonic break and new age constraints on the timing of the thrusting. Journal of the Geological Society of London, 152, 77-83. NGUURI, T. K. J., GORE, D. E., JAMES, S. J., WEBB, C., WRIGHT, T. G., ZENGENI, O., GWAVAVA, J. A., SNOKE, L. & Kaapvaal Seismic Group. 2001. Crustal structure beneath southern Africa and its implications for the formation and evolution of the Kaapvaal and Zimbabwe cratons. Geophysical Research Letters, 28, 2501-2504.

OBERTHLIR, T., DAVIS, D. W., BLENKINSOP,T. G. & HOHNDORF, A. 2002. Precise U - P b mineral ages, R b - S r and S m - N d systematics for the Great Dyke, Zimbabwe - constraints on crustal evolution and metallogenesis of the Zimbabwe Craton. Precambrian Research, 113, 293-305. ODELL, J. 1975. Explanation of the geological map of the country around Bangala dam. Rhodesian Geological Survey Bulletin, 42, 46. PELLETIER, A., GAPAIS, D., MI~NOT, R.-P. & PEUCAT, J.-J. 2002. Tectonique transpressive en terre Ad~lie au Pal~oprot~rozo'ique (Est Antarctique). Comptes Rendus Geoscience, 334, 505511. PLATT, J. P. & ENGLAND,P. 1994. Convective removal of the lithosphere beneath mountain belts: thermal and mechanical consequences. American Journal of Science, 294, 307-336. REY, P., VANDERHAEGHE, 0. & TEYSS1ER, C. 2001. Gravitational collapse of the continental crust: definition, regimes and modes. Tectonophysics, 342, 435 -449. RIDLEY, J. 1992. On the origins and tectonic significance of the Charnockite suite of the Archaean Limpopo Belt, NMZ, Zimbabwe. Precambrian Research, 55, 407-427. ROERING, C., VAN REENEN, D. D. et al. 1992. Tectonic model for the evolution of the Limpopo Belt. Precambrian Research, 55, 539-552. ROLLINSON, H. R. 1989. Garnet-orthopyroxene thermobarometry of granulites from the notch marginal zone of the Limpopo belt, Zimbabwe. In: DALY, J. S., CLIFF, R. A. & YARDLEY, B. W. D. (eds) Evolution of Metamotphic Belts. Geological Society, London, Special Publications, 43, 331-335. ROLL1NSON, H. R. 1993. A terrane interpretation of the Archaean Limpopo Belt. Geological Magazine, 130, 755-765. ROLLINSON, H. R. & BLENKINSOP, T. G. 1995. The magmatic, metamorphic and tectonic evolution of the Northern Marginal Zone of the Limpopo Belt in Zimbabwe. Journal of the Geological Society of London, 152, 65 -77. SANDIFORD, M. 1989a. Secular trends in the thermal evolution of metamorphic terrains. Earth and Planetary Science Letters, 95, 85-96. SAND1FORD, M. 1989b. Horizontal structures in granulite terrains: a record of mountain building or mountain collapse? Geology, 17, 449-452. STUART, G. W. & ZENGENI, T. G. 1989. Seismic crustal structure of the Limpopo mobile belt, Zimbabwe. Tectonophysics, 144, 323-335. T1KOFF, B. & TEYSSIER, C. 1994. Strain modelling of displacement-field partitioning in transpressional orogens. Journal of Structural Geology, 16, 1575-1588.

Synergistic effects of melting and deformation: an example from the Variscan belt, western France MICHAEL BROWN

Laboratory for Crustal Petrology, Department of Geology, UniversiO' of Maryland, College Park, MD 20742-421 l, U.S.A. (e-mail: mbrown @geol. umd. edu) Abstract" In the Vannes and St. Nazaire regions of the central part of the Domaine Sud-Armoricain, Variscan belt, western France, the lowermost tectonic unit is exposed as structural culminations composed of supracrustal migmatites that were deformed and metamorphosed along a multistep clockwise P-T path during Carboniferous time; peak P-T was around 800 ~ at 9 kbar. The strain field that emerged under subsolidus conditions during prograde metamorphism controlled the initial distribution of granite melt produced by suprasolidus mica breakdown; the limited retrograde reaction of peritectic garnet indicates that melt loss occurred around the metamorphic peak. A second episode of melt production occurred during the retrograde evolution due to a decompression event that led to interconnection of melt in a mesoscale network of deformation bands and formation of ductile opening-mode fractures, as evidenced by layer-parallel and transverse leucosomes linked with petrographic continuity to granite in dykes. The preservation of peritectic cordierite with only limited associated leucosome and the occurrence of pucker structures without leucosome both indicate that melt loss occurred during the second event. Dykes vary from centimetric (common) to hundreds of metres in width (rare), and exhibit scale-invariance over a limited range of measurements; larger dykes are inferred to have fed upper crustal plutons. Melt extraction may have been a self-organized critical phenomenon, but this remains to be demonstrated satisfactorily in nature. Fugitive melt was trapped in the vicinity of the brittle-ductile transition zone and emplaced laterally along horizons reactivated as extensional detachments. A feedback relation is postulated between dextral transtensive deformation, decompression melting and lower crustal doming, and between dome amplification, melt extraction and emplacement in developing extensional detachments and core complex formation.

The infrastructure of orogens is commonly exposed as a series of core complexes; typically these are composed of high-grade metamorphic rocks - migmatites and granulites - that record evidence of melting or have residual compositions. In contrast, late orogenic granites that include a crustal component are a common feature of the suprastructure of orogens (Brown 2001a, b). Many studies have documented an intimate relationship between presence of melt and localization of deformation, and the feedback relations between melting and deformation (e.g. Brown & Rushmer 1997; Brown & Solar 1998a). Further, rock deformation experiments have shown that the presence of melt weakens rock materials (e.g. Van der Molen & Paterson 1979; Rutter 1997; Gleason et al. 1999; Holyoke & Rushmer 2002), so that a synergistic relationship between melting and/or magmatic intrusion and deformation is to be expected

(e.g. Parsons & Thompson 1993; Hill et al. 1995; McKenzie & Jackson 2002), and it has been argued that melt plays an important role in the uplift and exhumation of orogenic belts (e.g. Hoilister 1993; Brown & Dalhneyer 1996; Vanderhaeghe & Teyssier 2001; Teyssier & Whitney 2002). For these reasons, the dynamic implications of crustal anatexis and melt extraction and emplacement need to be better understood. In anatectic migmatites and residual granulites, interconnection of leucosome may be demonstrated from grain- to vein-scale (Marchildon & Brown 2001, 2002; Sawyer 2001; Guernina & Sawyer 2003); this leucosome is inferred to record remnants of the melt flow network (Brown 1994, 2001a, b; Oliver & Barr 1997; Brown et al. 1999; Sawyer 2001; Guernina & Sawyer 2003; Marchildon & Brown 2003). However, the connection between melt flow networks and magma ascent conduits remains

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 205-226. 0305-8719/05/$15.00 ~(~ The Geological Society of London 2005.

206

M. BROWN

enigmatic, although theoretical and phenomenological models have been proposed (Petford & Koenders 1998; Bons et al. 2001, 2004) and ductile fracture has been postulated (Brown 2004). Orogens are complex systems, in which it is necessary to evaluate processes holistically at various scales; complex systems evolve toward an optimal dynamic structure to dissipate energy. This is expressed at the scale of the orogen by the internal self-organization of tectonic elements and the thermal evolution (Hodges 2000) and may be expressed at the scale of the outcrop in lower crustal rocks by the patterning of structures and leucosome networks in lower crustal migmatites and granulites (Brown & Solar 1999; Guernina & Sawyer 2003; Marchildon & Brown 2003). Furthermore, core complexes within orogens generally record evidence of a positive feedback relation between decompression and crustal anatexis, which suggests that the vertical transfer of lower crust during core complex formation is an efficient mechanism for advection of heat and cooling of orogens (Teyssier & Whitney 2002). The focus of this paper is on melting in continental crust, melt extraction and ascent, and the role of fugitive melt in exhumation of the lower crust of orogens. The synergistic effects of melting and deformation are illustrated using evidence from lower crustal migmatites and leucogranite associated with Upper Carboniferous core complex formation in the Domaine SudArmoricain in the Variscan belt of western France (Gapais et al. 1993; Brown & Dallmeyer 1996; Marchildon & Brown 2003).

The Variscan belt of western France The evolution of the European Variscides was complex - a consequence of interactions between Laurasia and Gondwana (Dewey & Burke 1973; Matte 1986, 1991; Rey et al. 1997). In western France, the Variscan cycle is interpreted to have comprised Ordovician rifting to form the Galicia-Southern Brittany Ocean, Silurian subduction and consequent ocean closure in the Devonian, and Carboniferous collision between a promontory of Gondwana and Armorica that involved closeto-orthogonal sinistral transpressive deformation (Brun & Burg 1982; Pin & Peucat 1986; Burg et al. 1987; Audren 1990; Jones 1991; Dias & Ribeiro 1995; Matte 2001). Subsequent Upper Carboniferous intracontinental displacement involved highly oblique dextral transtensive deformation and exhumation of lower crustal rocks associated with the emplacement of

leucogranites (e.g. Faure & Pons 1991; Gapais et al. 1993; Geoffroy 1993; Burg et al. 1994;

Brown & Dallmeyer 1996).

Geology of the Domaine Sud-Armoricain The rocks discussed here are part of the Domaine Sud-Armoricain (DSA; Fig. 1), a NW-SE-trending belt of supracrustal rocks deformed and metamorphosed during the Variscan orogeny (Vidal 1973; Autran & Cogn6 1980; Brun & Burg 1982; Peucat 1983; Gapais et al. 1993; Brown & Dallmeyer 1996). The DSA is separated from the Domaine Centre-Armoricain, Domaine Ligerien and Domaine Nantais by the southern branch of the South Armorican Shear Zone (SASZ; Fig. 1); together these domains comprise the Variscan belt of western France (Truffert et al. 2001). The four domains most probably represent tectonostratigraphic terranes juxtaposed during the Devonian, although there are different views concerning the tectonic evolution, the significance of thrusting versus normal-sense displacements and sinistral versus dextral displacements, and the amount of displacement between terrane elements (e.g. Brun & Burg 1982; Vauchez et al. 1987; Gapais et al. 1993; Faure et al. 1997; Shelley & Bossi6re 2000; Cartier et al. 2002). T e c t o n o - s t r a t i g r a p h i c units in the D S A

In the DSA, the protracted Palaeozoic tectonic evolution ultimately resulted in extensive deformation, metamorphism and partial melting of Neoproterozoic supracrustal rocks that are now exposed as a core of migmatitic gneiss domes. In the central part of the DSA, around Vannes and St. Nazaire (Fig. 1), these migmatites preserve a mineralogical record of metamorphism and melting along an overall clockwise P - T path (peak P - T of 9 kbar, 800 ~ Brown 1983; Jones 1988; Jones & Brown 1989, 1990), with a stepped retrograde P - T segment involving decompression and a second episode of melt generation (Brown & Dallmeyer 1996). Figure 2 is a P - T pseudosection constructed for a postulated residual composition (i.e. after some melt loss around the metamorphic peak) to illustrate the retrograde evolution relevant to the late orogenic deformation discussed in this paper; evidence for the full P - T evolution is discussed in detail in Johnson & Brown (2004). This high-grade metamorphic core is bounded to the north by the NW-SE-trending transcurrent SASZ (J6gouzo 1980), and is separated from overlying units composed of lower-grade rocks to the southwest by shallow-dipping detachments

SYNERGY BETWEEN MELTING AND DEFORMATION

207

I

50 km

I

2ow

48~+ DOMAINE SUD-ARMORICAIN

- ~

Q-T Tertiary& QuaternaryRocks

G ~ W !'--

MesozoicSedimentaryRocks

~ , ~ ~ .

V .. M \'>..~

~

Blueschists r~

,

Qui

Belle-IleGroup

~

i

\

OOMAINECENTRE"4RMORICAIN F

•~

DOMAINE

v

N-E

~

F'~ VilaineGroup

LIGURIAN

'~

"~.->"

Variscangranites

DOMAINE NANTAIS

47oN~.W I-

Anatecticrocks Eciogites(Essartscomplex)

~-.~

UndifferentiatedPaleozoicRocks QSZ

SASZ [

\ St.G

LSO i 15 km i

'"~,,

m

Fig. 1. Geological map of the Domaine Sud-Armoricain, a NW-SE-trending belt of supracrustal rocks deformed and metamorphosed during the Paleozoic Variscan orogeny. The southern branch of the South Armorican Shear Zone (SASZ) separates the Domaine Sud-Armoricain from the three domains to the northwest (unomamented with domain names); together these four domains comprise the Variscan belt of western France. Granites: PA, Pony Abb& P1, Ploemur; Qui, Quiberon; Que, Questembert; S, Sarzeau; Gu, Gu&ande. Place names: L, Lorient; G, Ile de Groix; BI, Belle Ile; V, Vannes; M, Golfe du Morbihan; St.N, St. Nazaire: BC, Bois de C6n6; St.G, St. Gildas; LSO, Les Sables d'Olonne. QSZ, Quiberon Shear Zone. The W - E sketch section is modified after Gapais et al. (1993).

and associated crustally derived granites (Gapais et al. 1993). Amphibolites of the adjacent Essarts

Complex contain relics of eclogite facies rocks (15-20kbar, 600-800~ Godard 1988); the Essarts Complex is separated from the highgrade metamorphic core by tectonic contacts (Godard 2001 ). The structurally overlying Vilaine Group is dominated by apparently unmigmatized interbedded metaclastic and meta-tuffaceous rocks with amphibolite horizons; it records uppermost amphibolite facies metamorphism around the Vilaine estuary (peak P - T of 7-9 kbar, 650 :C; Triboulet & Audren 1985, 1988), but in Vend6 the metamorphic grade decreases upward through the Group (e.g. Bossibre 1988; Goujou 1992). The lower part of the overlying Belle-Ile Group is characterized by highly strained, low-grade siliceous metavolcanic and

metavolcaniclastic rocks (the Vend6e porphyroids) and metapelites (e.g. Audren & Plaine 1986; Le H6bel et al. 2002a; Schultz et al. 2002); these rocks record moderate-P-low-T metamorphism (peak P - T of 8kbar, 350400 "*C; Le Hdbel et al. 2002b). Inferred to overlie these units are the blueschist rocks and eclogites of oceanic affinity exposed on the Ile de Groix and in the Bois de Cen6 (e.g. Quinquis & Choukroune 1981; Bernard-Griffiths et al. 1986; Shelley & Bossibre 1999; Schultz et al. 2001); these rocks record high-P-low-T metamorphism (Upper Unit peak P - T of 1618 kbar, 450-500 "C; Lower Unit peak P - T of 14-16 kbar, 400-450 ~ Triboulet 1974, 1991; Guiraud et al. 1987; Djro et al. 1989; Bosse et al. 2002; Ballbvre et al. 2003). Contacts between units are flat-lying. In the central part of the DSA, the Belle-Ile Group

208

M. BROWN

Fig. 2. MnNCKFMASHP-T pseudosection from Johnson and Brown (2004) constructed for a residual bulk composition that assumes melt loss of mol.% of the melt produced at the metamorphic peak (see Johnson & Brown 2004 for further details). This diagram is only appropriate for consideration of the retrograde phase equilibria. The principal effect of melt (and H20) loss is the elevation of solidus temperatures and an increased stability of garnet. Two possible decompression P-T segments consistent with petrological and thermobarometric constraints summarized by Johnson & Brown (2004) are shown in black. Dashed lines are tool.% melt isopleths for this residual bulk composition; the darkest fields are quinivariant (F = 5).

generally is separated from the underlying highgrade rocks by subhorizontal tabular plutons of two-mica leucogranite. S-C mylonitic fabrics defined by a flat-lying schistosity, a single set of dominant shear bands and a strong mineral elongation lineation are common within these teucogranites, particularly in their lower parts (e.g. Bouchez et al. 1981; Audren 1987; Gapais et al. 1993). These features are interpreted to indicate that granite emplacement was synkinematic (cf. Gapais 1989), and S-C fabrics indicate normal-sense displacement. Following Gapais et al. (1993), the first-order normal-sense feature that is approximately parallel to the coastline is termed the Quiberon Shear Zone (QSZ; Fig. 1). The overall architecture defines windows of high-T-low-P rocks that represent structural culminations surrounded by the highP - l o w - T units in structural depressions (Fig. 1). Age relations

Recent geochronology on the porphyroids and blueschists of the Belle-Ile Group indicates that

the peak of high-P-low-T metamorphism was at 370-360 Ma, with exhumation and cooling c. 350 Ma (e.g. Bosse et al. 2000; Le H~bel 2002). In contrast, geochronological data from the high-T-moderate-to-low-P migmatites and anatectic granites of the SBMB suggest a rapid evolution in the Upper Carboniferous, since age data on monazite (U-Pb), hornblende (4~ 39Ar), mica ( R b - S r and 4~ and apatite (fission tracks) all lie in the range c. 320290 Ma (e.g. Vidal 1973; Carpena et al. 1979; Peucat 1983; Brown & Dallmeyer 1996). Similar 4~ ages of c. 300 Ma on hornblende in amphibolites from the SBMB and the Vilaine Group from the Vilaine estuary (Brown & Dallmeyer 1996) suggest these two units have a common retrograde evolution, hnprecise U - P b ages determined on metamorphic and anatectic zircon have been used to argue for an earlier age of c. 420-380 Ma for the peak of high-T metamorphism (Peucat 1983; Brown & Dallmeyer 1996), but exactly what these data record is unclear in relation to information that suggests an important Upper Carboniferous evolution for these rocks, and the greater likelihood of inheritance in zircon in comparison with monazite (e.g. Rubatto et al. 2001). One implication of these data is that the supracrustal rocks of the lower crust may have contained melt for some tens of millions of years (cf. Gerdes et al. 2000). Ages of two-mica granites emplaced synkinematically along major strike slip and extensional shear zones in South Armorica help constrain the timing of displacement along these shear zones. Bernard-Griffiths et al. (1985) documented an apparent younging of R b - S r whole-rock isochron ages of these granites from north to south across the SASZ, from c. 345 Ma to c. 300 Ma. Older ages are from granites exposed north of (c. 3 4 5 - 3 4 0 M a ) and along (c. 330-320 Ma) the SASZ; they are interpreted to record the time of crystallization and likely record the age of emplacement. Therefore, the 3 3 0 - 3 2 0 M a ages may be used to constrain timing of the initiation of the dextral strike-slip displacement along the SASZ, since these granites exhibit subhorizontal mineral elongation lineations and dextral S-C fabrics (Berth~ et al. 1979); 4~ analysis on mica from granites along the SASZ yields ages of 310-300 Ma, which are interpreted to date recrystallization during ongoing subsolidus deformation (unpublished results of Ruffet 2001, quoted in Le H~bel 2002). Two-mica granites exposed south of the SASZ occur as shallow tabular plutons along contacts between tectono-stratigraphic units; these yield R b - S r ages of c. 3 0 5 - 3 0 0 M a ,

SYNERGY BETWEEN MELTING AND DEFORMATION consistent with 4~ ages on muscovite (Brown & Dallmeyer 1996), which serve to constrain timing of the initiation of extensional displacement based on S-C fabrics in the granites (Gapais et al. 1993), and suggest rapid exhumation and cooling.

Summary

The tectono-metamorphic evolution of the DSA is composed of two principal stages (cf. Brown & O'Brien 1997): an Upper Devonian to Lower Carboniferous high-P-low-T stage that involved thrusting and piling up of nappes during a subduction to continental collision transition; and an Upper Carboniferous high-T metamorphism and associated granite magmatism related to regional sinistral transpressive and dextral transtensive deformation and core complex formation. It has been speculated that postcollisional slab detachment occurred to motivate exhumation of the lower crustal units and core complex formation in the DSA (Brown & Dallmeyer 1996). However, interpretation of newly acquired geophysical data has led to the suggestion that a relict slab remains beneath Central Armorica (Judenherc et al. 2002, 2003; Bitri et al. 2003; Gumiaux et al. 2004). If the

209

slab did not break off or if detachment did not lead to the slab sinking into the mantle to allow asthenospheric heat to reach the base of the thinned lithosphere under the over-riding plate (van de Zedde & Wortel 2001), then peak metamorphic conditions may reflect thermal relaxation of a perturbed temperature profile enhanced by radiogenic heating, and exhumation must have been motivated by some mechanism other than slab detachment.

Anatectic rocks of the Domaine Sud-Armoricain Anatectic rocks of concern in this paper are exposed around Vannes in the Golfe du Morbihan, along the coast of the Presqu'~le de Rhuys and along the coast WSW of St. Nazaire (Fig. 1). These rocks are migmatites that comprise predominantly metatexite (Fig. 3) with minor diatexite (Brown 1983; Audren 1987; Jones & Brown 1990); they are traversed by centimetric to metric dykes of granite (Fig. 4), and several 'mega-dykes' up to 1 km wide. The petrological and structural characteristics of these rocks have been described in detail by a number of authors (Cogn6 1960; Brown 1983; Audren 1987; Jones 1988; Jones & Brown 1989, 1990;

Fig. 3. Features associated with melting and inferred melt-bearing structures in migmatites from Petit Mont, Morbihan. (a) Stromatic migmatite showing foliation boudinage; (b) stromatic migmatite showing compaction (C), dilation (D) and shear (S) bands infilled with leucosome; (c) steeply oriented surface parallel to So-Sl (note subhorizontal elongation lineation) to show vertical extent of leucosome in interboudin partitions.

210

M. BROWN

Fig. 4. Features associated with melt extraction in migmatites from Petit Mont, Morbihan. (a) View roughly NW of 0.5 m-wide dykes of granite in stromatic migmatite; (b) dykes of granite frozen close to site of formation in stromatic migmatite; (e) subhorizontal surface through stromatic migmatite to show synchronous centimetric granite dykes that intersect in a common steeply inclined 'pipe-like' channel structure (with Brunton compass for scale); there is only limited evidence for cross-cutting relationships, suggesting that material now in the dykes was molten at the same time.

Audren & Triboulet 1993; Brown & Dallmeyer 1996; Marchildon & Brown 2003; Johnson & Brown 2004). The migmatitic lower crustal unit followed a multistep clockwise P-T path (Brown 1983; Jones & Brown 1990; Audren & Triboulet 1993). This involved a two-stage prograde evolution to the metamorphic peak at 9 kbar, 800 C (Johnson & Brown 2004) with production of about 25 vol.% melt via low-aH20 muscoviteand biotite-consuming reactions producing

peritectic garnet (Fig. 5a). Johnson & Brown (2004) estimate that approximately 60% of the melt produced at the metamorphic peak was extracted and lost from the system, evidence of which is the variably incomplete retrogression of peritectic garnet associated with minimal leucosome (Fig. 5a). This was followed by a stepped retrograde evolution (Fig. 2) that began with erosion controlled conductive cooling, recorded by the partial reaction of garnet with melt to form biotite and sillimanite, was interrupted by

SYNERGY BETWEEN MELTING AND DEFORMATION

211

Fig. 5. Petrographic features in stromatic migmatite that support melt loss, Presqu'i'le de Rhuys, Morbihan. (a) Virtually pristine (unretrogressed) peritectic garnet in leucosome-poor stromatic migmatite, Petit Mont; (b) stromatic migmatite, Port Navalo Plage, note partially reacted garnet in foliation (lower right-hand side) and peritectic cordierite in interboudin partition (centre left) with minor leucosome; (c) stromatic migmatite, Port Navalo Plage, note peritectic cordierite in interboudin areas and adjacent to shear bands mostly without any associated leucosome; (d) stromatic migmatite, Port Navalo Plage, note cordierite in interboudin partition without much leucosome; (e) pucker structure in which layering is 'sucked' into interboudin partitions reflecting melt loss from these sites, note thickening of leucosome stromata in boudin neck, Petit Mont; (f) folded stromatic migmatite, Port Navalo Plage, thicker leucosome contains peritectic cordierite (throughout but particularly at bottom left-hand side) to suggest that decompression melting in which biotite breaks down to produce peritectic cordierite with melt was synchronous with F3 folding.

212

M. BROWN

a steep decompression segment that involved a second episode of melt production evidenced by the presence of centimetric peritectic cordierite associated with minimal leucosome, located in dilational sites but recording melt loss (cf. Bons 1999), and was terminated by rapid cooling below the solidus (Brown & Dallmeyer 1996). The metamorphic evolution may have taken on the order of 30 m.y. or more, with much of this evolution being suprasolidus; this is consistent with residence times for melt in anatectic crust of other studies (e.g. Rubatto et al. 2001 ). Structural evolution

Early progressive sinistral transpressive deformation in the migmatites is recorded by two sets of superimposed folds, inferred to have produced flat-lying structures in the lower crust prior to doming, and later dextral transtensive deformation by folding that defines the regional map pattern of foliation form lines (cf. Brun & Burg 1982; Audren 1987; Matte 2001; Cartier et al. 2002). Isoclinal intrafolial folds (F1) of compositional layering (So) have an associated axial planar foliation (S~); the resulting composite So-S1 foliation is typically parallel to compositional layering. Folding of So-S~ produced a generation of roughly N W - SE-trending regional synformal and antiformal F2 fold structures (Cogn6 1960; Audren 1987), but a new penetrative axial planar foliation generally is not developed. F 2 folds are close to tight, with subhorizontal hinge lines and subvertical axial surfaces, which accounts for the variable dip of the So-SI composite foliation. During the early part of the late progressive dextral transtensive deformation, which also involved doming, apparent extension related to decompression melting has resulted in widespread foliation boudinage (Fig. 3a) and formation of ductile dilation and shear bands (Fig. 3b) in multilayers in migmatite, and accumulation of melt in and loss of melt from the resulting interboudin partitions and shear surfaces, evidenced by more or less discontinuous cross-cutting leucosomes in which centimetric peritectic cordierite is locally developed (Fig. 5b). The absence of significant retrogression of the cordierite, together with the limited volume of associated leucosome (Fig. 5b), which is much less than the volume of melt predicted by mass balance, indicates loss of a substantial proportion of the melt accumulated at these sites (e.g. Johnson & Brown 2004). In some cases, cordierite in host migmatite is isolated from leucosome because all melt

has migrated away from these sites (Fig. 5c, d); this conclusion is further supported by pucker structures without leucosome and thickening of leucosomes in interboudin necks (Fig. 5e). The latter part of the third deformation produced open to tight, generally asymmetric F3 folds that have steep hinge lines and subvertical axial surfaces. Regional- and decametre-scale F3 folds account for the scatter in strike of the So-Sl foliation (Cogn6 1960; Audren 1987). Centimetre-scale F3 folds are developed only locally, primarily in metatexites, and preferentially in the short limbs of metric or decametric F3 folds or in discrete zones with a dextral sense of shear; an axial planar foliation, $3, is commonly observed where F 3 folds are centimetric. F3 folding was likely coeval with dextral displacement along the SASZ during the later stages of dextral transtension and the progressive localization of strain; in zones of high late-D3 strain, F3 folds have become disrupted and isolated as intrafolial relicts. All three phases of deformation fold migmatitic layering (cf. Jones 1988, 1991). A subhorizontal or shallow-pitching quartz-feldspar aggregate lineation, inferred to record a principal finite stretch, is ubiquitous throughout the area; it is particularly well defined on foliation-parallel surfaces. In sections parallel to stromatic layering and lineation, interboudin leucosomes are perpendicular to the lineation and continuous over tens of centimetres (Fig. 3c), consistent with boudinage and the development of meltbearing interboudin partitions being synchronous with the development or the modification/reactivation of the lineation. The trend of this lineation varies with the orientation of the dominant foliation (So-S1) on which it is observed; it is folded by the F3 folds together with the stromatic leucosomes. Cordierite is present in some of the folded leucosomes (Fig. 50, which suggests that F3 folding occurred during decompression or subsequent cooling, but under suprasolidus conditions that allowed melt migration from fold limbs to hinges and melt loss along the steeply inclined hinge lines (cf. the 'passive migration' mechanism of Barraud et al. 2004). The F3 folding indicates that by late in the D 3 deformation the strike of the migmatitic layering was oriented in the shortening sector of the D3 finite strain ellipsoid. However, early D3 layerparallel extension and decompression melting, as reflected in boudinage of compositional layering and foliation, is inferred to have occurred when the migmatitic layering was oriented in the lengthening sector of the D3 infinitesimal strain ellipsoid. Rotation of the migmatitic layering both counterclockwise and back to

SYNERGY BETWEEN MELTING AND DEFORMATION approximately horizontal about a N W - S E axis would bring the perpendicular to the interboudin partitions into the lengthening sector of the infinitesimal strain ellipsoid during the early stages of dextral transtension. In this orientation the lineation may represent an early D3 feature, or possibly a D 1 - D 2 feature modified during D3 deformation. Thus, the apparent contradiction of layer-parallel extension followed by layerparallel shortening and folding of the lineation is a direct consequence of doming and core complex formation during progressive dextral transtensive deformation, which rotated the layering clockwise and from shallowly dipping to steeply dipping. During this geometrically complex progressive deformation the layering is inferred to have migrated through the lengthening sector of the infinite and finite strain ellipsoids and across the surfaces of no infinitesimal and no finite longitudinal strain into the shortening sector of the finite strain ellipsoid. On the outcrop, leucosomes generally do not show evidence of solid-state deformation. In thin section, magmatic microstructures, such as rational faces of feldspar, and quartz or feldspar films inferred to pseudomorph intergranular melt are common in leucosomes, suggesting that much of the penetrative deformation that affected these rocks either pre-dated final melt crystallization, or was partitioned away from mostly crystalline leucosomes (Brown & Dallmeyer 1996; Marchildon & Brown 2003). Thus, migmatites of the Port Navalo area appear to have hosted melt continuously from prior to or during F1 folding until F3 folding. F o r m e r m e l t - b e a r i n g structures

Leucosome- and granite-filled structures record the melt flow network; these structures include ram- to cm-scale foliation-parallel and foliation-discordant leucosome stromata (reflecting compaction-enhanced layering, interboudin partitions and extensional shear surfaces) and cm- to m-scale dykes of granite (representing inferred ascent conduits). Petrographic continuity (mineralogy, mode and microstructure) of the infilling material from structure to structure implies that some proportion of this material crystallized from a continuous melt-beating network (Figs 3 and 6). Brown (2004) has suggested that leucosome networks in migmatites and residual granulites are analogous to deformation band networks composed of shear bands (Antonellini et al. 1994), compaction bands (Mollema & Antonellini 1996) and dilation bands (Du Bernard et al. 2002); they

213

represent a melt accumulation network (Figs 3b and 7a). Using quantitative data on the one- and two-dimensional distribution of inferred meltbearing structures, Marchildon & Brown (2003) investigated the transitions from fabricparallel to network to conduit flow in an attempt to understand the evolution of the architecture (connectivity and perrneability) of meltbearing systems. They obtained information about thickness and spacing distributions of layer-parallel leucosome from one-dimensional line traverses across layering on flat surface exposures of stromatic migmatite. Layer-parallel leucosome thicknesses mostly fall in the range 1-10 mm, with an upper limit around 2 0 30 mm. The number of thicker veins decreases abruptly with increasing thickness, which is inconsistent with scale-invariance; spacing distributions also are not scale-invariant. The lack of clustering of stromata in the data of Marchildon & Brown (2003) is inconsistent with an origin by fracturing, since fractures tend to be clustered or have regularly spaced distribution around a single large fracture. An origin by mass transfer down gradients in melt pressure generated by heterogeneous deformation of an anisotropic protolith is consistent with the scale-dependant nature of the data and is preferred (cf. Brown et al. 1995). Qualitative observation of inferred meltbearing structures in mutually perpendicular two-dimensional flat surface exposures from the same outcrop reveals anisotropy of the leucosome network related to the subhorizontal stretching lineation. On lineation parallel surfaces, traces of stromatic leucosomes are continuous and the leucosomes have smooth essentially straight edges (Fig. 7a), whereas on lineation normal surfaces, traces of stromatic leucosomes are more complex, with less continuity and more irregular outlines (Fig. 7b). Overall, the three-dimensional anisotropy suggests an apparent constrictional leucosome structure with the long axis parallel to the lineation (Fig. 7c). Marchildon & Brown (2003) report results of semi-quantitative analysis of the outlines of these leucosomes in two dimensions using the box counting method, which corroborate the inferred anisotropy, and show that leucosome morphology exhibits only limited scale invariance. This result is not surprising given the strong control exerted by the anisotropy on leucosome network morphology. The anisotropic nature of the stromatic leucosomes suggests that permeability in the plane of the foliation was larger parallel to lineation than perpendicular to lineation, which further

214

M. BROWN

Fig. 6. Petrographic continuity of leucosome and granite, Petit Mont, Morbihan. (a) Dyke of granite in stromatic migmatite; (b) detail of (a), located immediately below lens cap, to show connectivity of leucosome in stromatic migmatite with granite in dyke (e.g. left of lens cap); (e) granite dyke in stromatic migmatite; (d) close-up of (c), to the left of the pen tip, to show petrographic continuity between leucosome in migmatite and granite in dyke.

supports the synchroneity of deformation and leucosome topology. It may be inferred that melt flowed in the plane of the lbliation parallel to the lineation to interboudin partitions in response to gradients in melt pressure or that

diffusive mass transfer through interconnected melt enabled melt accumulation in interboudin partitions (cf. Hand & Dirks 1992). However, the growth of peritectic phases in deformation bands, particularly interboudin partitions

SYNERGY BETWEEN MELTING AND DEFORMATION

215

Fig. 7. Binary maps of inferred melt-bearing structures in (a) subhorizontal (D, dilation band; S, shear band; and C, compaction band) and (b) subvertical exposures, and (c) to show the three-dimensional relationship and apparent constrictional topology of the leucosome, Petit Mont, Morbihan (from Marchildon, N. and Brown, M. 2003. Spatial distribution of melt-bearing structures in anatectic rocks from southern Brittany, France: implications for melt-transfer at gram-scale to orogen-scale. Tectonophysics,364, 215-235. Copyright 2003 Elsevier B.V., with permission).

(Fig. 5), connotes an important contribution by diffusive mass transfer in the local concentration of melt.

Granite dykes Centimetric to metric granite dykes are abundant at outcrop; they commonly cross-cut structures in the migmatites (Figs 4 and 6). The volumetric importance of the dykes varies in space, but they may represent up to 20% by area of an outcrop. Individual dykes are planar and typically unaffected by subsequent deformation,

except locally, indicating that emplacement post-dated the episodes of folding described above. At map scale, 'mega-dykes' of several hundreds of metres width may represent the roots of eroded subhorizontal tabular granites (e.g. Port Navalo and lie aux Moines). Granite in these dykes varies from medium- to coarse-grained. It comprises quartz, which is locally intersertal, oscillatorily zoned plagioclase (An 15- 20), K-feldspar (microcline or perthite), biotite and muscovite. In addition, cordierite, garnet and andalusite may be present, and apatite, zircon and sometimes monazite occur

216

M. BROWN

as accessory phases. The microstructure is granitic, equigranular to porphyritic, with euhedral to subhedral feldspar and anhedral quartz. Within individual dykes, spatial variation in grain size is limited and finer-grained margins generally are not observed, suggesting approximate thermal equilibrium between melt in the fracture and host melt-bearing crust. The nature of the contacts between granite in the dykes and leucosome in the host migmatites is particularly instructive. On the outcrop, the granite dykes generally show discordant contacts with host migmatites, and individual dykes intersect one another (Fig. 4c). However, the modal mineralogy and microstructure of leucosome located in layer-parallel and cross-cutting structures are indistinguishable from the modal mineralogy and microstructure of granite in dykes (Fig. 6). This petrographic continuity is interpreted to mean that these structures once hosted a continuous melt-beating network and suggests that the material in leucosomes and in the dykes underwent final crystallization at the same time. This inference does not mean that leucosomes (or necessarily granite in dykes) have liquid compositions; it is clear that leucosomes comprise residual minerals (quartz and feldspar cores), cumulate material (feldspar rims) and material crystallized from derived liquid, and that dykes record minimum conduit widths compared with the likely aperture width at peak flow. The lack of evidence of dislocation creep, demonstrated by the preservation of magmatic microstructures and the limited plastic deformation of quartz, indicates that crystallization of melt was not followed by significant penetrative deformation. These observations are interpreted to indicate that a network of connected meltbearing structures, represented by layer-parallel and discordant leucosomes and granite dykes, existed in these rocks late in the deformation history. As temperature declined below the solidus, subsequent deformation was partitioned away from the core of the DSA and appears to have been localized into the weaker Upper Carboniferous leucogranites (Hanmer 1997), where significant subsolidus strain is recorded by mylonitic fabrics (Jtgouzo 1980; Jtgouzo & Rossello 1988; Gapais et al. 1993) and is reflected in 4~ mica ages. The petrographic continuity between leucosome in the deformation band networks and granite in dykes suggests that the ascent conduits represent either structures analogous to largescale dilation bands formed by opening-mode failure along zones of localized porosity increase or ductile opening-mode fractures formed by pore growth and coalescence of melt pockets

(e.g. Regenauer-Lieb 1999; Eichhubl et al. 2001; Du Bernard et al. 2002; Eichhubl & Aydin 2003). The dykes fill macroscopic fracture-like discontinuities with a preferred orientation independent of anisotropy, which suggests that dyke orientation is controlled by stress; this feature characterizes the process of dyke formation as a fracture phenomenon. Additional characteristic features of the dykes include the nature of the tips, which are blunter than expected for brittle fracture, the open zigzag geometry close to the tips (Fig. 4b), and the petrographic continuity between leucosome in host and granite in dykes. In particular, the zigzag propagation paths at the tips of dykes, which may originate from transverse structures such as interboudin partitions and shear surfaces, point to ductile fracture as the mode of formation of these conduits (Brown 2004). The thickness of dykes > 10 cm wide shows a power law distribution with an exponent of 1.11 (Fig. 8), comparable to values of 1.1 reported by Bons et al. (2004; who also reported an exponent of 1.9 for one suite of veins). This width-distribution relationship suggests the dykes may be scale-invariant, although the dataset is small (87 dykes) and the range of observations is only two orders of magnitude (cf. Bons et al. 2004). The largest dykes measured along the traverse were 3 m and 5.5 m wide, within the range of critical dyke widths for flowing magma to

1000

exponent = 1.11

r ii Q

z E

10

,,,,i

~.

J

1

10

100

Thickness (cm) Fig. 8. Cumulative frequency plots of dyke thickness (n = 87) from a traverse along the coastline around Port Navalo Plage and le Petit Mont, Morbihan. The number of dykes wider than a particular value is plotted against that value, and a best fit line through the data above 10 cm wide defines a power law relationship with an exponent of 1.11.

SYNERGY BETWEEN MELTING AND DEFORMATION advect heat faster than conduction through the walls and avoid freezing close to the source (Clemens 1998). Dykes of this size each may transport > 1 0 0 0 k m 3 of magma in c. 1200 years, assuming a horizontal length of 1 km (Clemens 1998). This fugitive melt is inferred to have fed crustally-derived plutons at higher crustal levels, such as the Ploemeur, Quiberon, Sarzeau, and Gu~rande granites.

Structural analysis of former melt-bearing structures and dykes Figure 9 shows the orientation of the primary metamorphic fabric and migmatite layering (So-S1), F3 folds, cumulative plots of the strike orientation of centimetric leucosome-filled shear bands and the orientation of centimetric

217

to metric granite dykes. The distribution and orientation of small-scale shear surfaces is approximately congruent with layer-parallel extension as recorded by petrographically continuous leucosome-filled interboudin partitions, assuming rotation of the migmatitic layering both counterclockwise and back to approximately horizontal about a N W - S E axis (as described earlier). In the DSA migmatites the dykes show a range of orientations; most trend approximately NNW and dip steeply to moderately to the W, defining a point maximum (Fig. 9). Audren (1987) interpreted similar results as reflecting the synchroneity of dyke formation with dextral strike-slip deformation along the SASZ, the NNW trend being at low angle to the short axis (i.e. minimum principal finite stretch) of the regional

Fig. 9. The orientation of structural elements related to Carboniferous deformation measured in outcrops around the Presqu'~le de Rhuys, Morbihan (equal area lower hemisphere projection).

218

M. BROWN

finite strain ellipsoid associated with this deformation. The geometrical relationship between the overall strike of granite dykes and the orientation of the SASZ is shown in Figure 10. It has been argued that in the transition from grain-scale melt flow to channelized flow in veins the direction of melt flow will be controlled by elements of the metamorphic fabric, particularly the foliation and lineation (Brown & Solar 1998a; Brown et aL 1999). In zones of highly oblique transtensive deformation, melt flow is expected to be either subhorizontal along the lineation in the plane of the foliation (Brown & Solar 1998a) or subvertical in discordant structures (cf. Sibson 1996; Sibson & Scott 1998). Marchildon & Brown (2003) have argued that melt flow in the DSA migmatites was along the layering parallel to the quartz-feldspar mineral aggregate lineation to interboudin partitions, shear surfaces and dykes. Upward transport of melt is inferred to have been principally through dykes or other backbone structures such as the 'pipe-like' structure formed by intersection of coeval but small dykes (Fig. 4c).

Discussion M i g m a t i t e - g r a n i t e relations

Granites emplaced in the upper crust are a necessary complement to melt-depleted lower

Fig. 10. Geometricrelationshipbetween the strike orientation of granite dykes (lighter grey ornament), the SASZ (darker grey ornament) and the crenulation cleavage associated with S3 folds.

crust (Fyfe 1973; Brown & Solar 1999); they represent a tangible result of the dissipation of strain and thermal energy. There is a tendency to separate the processes of melt generation and segregation from ascent and emplacement. However, this separation is artificial - a simple succession of events is unlikely and feedback relations are to be expected (Brown 2001a, b). Melt extraction may be a self-organized critical phenomenon in which the critical point is the threshold for melt escape from the anatectic lower crust (e.g. Handy et al. 2001; Bons et al. 2002). In actively deforming anatectic systems, the strain field that emerged under subsolidus conditions during prograde metamorphism will control the initial pattern of melt flow (Brown & Solar 1998a; Brown et al. 1999). As melt volume increases, ongoing deformation of the heterogeneous crust leads to development of an interconnected melt flow network and melt expulsion. In the DSA, the development of centimetric peritectic cordierite in interboudin partitions implies the second episode of melting was driven by decompression. The appropriate biotite dehydration-driven melting reaction involves an increase in molar volume of about 7% in crossing from the high P to low P side (Fig. 2, A or B). The dilational strain associated with progress of this reaction and increasing volume of melt, in combination with ongoing transtensive deformation and doming is likely to have motivated the interconnection of the deformation band network (cf. Grujic & Mancktelow 1998; Mandal et al. 2004). However, some parts of the remnant melt flow networks appear to exhibit scaledependent behaviour (e.g. Marchildon & Brown 2003), which conflicts with the scale-invariant behaviour expected from self-organized critical phenomena (e.g. Bons et al. 2004). It is likely that the scale-dependent behaviour is a function of intrinsic anisotropy in the system. Whether or not the extraction of melt is a self-organized critical phenomenon, the patterning of leucosome in residual lower crust provides evidence of the physical process by which this is accomplished in nature (Solar & Brown 2001; Brown 2004). Draining of melt from deformation bands and capture of small conduits by large conduits will be favoured by the lower melt pressure in the larger conduits. However, the anisotropy of stromatic leucosomes and their small thickness (millimetric to centimetric) impose a limit on effective melt transport, and in strike-slip (transpressive and transtensive) systems in particular, stromatic leucosomes cannot have been responsible for subvertical melt transport (Rutter

SYNERGY BETWEEN MELTING AND DEFORMATION 1997). Rather, in the migmatites of the DSA, stromatic leucosomes fed steeply dipping crosscutting leucosomes and granite dykes that formed synchronously with late Carboniferous (D3) dextral displacement along the SASZ. Melt flow appears to have occurred through a deformation band network at outcrop scale and in dykes at crustal scale. Dyke growth is governed primarily by the ratio of elastic to viscous response of the host rock. As viscosity contrast is reduced with increasing T and proportion of melt, the viscous response dominates, although the resulting intrusions are dyke-like. The petrographic continuity across apparently discordant contacts between granite in dykes and the leucosomes in the stromatic migmatites indicates that stromata, interboudin partitions, shear bands and dykes all contained melt during the period of this lateCarboniferous dextral transtensive deformation. Also, the dykes demonstrate that melt extraction and ascent through the crust had occurred at and below the crustal level exposed; that is, there is a flux of melt through any particular level of exposure in the source. The cumulative frequency distribution of dyke thicknesses suggests that dyke formation may be scale-invariant, which is consistent with the likelihood that dyke spacing and dyke width are controlled by the material properties at a scale larger than the migmatite anisotropy. In S Brittany, spacing of conduits is inferred to have been controlled by the strong intrinsic permeability anisotropy in layers with well-developed stromatic layering and the weak implied extrinsic permeability anisotropy between layers, which leads to variable focusing of melt that requires a larger collection zone for each conduit than would be the case for an isotropic protolith (cf. Eichhubl & Boles 2000). However, overall the limited scale of observations of dyke widths across only two orders of magnitude leaves ambiguity in this result, and scale invariance of melt extraction and ascent remains to be demonstrated convincingly in nature. The switch from ascent to emplacement may be caused by amplification of instabilities (Brown 2001a, b), where instabilities may occur within (e.g. caused by variations in permeability or magma flow rate) or surrounding the ascent column (e.g. caused by variations in the strength or state of stress). Alternatively, ascending magma may encounter a weaker layer in the crust or intersect the brittle-ductile transition zone or some other discontinuity, any of which may enable horizontal magma emplacement and pluton inflation (Roman Berdiel et al. 1995, 1997; Brown & Solar 1998b; Handy

219

et al. 2001). In the case of the DSA, magma

ascent appears to have stalled around the brittle-ductile transition zone and to have been emplaced horizontally approximately along the interface between the infrastructure and the suprastructure. In the DSA, the lower crustal unit is narrow and parallel to the SASZ, which is a major crustal-scale tectonic discontinuity, and leucogranite melt ascent and emplacement was diachronous across the DSA. To the northwest side of the SASZ older plutons are spatially related to the SASZ, which could have been a major backbone element for early melt extraction from anatectic lower crust. During emplacement, leucogranite magma expanded into the Domaine Centre-Armoricain to form plutons (Martelet et al. 2004). Along the SASZ, granites occur as steeply oriented tabular plutons. To the southwest side of the SASZ younger granites occur as thin subhorizontal tabular plutons emplaced along horizons now reactivated as extensional detachments. The system of structures observed in lower crustal migmatites is inferred to have accommodated large-scale buoyancy-driven melt transport to feed Upper Carboniferous two-mica leucogranites. The importance of the two-mica granites in focusing deformation along regional-scale shear zones and detachments has been noted by a number of authors (e.g. Strong & Hanmer 1981; Lagarde et al. 1990; Gapais et al. 1993; Brown & Dallmeyer 1996). The temporal migration of the focus of Carboniferous leucogranite magmatism from earlier ascent and emplacement spatially related to the SASZ (Strong & Hanmer 1981) to later ascent through the migmatites and emplacement at detachments between units (Gapais et al. 1993) is inferred to reflect the change in geodynamics from sinistral transpression to dextral transtension. This migration implies that early melt flow was subhorizontal toward the SASZ. In its present configuration, the SASZ does not extend downward to the Moho, but has been displaced to the NE by thrusting (Bitri et al. 2003; Martelet et al. 2004). However, images of the upper mantle based on seismic tomography show that the SASZ may be traced through the whole lithosphere (Judenherc et al. 2003). Thus, at the time of earlier Carboniferous leucogranite magmatism, it is likely that the SASZ cut through the suprasolidus lower crust and provided a channel by which melt could escape. In contrast, the synchroneity among the second melting event recorded in the DSA lower crustal migmatites, represented by the cordierite-bearing granite-leucosome network,

220

M. BROWN

and the two-mica granites, such as the Ploemeur, Quiberon, Sarzeau and Gu6rande plutons emplaced along extensional detachments at higher crustal levels (Brown & Dallmeyer 1996), suggests that there was melt transfer directly from the anatectic zone, represented by the lower crustal migmatites, via dykes with rare 'mega-dykes' representing the roots or feeders to the sites of emplacement of former plutons now lost to erosion. The inference is that Upper Carboniferous exhumation of the DSA was facilitated by melt trapped at horizons that became reactivated as extensional detachments (Gapais et al. 1993), and segregation, ascent and emplacement of crustally-derived leucogranite magma was intimately linked to regional-scale late orogenic deformation and core complex formation.

Core complex formation

The distribution of melt at depth in an orogen will influence the strain distribution in the overlying subsolidus crust (e.g. Brunet al. 1994). In the DSA, the lower crustal units are exposed in structural culminations, interpreted to be the result of transtension that allowed the middle and lower crustal units to rise up, initiating several dome-like structures (Audren 1987; Tirel et al. 2003) and imposing boudin-like structures at the crustal scale (Gapais et al. 1995). As the weak melt-beating crust of the DSA responded to the regional transtension by rising, so the migmatitic layering and fabric were rotated from subhorizontal in the lower crust to the present subvertical orientation. Brown & Dallmeyer (1996) suggested that the Carboniferous leucogranites and the formation of a core complex in the DSA are the crustal record of slab detachment, possibly triggered by the change from close-to-orthogonal sinistral transpressive to highly oblique dextral transtensive deformation. Although slab detachment may not have occurred (see discussion above), the second melting event that led to melt loss from lower crust and core complex formation may still reflect a change in the deformation regime. The trigger for the second melting episode is inferred to be the change from sinistral transpression to dextral transtension. The reasoning behind this inference is that a change from transpression to transtension that involves reversal of the sense of shear requires a switch among the principal stresses. Such a switch is inferred to induce incipient sites of dilation and boudinage, and in appropriate circumstances the small

decrease in P at spaced intervals along competent layers in extension may initiate a metamorphic reaction with positive d P / d T slope (in this case the biotite dehydration induced melting reaction involving cordierite as a peritectic product; see Fig. 2). This certainly appears to have been what has happened in the DSA migmatites, since the peritectic cordierite is located in interboudin partitions that clearly must have been sites of melting to generate the cordierite, and also sites of melt loss to preserve the cordierite. The production of small amounts of melt occurs over a very small drop in P for this particular reaction, but the effective reduction in density and rheology throughout the volume of suprasolidus orogenic crust implied by the exposed lower crustal migmatite unit in the DSA likely was sufficient to facilitate the initial rise of this unit during the early stage of transtensive deformation. This initiated a feedback relation between lower crustal doming and decompression melting, and between dome amplification and melt extraction and emplacement in developing extensional detachments. The increase in volume associated with the melting reaction in combination with doming in response to the dextral transtensive deformation drives interconnection of the deformation band network, parts of which likely are reactivated structures from the earlier melting event. The dextral shear component of the deformation may have helped to facilitate extraction of the melt generated during decompression. This process likely is terminated once sufficient melt has accumulated at the base of the Belle Ile Group to allow detachment of upper crustal units from the lower crustal unit, which enabled rapid tectonic exhumation and fast cooling of the migmatite source, and terminated reaction and froze any residual melt. Marchildon and Brown (2003) suggested that melt loss at this stage accumulated at shallower crustal levels and was emplaced along the lower contact of the Belle-Ile Group to facilitate extensional detachment of the upper crustal units. This level probably was close to the Upper Carboniferous brittle-ductile transition zone (e.g. Cagnard et al. 2004); the strength peak in the crust at this level probably represents a natural magma trap, particularly when the effects of magma pressure in relation to the regional stress field are considered (Brisbin 1986; Vigneresse et al. 1999), unless the melt begins ascent with a magmastatic head sufficient to break through this barrier. Displacement related to the detachment is recorded by the late-magmatic-early-subsolidus S-C fabrics

SYNERGY BETWEEN MELTING AND DEFORMATION present in the Quiberon, Sarzeau and Gu&ande plutons (Bouchez et al. 1981; Audren 1987; Gapais et al. 1993; Brown & Dallmeyer 1996). In contrast, at the southeast extremity of the DSA around Les Sables d'Olonne, the Upper Carboniferous deformation was accommodated by pervasive extension beneath the brittleductile transition zone and localization and core complex formation did not occur (Cagnard et al. 2004). This difference may be related to degree of melting in the lower crust, since core complex formation appears to require local weak heterogeneities in the lower crust and in the absence of such weak heterogeneities distributed thinning occurs (Brunet al. 1994). Since the pressure of metamorphism in the core complexes around Vannes and St. Nazaire is not significantly different than that at the structural base of the section at Les Sables d'Olonne, this cannot relate to differential thickening during the contractional (transpressive) phase of orogenesis but may instead relate to the bulk composition and/or fertility of the supracrustal lithologies in the lower crust here. Thus, at Les Sables d'Olonne the higher proportion of quartz-rich units in the supracrustal sequence may have limited the degree of melting, which, in turn, means that the lower crust did not provide sufficient fugitive melt to facilitate localization at the brittle-ductile transition zone and core complex formation, or the bulk composition may have been inappropriate for the cordieriteproducing biotite dehydration-driven melting reaction to occur. In summary, the following three features of the late Carboniferous evolution of the DSA support a feedback relation between transtension, a second melting event in the lower crustal unit, leucogranite ascent and emplacement, exhumation and core complex formation: (1) the decompression step in the retrograde P-T path followed by the DSA migmatites (Brown & Dallmeyer 1996; Johnson & Brown 2004); (2) the occurrence of subhorizontal tabular synkinematic plutons of leucogranite marking extensional detachments that separate high-T migmatites of the lower crustal unit from lower-T rocks of the upper crustal units (Gapais et al. 1993); and, (3) concordance of cooling ages among successively lower-T thermochronometers that indicates rapid cooling of lower crustal units and leucogranites during the late Carboniferous (Brown & Dallmeyer 1996). I thank D. Gapais and K. McCaffrey for reviews of an earlier version of this paper, T. Johnson and R. Weinberg for informal reviews, and anonymous and C. Teyssier for fornlal reviews; review comments

221

stimulated significant improvement of the paper. This work has been supported by the NSF (EAR-0003531) and the University of Maryland; Nathalie Marchildon and Barry Reno are thanked for help with preparation of the figures.

References ANTONELLINI, M. A., AYD1N, A. & POLLARD, D. D. 1994. Microstructure of deformation bands in porous sandstones at Arches National Park, Utah. Journal of Structural Geology, 16, 941-959. Au DREN, C. 1987. Evolution structurale de la Bretagne m6ridionale au Pal4ozoique. Socidtd Gdologique et Miniralogique de Bretagne Mdmoire, 31. AUDREN, C. 1990. Evolution tectonique et m4tamorphique de la chaTne varisque en Bretagne m6ridionale. Schweizerische Mineralogische und Petrographische Mittelunge, 70, 17-34. AUDREN, C. & PLAINE, J. 1986. Notice explicative de la feuille Belle-Ile-en-Mer et Iles Houat et Hoedic (Carte G6ologique de la France 1:50,000). Orldans: Editions du B.R.G.M. AtJDREN, C. & TRmOULET, C. 1993. P-T-t-deformation paths recorded by Kinzigites during diapirism in the western Variscan Belt (Golfe du Morbihan, southern Brittany, France). Journal of Metamorphic Geology, 11, 337-357. AUTRAN, A. & COGNt~, J. 1980. La zone interne de l'orog~ne varisque dans l'Ouest de la France et sa place dans le d~veloppement de la chaine hercynienne. Mdmoire BRGM, 108, 90-111. BALLI~VRE, M., PITRA, P. 8`: BOHN, M. 2003. Lawsonite growth in the epidote blueschist from the Ile de Groix (Armorican Massif, France): A potential geobarometer. Journal of Metamorphic Geology, 21, 723-735. BARRAUD, J., GARDIEN, V., ALLEMAND, P. & GRANDJEAN, P. 2004. Analogue models of meltflow networks in folding migmatites. Journal of Structural Geology, 26, 307-324. BERNARD-GRIFFITHS,J., PEUCAT, J.-J., SHEPPARD, S. & VIDAL, P. 1985. Petrogenesis of Hercynian leucogranites from the southern Armorican Massif; contribution of REE and isotopic (St, Nd, Pb and O) geochemical data to the study of source rock characteristics and ages. Earth and Planetary Science Letters, 74, 235-250. BERNARD-GRIFFITHS, J., CARPENTER, M. S. N., PEUCAT, J.-J. 8,: JAHN, B. M. 1986. Geochemical and isotopic characteristics of blueschist facies rocks from the Ile de Groix, Armorican Massif (Northwest France). Lithos, 19, 235-253. BERTHE, D., CHOUKROUNE,P. & JI~GOUZO, 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. BITRI, A., BALLI~VRE, M. et al. 2003. Imagerie sismique de la zone de collision hercynienne dans le sud-est du massif Armorican (Projet Armor 2/ G6ofrance 3D). Comptes Rendus Geosciences, 335, 969-979.

222

M. BROWN

BONS, P. D. 1999. Apparent extensional structures due to volume loss. Proceedings of the Estonian Academy of Sciences, Geology, 48, 3-14. BONS, P. D., DOUGHERTY-PAGE,J. & ELBURG, M. A. 2001. Stepwise accumulation and ascent of magmas. Journal of Metamorphic Geology, 19, 627-633. BONS, P. D., ELBURG, M. A., VAN MILLIGEN, B. P., SOESOO, A. & PASSCHIER, C. W. 2002. A numerical model for self-organized criticality in melt accumulation and ascent in partially molten rocks. In: PREISS, V. P. (ed) Geoscience 2002: Expanding Horizons. Abstracts of the 16th Australian Geological Convention, Adelaide Convention Center, Adelaide, SA, Australia. July 1-5, 2002, No. 67, 206. BONS, P. D., ARNOLD, J., ELBURG, M. A., CALDER, J. & SOESOO, A. 2004. Melt accumulation from partially molten rocks. Lithos, 78, 25-42. BOSSE, V., FI~RAUD, G., BAELEVRE, M. & PEUCaT, J.-J. 2000. Petrological and geochronological constraints in the Ile de Groix blueschist (Armorican Massif, France). Basement Tectonics, 15, 63-66. BOSSE, V., BALLI~VRE,M. & VIDAL, O. 2002. Ductile thrusting recorded by the garnet isograd from the blueschist-facies metapelites of the Ile de Groix, Armorican Massif, France. Journal of Petrology, 43, 485-510. BOSSlERE, G. 1988. Evolutions chimico-min~ralogiques du grenat et de la muscovite au voisinage de l'isograde biotite-staurotide dans un mGtamorphisme prograde de type barrovien: un exemple en Vendee littorale (Massif Armoricain). Comptes Rendus de l'Acaddmie des Sciences, Paris, Sdrie //, 306, 135-140. BOUCHEZ, J.-L., GU1LLET, P. & CHEVALIER, F. 1981. Structures d'Gcoulement li~es ~ la mise en place du granite de GuGrande (Loire-Atlantique, France). Bulletin de Socidtd Gdologique de France, 7, 387-399. BmSB1N, W. C. 1986. Mechanics of pegmatite intrusion. American Mineralogist, 71, 644-651. BROWN, M. 1983. The petrogenesis of some migmatites from the Presqu'i'le de Rhuys, southern Brittany, France. In: ATHERTON, M. P. & GRIBBLE, C. D. (eds) Migmatites, Melting and Metamorphism. Shiva Publishing, Nantwich, 174- 2OO. BROWN, M. 1994. The generation, segregation, ascent and emplacement of granite magma: the migmatiteto-crustally-derived granite connection in thickened orogens. Earth Science Reviews, 36, 83-130. BROWN, M. 2001a. Crustal melting and granite magmatism: key issues. Physics and Chemistry of the Earth (A), 26, 201-212. BROWN, M. 200lb. Orogeny, migmatites and leucogranites: a review. Proceedings of the Indian Academy of Sciences (Earth and Planetar)' Science), 110, 313-336. BROWN, M. 2004. The mechanism of melt extraction from lower continental crust of orogens. Transactions of the Royal SocieO' of Edinburgh, Earth Sciences, 95, 35-48.

BROWN, M. & DALLMEYER,R. D. 1996. Rapid Variscan exhumation and the role of magma in core complex formation: southern Brittany metamorphic belt, France. Journal of Metamorphic Geology, 14, 361-379. BROWN, M. & O'BRIEN, P. J. 1997. Evolution of metamorphic belts: a changing view. In: XIANGHN, Q., ZHENDONG, Y. & HALL, H. C. (eds) Precambrian Geology and Metamorphic Petrology. Proceedings of the 30th International Geological Congress, VSP, Zeist, The Netherlands, 17, 217-231. BROWN, M. & RUSHMER, T. 1997. The role of deformation in the movement of granite melt: views from the laboratory and the field. In: HOLNESS, M. B. (ed) Deformation-Enhanced Fluid Transport in the Earth's Crust and Mantle. The Mineralogical Society Series: 8. Chapman and Hall, London, 111-144, BROWN, M. & SOLAR, G. S. 1998a. Shear zone systems and melts: Feedback relations and selforganization in orogenic belts. Journal of Structural Geology, 20, 211-227. BROWN, M. & SOLAR, G. S. 1998b. Granite ascent and emplacement during contractional deformation in convergent orogens. Journal of Structural Geology, 20, 1365-1393. BROWN, M. & SOLAR, G. S. 1999. The mechanism of ascent and emplacement of granite magma during transpression: a syntectonic granite paradigm. Tectonophysics, 312, 1-33. BROWN, M., AVERKIN, Y. A., MCLELLAN, E. L. & SAWYER, E. W. 1995. Melt segregation in migmatites. Journal of Geophysical Research - Solid Earth, 100, 15 655-15 679. BROWN, M. A., BROWN, M., CARLSON, W. D. & DENISON, C. 1999. Topology of syntectonic melt flow networks in the deep crust: inferences from three-dimensional images of leucosome geometry in migmatites. American Mineralogist, 84, 1793-1818. BRUN, J.-P. & BURG, J.-P. 1982. Combined thrusting and wrenching in the Ibero-Armorican arc: a corner effect during continental collision. Earth and Planetar)' Science Letters, 61, 319-332. BRUN, J.-I., SOKOUTIS, D. & VANDENDRIESSCHE,J. 1994. Analog modeling of detachment fault systems and core complexes. Geology, 22, 319-322. BURG, J.-P., BALI~, P., BRUN, J.-P. & GIRARDEAU, J. 1987. Stretching lineation and transport direction in the Ibero-Armorican arc during the SiluroDevonian collision. Geodynamica Acta, 1, 71-87. BURG, J.-P., VAN DEN DRIESSCHE, J. & BRUN, J.-P. 1994. Syn- to post-thickening extension in the Variscan Belt of western Europe: modes and structural consequences. Gdologie de la France, 1994, 33-51. CAGNARD, F., GAPAIS, D., BRUN, J.-P., GUMIAUX, C. & VAN DEN DRIESSCHE, J. 2004. Late pervasive crustal-scale extension in the south Armorica Hercynian belt (Vendee, France). Journal of Structural Geology, 26, 435-440. CARPENA, J., CHAILLOU, D., CHAMBAUDET, A. & POUPEAU, G. 1979. Fission track geochronology

SYNERGY BETWEEN MELTING AND DEFORMATION of the Hercynian Platform in France. In: Report of the lOth Annual Conference on Solid State Nuclear Track Detectors. Space Biophysics of the Parliament of the Council of Europe, Lyon. CARTIER, C., FAURE, M. & LARDEUX, H. 2002. Discussion on 'a new model for the Hercynian orogen of Gondwana France and Iberia' by D. Shelley & G. Bossi~re, Journal of Structural Geology 2000, 22, 757-776. Journal of Structural Geology, 24, 1659-1663. CLEMENS, J. D. 1998. Observations on the origins and ascent mechanisms of granitic magmas. Journal of the Geological Society, 155, 843-851. COGNE, J. 1960. Schistes cristallins et granites en Bretagne mGridionale: le domaine de l'Anticlinal de Cornouaille. Mimoire du Service de la Carte Giologique de France, 1, 382. DEWEY, J. & BURKE, K. C. 1973. Tibetan, Variscan and Precamhrian basement reactivation: products of continental collision. Journal of Geology, 81, 683 -692. DIAS, R. & RIBEIRO, A. 1995. The Ibero-Armorican arc: A collision effect against an irregular continent? Tectonophysics, 246, 113-128. DJRO, S. C., TR1BOULET, C. & AUDREN, C. 1989. Les chemins pression- temperature- temps-dGformation-espace ( P - T - t - d - e ) dans les micaschistes associGs aux schistes bleus de l'~le de Groix, Bretagne M~ridionale, France. Schweizerische Mineralogische und Petrographische Mitteilungen, 69, 73-90. Du BERNARD, X., E1CHHUBL, P. & AYDIN, A. 2002. Dilation bands: a new form of localized failure in granular media. Geophysical Research Letters, 29, DOI: 10.1029/2002GL015966. EICHHUBL, P. & AYDIN, A. 2003. Ductile openingmode fracture by pore growth and coalescence during combustion alteration of siliceous mudstone. Journal of Structural Geology, 25, 121 - 134. EICHHUBL, P. & BOLES, J. R. 2000. Focused fluid flow along faults in the Monterey Formation, Coastal California. Geological Society of American Bulletin, 112, 1667-1679. EICHHUBL, P., AYDIN, A. & LORE, J. 2001. Openingmode fracture in siliceous mudstone at high homologous temperature - effective of surface forces. Geophysical Research Letters, 28, 1299-1302. FAURE, M. & PONS, J. 1991. Crustal thinning recorded by the shape of the Namurian-Westphalian leucogranite in the Variscan Belt of the Northwest Massif Central, France. Geology, 7, 730-733. FAURE, M., LELOIX, C. & ROIG, J.-Y. 1997. L'E~volution polycyclique de la chaine hercynienne. Bulletin de la Sociiti Gdologique de France, Huitikme Sirie, 168, 695-705. FYFE, W. S. 1973. The granulite facies, partial melting and the Archaean crust. Philosophical Transactions of the Royal Society of London, 273A, 457-461. GAPA1S, D. 1989. Shear structures within deformed granites - mechanical and thermal indicators. Geology, 17 (12), 1144-1147. GAPAIS, D., LAGARDE, J.-L., LE CORRE, C., AUDREN, C., JIEGOUZO, P., CASAS SAINZ, A. &

223

VAN DEN DRIESSCHE, J. 1993. La zone de cisaillement de Quiberon: tGmoin d'extension de la cha~ne varisque en Bretagne mGridionale au CarbonifGre. Comptes Rendus de l'Academie Sciences Paris, Serie 2, 316, 1123-1129. GAPAIS, D., VAN DEN DRIESSCHE, J., BRUN, J.-P. & RICHER, C. 1995. Pervasive ductile spreading and boudinage in extending orogens; evidence from the French Variscan Belt. TERRA Abstracts, 7, 121. GEOFFROY, L. 1993. Tectonique tardi-varisque en failles normales ductiles en VendGe Littorale, Massif Armoricain. Cornptes Rendus des S~ances de l'Acaddmie des Sciences, 317, 1237-1243. GERDES, A., WORNER, G. & HENK, A. 2000. Postcollisional granite generation and H T - L P metamorphism by radiogenic heating: the Variscan South Bohemian Batholith. Journal of the Geological Society, London, 157, 577-587. GEEASON, G. C., BRUCE, V. & GREEN, H. W. 1999. Experimental investigation of melt topology in partially melted quartzo-feldspathic aggregates under hydrostatic and non-hydrostatic stress. Journal of Metamorphic Geology, 17, 705-722. GODARD, G. 1988. Petrology of some eclogites in the Hercynides: The eclogites from the southern Armorican Massif, France. In: SMITH, D. C. (ed) Eclogites and Eclogite-Facies Rocks. Elsevier, Amsterdam, 451-519. GODARD, G. 2001. Les Essarts eclogite-bearing metamorphic complex (VendGe, Southern Armorican Massif, France). Gdologie de la France, 2001 (1-2), 19-51 & 2 maps. GouJou, J. C. 1992. Analyse pGtro-structurale dans un avant-pays m~tamorphique: influence du plutonisme tardi-orogGnique varisque sur l'encaissant 6pi fi mGsozonal de VendGe. Documents du BRGM, 216. GRUJIC, D. & MANCKTELOW,N. S. 1998. Melt-bearing shear zones: analogue experiments and comparison with examples from southern Madagascar. Journal of Structural Geology, 20, 673-680. GUERNINA, S. & SAWYER, E. W. 2003. Large-scale melt-depletion in granulite terranes: An example from the Archean Ashuanipi Subprovince of Quebec. Journal of Metamorphic Geology, 21, 181-201. GUIRAUD, M., BURG, J.-P. & POWELL, R. 1987. Evidence for a Variscan suture zone in the VendGe, France: a petrological study of blueschist facies rocks from Bois de Cen& Journal of Metamorphic Geology, 5, 225-237. GUMIAUX, G., JUDENHERC, S., BRUN, J.-P., GAPAIS, D., GRANET, M. & POUPINET, G. 2004. Restoration of lithosphere-scale wrenching from integrated structural and tomographic data (Hercynian belt of western France). Geology, 32, 333 -336. HAND, M. & DIRKS, P. H. G. M. 1992. The influence of deformation on the formation of axial-planar leucosomes and the segregation of small melt bodies within the migmatitic Napperby Gneiss, central Australia. Journal of Structural Geology, 14, 591-604.

224

M. BROWN

HANDY, M. R., MULCH, A., ROSENAU, N. & ROSENBERG, C. L. 2001. The role of fault zones and melts as agents of weakening, hardening and differentiation of the continental crust: a synthesis. In: HOLDSWORTH, R. E., MAGLOUGHLIN, J., KNIPE, R. J., STRACHAN, R. A. & SEARLE, R. C. (eds). The Nature and Tectonic Significance of Fault Zone Weakening. Geological Society Special Publications, 186, 305-332. HANMER, S. 1997. Shear zone reactivation at granulite facies: The importance of plutons in the localization of viscous flow. Journal of the Geological Society, 154, 111 - 116. HILL, J., BALDWIN, S. L. &LlSTER, G. S. 1995. Magmatism as an essential driving force for formation of active metamorphic core complexes in eastern Papua New Guinea. Journal of Geophysical Research, 100, 10 441-10 451. HODC,ES, K. V. 2000. Tectonics of the Himalayas and southern Tibet from two prospectives. Geological Society of America Bulletin, 112, 324-350. HOLmSTER, L. S. 1993. The role of melt in the uplift and exhumation of orogenic belts. Chemical Geology, 108, 31-48. HOLYOKE, III, C. W. & RUSHMER, T. 2002. An experimental study of grain scale melt segregation mechanisms in two common crustal rock types. Journal of Metamorphic Geology, 20, 493-512. Jr~GOUZO, P. 1980. The South Armorican shear zone. Journal o[ Structural Geology, 2, 39-47. JEGOUZO, P. & ROSSELLO, E. A. 1988. La branch6 nord du cisaillement Sud-Armoricain (France): un essai d'6valuation du d6placement par l'analyse des mylonites. Comptes Rendus de l'Academie Sciences Paris, Serie 2,307, 1825-1831. JOHNSON, T. & BROWN, M. 2004. Quantitative constraints on metamorphism in the Variscides of southern Brittany - a complementary pseudosection approach. Journal of Petrology, 45, 1237-1259. JONES, K. A. 1988. The metamorphic petrology of the southern Brittany Migmatite Belt, France. PhD dissertation, Kingston Polytechnic University. JONES, K. A. 1991. Paleozoic continental margin tectonics of southern Armorica. Journal of the Geological Society, London, 148, 55-64. JONES, K. A. & BROWN, M. 1989. The metamorphic evolution of the southern Brittany Migmatite Belt. In: DALY, J. S., CLIFF, R. A. & YARDLEY, B. W. D. (eds) Evolution of Metamorphic Belts. Geological Society Special Publications, 43, 501-505. JONES, K. A. & BROWN, M. 1990. High-temperature 'clockwise' P-T paths and melting in the development of regional migmatites: an example from southern Brittany, France. Journal of Metamorphic Geology, 8, 551-578. JUDENHERC, S., GRANET, M., BRUN, J.-P., POUP1NET, G., PLOMEROVA, J., MOCQUET, A. & ACHAUER, U. 2002. Images of lithospheric heterogeneities in the Armorican segment of the Hercynian Range in France. Tectonophysics, 358, 121-143.

JUDENHERC, S., GRANET, M., BRUN, J.-P. & POUPINET, G. 2003. The Hercynian collision in the Armorican Massif; evidence of different lithospheric domains inferred from seismic tomography and anisotropy. Bulletin de la Socidtd Gdologique de France, 174, 45-57. LAGARDE, J.-L., BRUN, J.-P. & GAPAIS, D. 1990. Formation des plutons granitiques par injections et expansion lat&ale dans leur site de raise en place: une alternative au diapirisme en domaine ~pizonal. Comptes Rendus des Sdances de l'Acadimie des Sciences, 310, 1109-1114. LE HEI3EL, F. 2002, Ddformation continentale et histoire des fluides au cours d'un cycle subduction, exhumation, extension. Exemple des porphyroi'des sud-armoricains. Unpublished PhD thesis, University of Rennes I. LE HI2BEL, F., GAPAIS, D., FOURCADE, S. &; CAPDEVILA, R. 2002a. Fluid-assisted large strains in a crustal-scale ddcollement (Hercynian belt of south Brittany, France). Geological Society Special Publications, 200, 85-101. LE HI~BEL,F., VIDAL, 0., KIENAST,J. R. ~; GAPAIS, D. 2002b. Les 'Porphyroides' de Bretagne m6ridionale: une unit6 de HP-BT dans la chaine hercynienne. Comptes Rendus G&)science, 334, 205-211. MANDAL, N., MISRA, S. & SAMANTA, S. K. 2004. Role of weak flaws in nucleation of shear zones: an experimental and theoretical study. Journal of Structural Geology, 26, 1391 - 1400. MARCHILDON, N. & BROWN, M. 2001. Melt segregation in late-tectonic anatectic migmatites: an example from the Onawa contact aureole, Maine, U.S.A. Physics and Chemistry of the Earth, 26, 225 -229. MARCHILDON, N. &; BROWN, M. 2002. Grain-scale melt distribution in two contact aureole rocks: implications for controls on melt localization and deformation. Journal of Metamorphic Geology, 20, 381-396. MARCHILDON, N. & BROWN, M. 2003. Spatial distribution of melt-bearing structures in anatectic rocks from southern Brittany, France: implications for melt-transfer at grain-scale to orogen-scale. Tectonophysics, 364, 215-235. MARTELET, G., CALCAGNO, P., GUMIAUX, C., TRUFFERT, C., BITRI, A., GAPAIS, D. & BRUN, J.-P. 2004. Integrated 3D geophysical and geological modelling of the Hercynian Suture Zone in the Champtoceaux area (south Brittany, France). Tectonophysics, 382, 117-128. MATTE, P. 1986. Tectonics and plate tectonic model for the Variscan belt of Europe. Tectonophysics, 126, 329-374. MATTE, P. 1991. Accretionary history and crustal evolution of the Variscan belt in Western Europe. Tectonophysics, 196, 309-337. MATTE, P. 2001. The Variscan collage and orogeny (480-290 Ma) and the definition of the Armorica microplate: tectonic approach. Terra Nova, 13, 122-128. MCKENZIE, D. P & JACKSON, J. 2002. Conditions for flow in the continental crust. Tectonics, 21, 1055, doi: 10.1029/2002TC001394.

SYNERGY BETWEEN MELTING AND DEFORMATION MOLLEMA, P. N. & ANTONELLINI,M. A. 1996. Compaction bands; a structural analog for anti-mode I cracks in Aeolian sandstone. Tectonophysics, 267, 209-228. OLIVER, N. H. S. & BARR, T. D. 1997. The geometry and evolution of magma pathways through migmatites of the Halls Creek Orogen, Western Australia. Mineralogical Magazine, 61, 3-14. PARSONS, T. & THOMPSON, G. A. 1993. Does magmatism influence low-angle normal faulting? Geology, 21, 247-250. PETFORD, N. & KOENDERS,M. A. 1998. Self-organization and fracture connectivity in rapidly heated continental crust. Journal of Structural Geology, 20, 1425-1434. PEUCAT, J.-J. 1983. Grochronologie des roches mrtamorphiques (Rb-Sr et U-Pb). Mdmoire de la Socidtd Gdologique et Mindralogique de Bretagne, 28, 1-158. PIN, C. & PEUCAT, J.-J. 1986. Ages des ~pisodes de m6tamorphisme pa16ozoi'que dans le Massif central et le Massif Armoficain. Bulletin de Socidtd Gdologique de France, 8, 461-469. QUINQU1S, H. & CHOUKROUNE, P. 1981. Les schistes bleus de File de Groix dans la cha~ne hercynienne: implications cinrmatiques. Bulletin de la Socidtd Gdologique de France, 23, 409-418. REGENAUER-LIEB, K. 1999. Dilatant plasticity applied to Alpine collision: Ductile void growth in the intraplate area beneath the Eifel volcanic field. Journal of Geodynamics, 27, 1-21. REY, P., BURG, J. P. & CASEY, M. 1997. The Scandinavian Caledonides and their relationship to the Variscan belt. In: BURG, J. P. & FORD, M. (eds) Orogeny Through Time. Special Publications Geological Society, 121, 179-200. ROMAN BERDIEL, T., GAPAIS, D. & BRUN, J.-P. 1995. Analog models of laccolith formation. Journal of Structural Geology, 17, 1337-1346. ROMAN BERDIEL, T., GAPAIS, D. & BRUN, J.-P. 1997. Granite intrusion along strike-slip zones in experiment and nature. American Journal of Science, 297, 651-678. RUBATTO, D., WILLIAMS, I. S. & BUICK, I. S. 2001. Zircon and monazite response to prograde metamorphism in the Reynolds Range, central Australia. Contributions to Mineralogy and Petrology, 140, 458-468. RUTTER, E. H. 1997. The influence of deformation on the extraction of crustal melts: a consideration of the role of melt-assisted granular flow. In: HOLNESS, M. B. (ed) Deformation-Enhanced Fluid Transport in the Earth's Crust and Mantle. Mineralogical Society Series, 8, 82-110. SAWYER, E. W. 2001. Melt segregation in the continental crust: distribution and movement of melt in anatectic rocks. Journal of Metamorphic Geology, 19, 291-309. SCHULTZ, B., TRIBOULET, C., AUDREN, C., PFEIFER, H.-R. & GILG, A. 2001. Two-stage prograde and retrograde Variscan metamorphism of glaucophane-eclogites, blueschists and greenschists from the Ile de Groix (Brittany, France). International Journal of Earth Sciences, 90, 871-889.

225

SCHULTZ, B., AUDREN, C. & TRIBOULET, C. 2002. Oxygen isotope record of fluid-rock-SiO2 interaction during Variscan progressive deformation and quartz veining in the meta-volcano sediments of Belle-Ile (Southern Brittany). Journal of Structural Geology, 24, 1281-1297. SHELLEY, O. & BOSSII~RE,G. 1999. Ile de Groix: retrogression and structural developments in an extensional rrgime. Journal of Structural Geology, 21, 1441-1455. SHELLEY, D. & BOSSII~RE, G. 2000. A new model for the Hercynian Orogen of Gondwanan France and Iberia, Journal of Structural Geology, 22, 757 -776. SmSON, R. H. 1996. Structural permeability of fluiddriven fault - fracture meshes. Journal of Structural Geology, 18, 1031 - 1042. SIBSON, R. H. & SCOTT, J. 1998. Stress/fault controls on the containment and release of over pressured fluids: examples from gold-quartz vein systems in Juneau, Alaska; Victoria, Australia and Otago, New Zealand. Ore Geology Reviews, 13, 293-306. SOLAR, G. S. & BROWN, M. 2001. Petrogenesis of migmatites in Maine, USA: possible source of peraluminous leucogranite in plutons? Journal of Petrology, 42, 789-823. STRONG, D. F. & HANMER, S. K. 1981. The leucogranites of southern Brittany; origin by faulting, frictional heating, fluid flux and fractional melting. Canadian Mineralogist, 19, 163-176. TEYSSIER, C. & WHITNEY, D. 2002. Gneiss domes and orogeny. Geology, 30, 1139-1142. TIREL, C., BRUN, J.-P. & BUROV, E. 2003. Thermomechanical modelling of extensional gneiss domes. Deformation Mechanisms, Rheology and Tectonics. G~osciences Rennes, St. Malo, France, April 14-16, 2003, Abstract Volume, 160. TRmOULET, C. 1974. Les glaucophanites et roches associres de File de Groix (Morbihan, France): 6tude min~ralogique et prtrogrn6tique. Contributions to Mineralogy and Petrology, 45, 65-90. TRIBOULET, C. 1991. Etude gdothermo-baromrtrique compar6 des schistes bleus palrozo'iques de l'Ouest de la France (Ile de Groix, Bretagne Mrridionale et Bois de Cent, Vendre). Comptes Rendus des Sdances de l'Acaddmie des Sciences, 312, 1163-1168. TRIBOULET, C. & AUDREN, C. 1985. Continuous reactions between biotite, garnet, staurolite, kyanitesillimanite-andalusite and P-T-time-deformation path in micaschists from the estuary of the river Vilaine, South Brittany, France. Journal of Metamorphic Geology, 3, 91-105. TRIBOULET, C. 8,: AUDREN, C. 1988. Controls on P-T-t deformation path from amphiboles during progressive metamorphism with basic rocks (estuary of the river Vilaine, South Brittany, France). Journal of Metamorphic Geology, 6, 117-133. TRUFFERT, C., GUMIAUX, C. ET AL. 2001. Lev~ g~ophysique arroport6 dans le Sud-Est du Massif armoricain (programme GroFrance3D Armor2). Magn~tisme et Radiom~trie spectrale. Comptes

226

M. BROWN

Rendus des Sdances de l'Acaddmie des Sciences, 333, 263-270. VAN DE ZEDDE, D. M. A. & WORTEL, M. J. R. 2001. Shallow slab detachment as a transient source of heat at midlithospheric depths. Tectonics, 20, 86-882. VAN DER MOLEN, I. & PATERSON, M. S. 1979. Experimental deformation of partially-melted granite. Contributions to Mineralogy and Petrology, 70, 299-318. VANDERHAEGHE, 0. & TEYSSIER, C. 2001. Partial melting and flow of orogens. Tectonophysics, 342, 451-472.

VAtJCHEZ, A., MAILLET, D. & SOUGY, J. 1987. Strain and deformation mechanisms in the Variscan nappes of Vend6e, South Brittany, France. Journal of Structural Geology, 9, 31-40. VIDAL, P. 1973. Premieres donn~es g6ochronologiques sur les granites hercyniens du Sud du Massif Armoricain. Bulletin de Soci~td Giologique de France, 7, 239-245. VIGNERESSE, J.-L., TIKOFF, B. & AMI~GLIO, L. 1999. Modification of the regional stress field by magma intrusion and formation of tabular granitic plutons. Tectonophysics, 302, 203-224.

A critical assessment of the tectono-thermal memory of rocks and definition of tectono-metamorphic units: evidence from fabric and degree of metamorphic transformations M. IOLE SPALLA 1'2, MICHELE ZUCALI 1, SILVIA DI PAOLA 1 & GUIDO GOSSO 1'2

1Dipartimento di Scienze della Terra 'A. Desio', Universitdt di Milano, Via Mangiagalli 34, 20133 Milano, Italy (e-mails: [email protected], [email protected], [email protected], [email protected]) 2C.N.R. - IDPA, Via Mangiagalli 34, 1-20133 Milano, Italy Abstract: A correlation procedure of scattered tectonic and metamorphic imprints in the reactivated crust is elaborated from recent analytical work in three Alpine metamorphic complexes. It consists of: interpretation of the time-sequence of tectonic fabrics and test of their kinematic coherence; determination of paragenetic compatibility among the mineralogical support of mesoscopic fabrics; cross-validation of mineral transformation overprints; construction of P-T-d-t paths using a time-sequence of parageneses. The representation of structural and metamorphic information conveys the full tectono-metamorphic history on maps displaying combined tectonic and metamorphic effects. Shape and size definition of metamorphic units, now individuated mainly using their lithological homogeneity and dominant metamorphic imprint, is improved. The analysis of interactions between fabric and metamorphic imprint distributions, proposed in three Alpine examples, shows that the dominant metamorphic imprint does not coincide with Tmax-Prmaxof each inferred P-T-d-t loop; the dominant metamorphic imprint is that given by the mineralogical support of the most pervasive fabric. Different metamorphic imprints may dominate in adjacent areas of a single tectono-metamorphic unit (TMU), or equivalent metamorphic imprints may dominate in different TMUs. Therefore, lithostratigraphic setting and dominant metamorphic imprint are inefficient to contour TMUs in terrains with polyphase deformation and metamorphism, without considering multiscale heterogeneity of superposed synmetamorphic fabrics.

Coupling and decoupling of crustal slices, with their associated size variation, work in competition within subduction-collision zones to form the tectonic units of metamorphic belts. Since their structural and metamorphic evolutions are tracers of their transit throughout different levels of the lithosphere, the definition of their contours may be improved to individualize the volumes carrying distinct structural and metamorphic histories (tectono-metamorphic unit = TMU). Size definition of a TMU, and careful reconstruction of its tectono-thermal evolution, are critical to infer geological processes as (i) tectonic erosion or accretion at the trench margins, (ii) continental collision or deep subduction of continental crust (ablative subduction), (iii) exhumation velocity variation and its influence on rapid and effective meta-stabilization of HP- and UHP-LT assemblages (e.g. England & Thompson 1984; Lallemand 1999; Tao & O'Connel 1992; Cloos 1993; Ernst 2001).

A reliable assessment of the effect of these different processes requires a careful examination of the tectono-metamorphic memory of units traditionally distinguished in tectonically reactivated crust and mountain belts; therefore an analytical procedure is presented, aiming to individuate portions of crust that underwent an internally coherent tectonic and metamorphic evolution. At present, in spite of remarkable progress of analytical techniques in structural geology and petrology (e.g. Passchier et al. 1990; Brown 2001), the recognition of specific P-T re-equilibration stages of successive groups of structures remains a problem because the extent, degree and timing of metamorphic re-equilibrations and fabric changes are often inadequately known to delimit TMUs. Even in relatively well-explored belts, as the European Alps, metamorphic complexes have been contoured on the basis of their lithological homogeneity

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 227-247. 0305-8719/05/$15.00

9 The Geological Society of London 2005.

228

M. IOLE SPALLA ET AL.

(suggesting for example continental or oceanic affinities) or of their dominant metamorphic imprint. On the other hand, recent pressure (P)-temperature (7) paths reconstructions reveal different sequences of metamorphic overprints within a single nappe (or complex; e.g. Pognante 1991; Spalla et al. 1996). In this light, rocks recording the maximum number of re-equilibration steps (i.e. rocks with long memory) are fundamental to reconstruct exhaustively the structural and metamorphic evolution. Fortunately, deformation and metamorphic re-equilibration are heterogeneous processes and rocks may retain a valuable memory that permits the preservation of relics, even at very high T. On the other hand, these heterogeneities make spatial and chronological correlations rather difficult. Consequently, setting up a correlation method that reconstructs well-constrained successions of structural and metamorphic re-equilibration stages, hidden as scanty traces in the tectonothermal memory, is fundamental to separate crustal volumes that underwent significantly different deformation and metamorphic histories. T is currently accepted as the main driving force of reaction kinetics, as suggested by the Arrhenius exponential relation, and consequently the dominant metamorphic imprint of a metamorphic unit is admitted to have been imposed when rocks reached conditions of Tmax-PTmax. In these models (e.g. England & Richardson 1977; Spear et al. 1984) the different T,.... PTmax conditions, recorded in adjacent crustal volumes showing different dominant metamorphic imprints, are not aligned along the same geotherm, but locate on the P - T space the 'metamorphic field gradient'. Tma• culminations of each P - T loop are therefore diachronous (implying different exhumation times) and are considered useful to identify different metamorphic units and to assess their size. In this view, the catalysing effect of grainscale deformation on metamorphic reaction rates is completely disregarded, though the effect of strain on metamorphic recrystallization be well documented at several scales, in many orogenic belts and in various metamorphic environments (e.g. Lardeaux et al. 1982; Mcrk 1985; Pognante 1985; Koons 1986; Austrheim 1990; Rubie 1990; Spalla 1993). For this reason, we take into account the role of fabric evolution on reaction progress to individuate rock volumes characterized by a potentially 'long tectono-thermal memory' that may change throughout adjacent volumes as a function of deformation heterogeneity. Therefore, we will compare the pressure-temperature-relative time of deformation paths (P-T-d-t paths of Spalla

1993; Johnson & Vernon 1995; Passchier & Trouw 1996), reconstructed in three examples from the Western and Central Austroalpine and Central Southalpine domains of the Alps. P-T-d-t paths were inferred using P - T estimates obtained from nfineral assemblages marking superposed and kinematically incompatible tectonic fabrics. They were represented on P - T diagrams by chronologically ordered re-equilibration steps corresponding to D1, D2 . . . . Dn deformation stages. The three examples were chosen in domains that underwent polyphase or polyphase and polycyclic tectono-metamorphic evolution during pre-Alpine and Alpine times, resulting from different geodynamic settings (subduction and continental collision or lithospheric extension) as suggested by strongly contrasted thermal regimes, from HP-LT to HT-LP. We qualitatively evaluate, on maps of structural versus metamorphic re-equilibration, the interaction between distribution of highly evolved fabrics and pervasive growth of metamorphic minerals, to point out the dependence of the degree of metamorphic transformation on deformation gradients. Such a mapping technique improves the structural mapping by Johnson & Duncan (1992) or Connors & Lister (1995) and enables speculation on the physical significance of the metamorphic field gradient and on its use for the definition of TMU boundaries. In addition, TMU boundaries are compared with those of units drawn from lithostratigraphic criteria.

Correlation tools The identification of a TMU requires individuating rock volumes with coherent tectonometamorphic memory and this needs an efficient correlation method. Correlation of structural and metamorphic re-equilibration stages between adjacent rock volumes, in terrains that experienced complex structural and metamorphic evolutions, needs a joint use of different tools: correlation of superposed fabric elements on a foliation trajectory map; microstructural analysis to infer stable mineral assemblages marking superposed fabrics; absolute age data (Turner & Weiss 1963; Park 1969; Hobbs et al. 1976; Van Roermund et al. 1979; Williams 1985; Passchier et al. 1990; Johnson & Duncan 1992; Johnson & Vernon 1995; Spalla et al. 2000; Di Vincenzo & Palmeri 2001). Deformation heterogeneity makes structural correlations difficult, if exclusively based on classical geometric criteria, because it generates strong variations in deformation style across

T E C T O N O - T H E R M A L M E M O R Y OF ROCKS

adjacent rock volumes (e.g. from undeformed to mylonitic textures). The analytical procedure (Fig. 1), adopted to improve correlations, comprises the following (1)

construction of a map of the total deformation field, in which configuration of lithostratigraphy and all the fabric elements is reported; test of kinematic compatibilities between all structures and their superposition sequence, to separate successive deformation imprints; recognition of mineral compatibilities (assemblages) among the mineralogical support of successive planar and/or linear fabrics; cross-control of superposed mineral transformations in the superposed fabric sequences of adjacent heterogeneously deformed volumes.

(2)

(3)

(4)

In accordance, all the structural-petrographic maps presented in this contribution contain

7

. . . . . . . . . . . . . . . . . . .

.

229

grids of superposed foliation trajectories and indicate the associated metamorphic conditions and the fabric gradients. Foliation trajectory symbols represent interpolations of seams of foliation traces containing intbrmation on the relative chronology of structural imprints (Williams 1971; Myers 1978; Gosso et al. 1983; Johnson & Duncan 1992; Connors & Lister 1995). To estimate the degree of fabric evolution, the degree of grain-scale reorganization of the prevailing planar fabric may be regionally evaluated, either with reference to six stages of the decrenulation model (Fig. 2; Bell & Rubenach 1983: Passchier & Trouw 1996), or by differentiating areas with undeformed, normally foliated or mylonitic fabrics. In the three examples, both pre-Alpine and Alpine tectonic fabrics reach variable degrees of evolution, up to stage 5 - 6 (Bell & Rubenach 1983) in metapelites or from undeformed to mylonitic in metaintrusives. The area extent of a new differentiated foliation (evolutionary stages 4 to 6) was evaluated for each successive group of structures, together with the modal amount of associated metamorphic

t

. . . . . . .

ow to delimit tectonic t ~ lithostratiaraphy~ J ~uHnitswithin metamorphicF~L ~ - -

I

! ~ |

~

terrains? ~ ~

S~qtlellC~ OT ~IID'@.F'DO_q~

!

| !

| ~-[-5-~-7imprints

. . . . . . .

~ ~ oorlllflLlrlI

.|~ ~

| aeTorI~GTIOR

.

I

~metamorphic ]| I [~ ~ ~

.

.

.

.

~

" ~

l ~

'

not sufficient because

I

*

~

portions

l

of

the same may follow

lithostratigraphic unit

different lithostratigraphfc units may foflowthe same tectonic evolution; a renewed lithostratigraphy may arise

fro~stroi~ocolis~tio~(t~ansposition)

| 1) to evaulate the role af fabric gradients on reaction progress for individuating rocks with tong structural metamorphic~ I m~mo~i film whilh struclural and metamorphic evolutions may be reconstructed starting from detailed maps l | 2) to contour rock volumes recording coherent sequencesof tectono metamorphic imprints (: coherent tectano metamorphic~ g tectonometamorphic un ~t

Lm .e m~ 1

ll

1

~

~

.

.

.

.

.

.

1

. . . . .

~

....

~

~

. . . . . . . . .

Fig. 1. Methodology for tectonic analysis in the metamorphic crust. The main steps of the analytical procedure for an improved definition of tectonic units and metamorphic complexes are reported in the lower box and explained in the text.

230

M. IOLE SPALLA ET AL.

2

Fig. 2, Stages 1 to 6 of fabric evolution during the superposed foliation development, redrawn after Bell & Rubenach (1983). 1, original foliation; 2, crenulation; 3, crenulation and incipient differentiation; 4, differentiated crenulation cleavage; 5, obliteration of relict crenulation in Q-domains; 6, new continuous foliation.

assemblages, with the aim to display on a synthetic map the relationships between fabric evolution and degree of metamorphic transformation. Maps of heterogeneity of deformation history may be drawn to further constrain the tectono-metamorphic evolution. Finally, specific metamorphic conditions of successive deformation stages were inferred from the compatibility of mineral assemblages forming the mineralogical support of fabric elements coeval in different bulk compositions. In summary, this representation technique directly conveys the inferred P-T-d-t evolution(s), and visualizes rock memory variations, offering an immediate insight on TMU contours.

The example of the polycyclic basement of the Sesia-Lanzo Zone Geological setting

In the axial portion of the Western Alps, between the Penninic Front and the Insubric Lineament (PF and IL in Fig. 3), an HP-LT metamorphic imprint widely affects the pre-Alpine continental crust. Austroalpine continental fragments, sandwiched with slices of meta-ophiolites, are interpreted as exhumed remnants of the southern plate margin that was subducted during Alpine convergence. The Sesia-Lanzo Zone (SLZ; Fig. 4) is the widest portion of Austroalpine continental crust recording early Alpine eclogitefacies metamorphism. It has been subdivided into a lower (Gneiss Minuti Complex = GMC; Eclogitic Micaschists Complex = EMC) and an

upper element (II Dioritic-Kinzigitic Zone-IIDK; e.g. Dal Piaz et al. 1972; Compagnoni et al. 1977; Pognante et al. 1987). The upper unit records a high-pressure blueschist metamorphic imprint and its mylonitic contact with the lower unit developed under eclogite or blueschist facies conditions (Lardeaux et al. 1982; Pognante et al. 1987). In the lower unit, the Alpine polyphase deformation developed during a LT eclogite imprint, followed by a blueschist re-equilibration and by a low-pressure greenschist facies overprint (e.g. Lardeaux et at. 1982; Vuichard & Ball~vre 1988; Castelli 1991; Pognante 1991 and references therein). The HP-LT metamorphic imprint is early Alpine (130-70Ma; e.g. Oberhaensli et al. 1985; Stoeckhert et al. 1986; Hunziker et al. 1992; Rubatto et al. 1999). The GMC and EMC, both pervasively eclogitized, strongly differ in the volume percentage of greenschist re-equilibration. The GMC represents the continentocean tectonic boundary and is dominantly re-equilibrated under greenschist facies conditions generally associated with mylonitic textures (Sttinitz 1989; Spalla et al. 1991). The EMC constitutes the innermost part of the SLZ and its greenschist facies overprint is confined to discrete shear zones, more pervasively developed towards its inner boundary with the Southern Alps. Contrasting dominance of HP or greenschist metamorphic imprints in EMC and GMC, respectively, led Sttinitz (1989) and Spalla et al. (1991) to the still debated interpretation that the two complexes correspond to a large-scale strain partitioning effect, taking place during greenschist facies re-equilibration. The peculiar low T I P Alpine ratio, during the eclogitic peak (c. 10 ~ -]) and during the exhumation path (_<14 ~ km-1), favoured the preservation of pre-Alpine relic assemblages, which in the three complexes indicate a metamorphic evolution from granulite to greenschist facies conditions (Compagnoni et al. 1977; Lardeaux & Spalla 1991; Rebay & Spalla 2001 and references therein). In the following, the tectono-metamorphic evolution of the SLZ is reconstructed in the EMC using the proposed structural and metamorphic correlation approach, from Monte Mucrone to lower Val d'Aosta (Fig. 3). D e f o r m a t i o n - m e t a m o r p h i s m interactions

As summarized above and in Table 1, the EMC underwent a polycyclic and polyphase tectonometamorphic evolution, lasting from pre-Alpine times to the Tertiary. In the map displaying the structural-petrographic evolution of the

TECTONO-THERMAL MEMORY OF ROCKS

231

Fig. 3. Tectonic sketch map of the Alpine chain: shaded areas correspond to continental crust. Stars locate examples illustrated in the text: Monte Mucrone-Monte Mars area in the Western Austroalpine domain of the Sesia-Lanzo Zone (a), Lake Como area in the Central Southalpine domain (b), and Mortirolo Pass area in the Central Austroalpine domain of the Languard-Campo Nappe (c).

EMC basement between Monte Mucrone and Mombarone (Fig. 4), five superposed groups (D1-D5) of Alpine syn-metamorphic structures are distinguished (Zucali 2002; Zucali et al. 2002b). Details on superposed structures and their relationships, on the mineral assemblages supporting successive fabrics, and on P - T conditions in which they developed are summarized in Table 1 and on the P-T-d-t path of Figure 5. In spite of the extremely pervasive structural reworking, small lenses ( 1 - 1 0 0 m in size) escaped Alpine deformation and preserve preD1 tectonic and igneous fabrics, where the modal amount of granulitic or igneous pre-Alpine minerals is _<15%. Generally, in undeformed rocks, incipient or pervasive metamorphic transformations are characterized by fine-grained aggregates replacing medium to large grain minerals (as for example white mica + epidote + clinopyroxene or kyanite fine-grained pseudomorphs on plagioclase). The preservation of the protolith texture suggests that strain localization due to grain size reduction (reaction softening) was not effective in all these cases, in contrast with local observations in the Monte Mucrone granite (Rubie 1990).

D1 planar fabrics are evolved up to stages 5 - 6 (Bell & Rubenach 1983), whereas D2 planar fabrics evolve from crenulation cleavage (stage 4) to continuous foliation, locally mylonitic (stages 5-6). D3 structures consist of a crenulation (stages 2-3). Eclogite facies minerals mark D1, D2 and D3 fabrics. $4, marked by blueschist facies assemblages, is mylonitic (stage 6) but extremely localized in 10 cm thick domains. At the micro-scale, D5 consists of gentle waving to crenulation (stages 2-3), only locally evolved up to a continuous foliation (stage 6) in narrow zones (up to 2 m thick). Greenschist facies minerals underline the $5 differentiated foliation. In the map of structural versus metamorphic reequilibration (Fig. 5), areas where a specific fabric is dominant are recognizable by the distribution of high grain-scale deformation domains, represented for each group of structures. Details on the degree of metamorphic transformation (metamorphic imprint in percentage of synmetamorphic minerals) are also shown and outline adjacent rock volumes characterized by different memory of the whole tectono-metamorphic evolution. In this map, the dominance of the eclogitic metamorphic imprint is easily perceived, thanks

232

M. IOLE SPALLA E T AL.

Fig. 4. (a) Tectonic sketch of the Sesia-Lanzo Zone. (b) Structural-petrographic map of the Sesia Lanzo Zone in the Monte Mucrone-Monte Mars area, redrawn after Zucali (2002) and Zucali et al. (2002b). to the wide distribution of high grain-scale D2 deformation domains (widespread stages 5 - 6 = HD2). The only exceptions are zones where preD 1 fabrics dominate: the pre-Alpine mineralogical

relics are localized in these domains. In low grain-scale D2 deformation domains (widespread stage 4 = LD2), the S1 foliation (widespread stages 5 - 6 ) occupies the 7 5 - 8 0 % ,

+& o~ r - -

+~

0

=4 1= .--

0

-H

+~-

~+

II

._~ + 0

~

2~

N

+

.~9 ~ ,~

+

+,~

.~

~+~=~ +1 ~

+~

.~

~

~'~ ~.~ +

= .. + l =~;5 ~ .~E

~E

ii d

+

rj

+ +

'~ +

~9 " ~

~3

~.~ ~'~

I

I=

II

u.~J

..

~+==

~.P

~1"

E.~

_~:<~

~.

+

r.-. -~---

=(.~

~

~ o~

q

~+'~

~.

1H ~,

+~

-I- E -

r~-

+1

~

~

+1 o~

9-

II

--}-

+~. +

ou

E

0

+ =

.=

=

+1 ~.

~

-

+

k3-8

+1 o~

T~,,, o ~

L~o'

E ~

.ff~

+1~

~

II

o

-tll '~

~+~

~

.~+~+,

o

~ ~ov~

+l

+-

0 -l-l

<

~.~

+~

~

o <~

+~ '~

b r~

9~

9~

,~

~1-

~a ~d

.~ + +~

V

eq

~+~ ~dV

-~+~

b~

2~

eq

~6 +

~

+

++ ~4 ~ll

Bo~

o+~

4 ~

~+ ~ , o~

b~

~

~

~ll ~ ~

~o 0

tm

.,=

"~

.= ~,

<

<

~ q

+ .o

p

4'- "~ 4- ~

~+

VI [..,,. ~

"==--

~o

+

<+~ P

0o(D

o

I.

-~ + ~ , ~

~

~~o "

o~,

s r

& V I ~ b:-,~ /',1

~~" + i"q

-I- +i

v,

r~

~'~

+ , 2 '~

oo ro on tt) /oVl

~,w ~ p-_. o~

~ ~

e . _~

,.IN 9~ O 6~ I

o

J

'-"

.~ ,..-4

i;

9

Oh

-~

--4 =j o, O'~ o r

o

""

r-,

~ ~O ~., ~

U~ o

O'~ oO

+

+

~G (,.)

~+~

~ ~

~

+U

~.o

~

~'

it,,

~Vl

t~ <

+~

v,

+r.2

[~

~.~D~ I oo ~ V l

:=

~3

~-~ ,,~

~:.q

E

e~

'-'d

r

(J

236

M. IOLE SPALLA E T AL.

Fig. 5. Tectono-metamorphic outline of Monte Mucrone-Monte Mars area, redrawn after Zucali et al. (2002b). The metamorphic history (P-T-d-t path) is shown together with the map of deformation imprints, related fabric evolution and degree of metamorphic transformation. The degree of new planar synmetamorphic fabric evolution (dominant fabric), the deformation imprint and metamorphic transformation (metamorphic imprint in percentage of syntectonic minerals) have been estimated and reported in the two tables.

TECTONO-THERMAL MEMORY OF ROCKS while the $2 crenulation cleavage slightly (15-20%) overprints S1; also in this case, the dominant metamorphic imprint will be eclogitic, since the eclogitic mineral assemblage marks S 1. In contrast, the blueschist and greenschist metamorphic imprints are never dominant at the map scale, except in thin domains with high grain-scale D4 and D5 deformation (extremely localized stage 6), due to the low degree of fabric evolution reached during these deformation stages (widespread stages 2-3). Summary and discussion

In this portion of SLZ, a positive correlation between degree of fabric evolution and progress of metamorphic transformation occurs. The dominant metamorphic imprint, which in this case matches the Tmax-PTmaxof the Alpine P-T evolution, is eclogitic and corresponds to the dominant regional scale fabric ($2). Even if the highest T mineral assemblages mark pre-D1 fabrics, their mineral relics are scarce and confined to rock-volumes that escaped development of Alpine fabrics. The map of heterogeneity of deformational (and metamorphic) imprints also stresses the need for analysing a critical volume before attributing a regional significance to structural correlations, since the recorded deformation sequence within each of the adjacent areas may change due to broad-scale deformation heterogeneity. As an example, we can compare the prevailing deformation sequence in adjacent subareas: in the Monte Mars region, the sequence corresponds to the pre-D1 and D2 deformations; in the area of Monte Rosso, the successive deformations are all recorded, except D3, while in the area of Monte Bechit-Monte Roux, D1 to D5 are recognized. Therefore, the confidence of such a structural-petrographic map as a correlation tool is strongly scale dependent. This correlation procedure, applied at the scale of numerous GMC, EMC and II DK junctions, would greatly help to clarify which of the three complexes coincides with crustal-scale strain partitioning zones or with different TMUs.

Example of the Lake Como Southalpine deep crust Geological setting

The South-Alpine Domain of the Central Alps is the hinterland of the Alpine arc and represents the southern plate margin deformed during Alpine convergence at very shallow structural levels. It is separated from the internal belt of

237

the Alps (Penninic and Austroalpine nappes), representing the suture zone (e.g. Caron et al. 1984), by the Insubric-Tonale tectonic line (IL in Fig. 3) and involves thick-skinned preAlpine basement and Permian-Mesozoic cover sheets (e.g. De Sitter & De Sitter-Koomans 1949; Schumacher & Laubscher 1996). On the two sides of Lake Como, the basement is split into two south-vergent Alpine thrust elements (Musso and Val Colla-S.Marco units) separated by a cataclastic Alpine fault zone (Musso Fault Zone= MFZ; Schumacher & Laubscher 1996) that was already active at depth during pre-Alpine times (Bertotti et al. 1993; Gosso et al. 1997). These two shallowlevel Alpine tectonic units comprise three preAlpine metamorphic zones: the Musso unit coincides with the Domaso-Cortaf6 Zone (DCZ; Fumasoli 1974; Bocchio et al. 1980), a basement portion dominated by a homogeneous metamorphic pre-Alpine evolution; whereas the Val Colla-S.Marco unit comprises the DervioOlgiasca and Monte Muggio Zones (DOZ and MMZ; Gosso et al. 1997 and references therein), which underwent heterogeneous pre-Alpine metamorphic evolutions (Table 1). They are separated by a greenschist-facies mylonitic belt reactivated by cataclastic deformation, the Lugano-Val Grande Fault Zone (LVGFZ; Liassic synsedimentary normal fault; Bertotti et al. 1993). DCZ, DOZ and MMZ contain metapelites with interlayered amphibolites, quartzites, marbles, calcareous schists and metagranitoids; pegmatites occur exclusively in the DOZ biotite-sillimanite schists (Spalla et al. 2000). A Permian-Mesozoic sedimentary cover unconformably overlies the leucocratic metagranitoids of the MMZ south of the LVGFZ; slices of Permian-Mesozoic dolostones, conglomerates and siltstones rest tectonically on the DCZ basement located north of the MFZ greenschist mylonites. Non-metamorphic PermianMesozoic sediments are lacking over the DOZ metamorphic basement. The P-T-d-t paths (Fig. 6) are characterized by intermediate T/depth ratio (Barrovian metamorphic gradient), during the amphibolite facies P-peak (c. 20 ~ k m - b and during the exhumation path for MMZ and DCZ; whereas DOZ exhumation occurred under high T/depth ratio (c. 50 ~ km -1) during the earlier stage of decompression. D e f o r m a t i o n - m e t a m o r p h i s m interactions

The structural and metamorphic evolutions, together with the available radiometric data for DCZ, DOZ and MMZ are synthesised in

238

M. IOLE SPALLA ET AL.

Fig. 6. Tectono-metamorphic outline of Lake Como pre-Alpine basement, redrawn after Spalla et al. (2000). Metamorphic histories (P-T-d-t paths) of rocks from DCZ, DOZ and MMZ tectono-metamorphic units show contrasting structural and thermal evolutions recorded in three zones from Variscan subduction to Permian-Triassic continental rifting; see Table 1 for the relative radiometric ages. Unperturbed (Vi) and maximally relaxed geotherm (Voo) are after England & Thompson (1984); aluminum silicate triple point after Holdaway (1971). Areas with different dominant metamorphic imprints are distinguished and their conditions indicated with corresponding shading on each P-T-d-t path. The dominant metamorphic imprint coincides with Tmax-PTmax only in the northern portion of DOZ, where the associated fabric ($2) is the most evolved.

TECTONO-THERMAL MEMORY OF ROCKS Table 1 and in the structural-metamorphic map containing the regional grid of superposed foliations and indications of the metamorphic environments of their development and the P-T-d-t evolutions (Fig. 6). P-T-d-t paths have been used as a discrimination tool, instead of lithostratigraphic associations or of the dominant metamorphic imprint, to distinguish three different metamorphic histories in MMZ, DOZ and DCZ (Diella et al. 1992; Bertotti et al. 1993; di Paola & Spalla 2000; Spalla et al. 2002) that therefore appear as TMUs. Muggio TMU (MMZ). MMZ basement rocks underwent a polyphase tectonometamorphic evolution pre-dating the Permian sedimentary sequences (Bertotti et al. 1993; Gosso et al. 1997). In the map displaying the structural-petrographic evolution of the MMZ, D1 and D2 Variscan synmetamorphic structures and related metamorphic imprints are represented: successive planar fabrics (S1 and $2) are marked by mineral assemblages corresponding to amphibolite and greenschist facies conditions, respectively (di Paola & Spalla 2000; Spalla et al. 2002). At the micro-scale, S1 foliation is spaced or continuous (Passchier & Trouw 1996), evolved up to stages 5-6, whereas $2 foliation is evolved up to stage 4, or more generally stage 5 of Bell & Rubenach (1983) and may be locally mylonitic (stage 6). These characters favoured the attribution of the MMZ to the low-grade zone in a Southalpine metamorphic zonation scheme (Vai et al. 1981). The dominance of the greenschist metamorphic imprint in the MMZ basement, is a consequence of the high grain-scale $2 evolution.

239

case, $1 is generally evolved up to stages 5 - 6 (no relics of a previous oriented fabric where the foliation is spaced). On the contrary, where the $2 planar fabric is mylonitic, or evolved up to stages 5 and 6 with s-c geometries, biotitesillimanite-bearing metapelites dominate (Fig. 6). In this case, the few remnants of the earlier tectono-metamorphic stage consist of rare mineral relics, and the HT-LP metamorphic imprint dominates. Towards the LVGFZ, a generally continuous $3 foliation is the most penetrative fabric element and syn-D3 greenschist facies mineral assemblages are widespread (Fig. 6).

Monte

Three groups of synmetamorphic structures were recorded in DOZ rocks during Variscan and PermianLiassic times (Diella et al. 1992; di Paola & Spalla 2000). Mineral assemblages corresponding to amphibolite, amphibolite/granulite (HTLP) and greenschist facies conditions mark S1 (Variscan), and $2 and $3 (Permian-Liassic) planar fabrics, respectively (see Fig. 6). The DOZ encompasses portions where the most pervasive fabric ($2) reached different evolutionary stages: where St-beating metapelites are abundant, the $2 planar fabric has evolved up to stages 3 and 4 (Bell & Rubenach I983). Here, the HT-LP minerals (Ti-rich biotite and sillimanite) growing along $2 are scarce and syn-D1 amphibolite-facies metamorphic imprint dominates (Fig. 6). In this Dervio-Olgiasca TMU (DOZ).

Domaso-Cortafb TMU (DCZ). Three superposed planar fabrics (S1, $2 and $3) developed in DCZ during its pre-Permian metamorphic evolution (di Paola & Spalla 2000; di Paola et al. 2001), each of them under different metamorphic conditions: epidote amphibolite-facies conditions for D 1, amphibolite-facies during D2 and greenschist-facies for D3 planar fabrics. D3 metre- to kilometre-scale folds associated with a spaced axial planar foliation (stage 4), or with grain-scale reactivation of the earlier $2, are the dominant structural character (di Paola et al. 2001); in places $3 is mylonitic (stage 6). The syn-D3 greenschist metamorphic imprint is dominant in DCZ rocks, and epidote-amphibolite or amphibolite facies mineral assemblages dominates in centimetre- to metre-scale domains that escaped D3 pervasive structural reworking. In particular, epidote-amphibolite facies mineral assemblages are preserved exclusively northwest of Domaso (Fig. 6), where the D2 microstructure is poorly evolved (stages 2-3). Summary and discussion

In the three TMUs (DCZ, DOZ and MMZ) the foliation trajectory map (Fig. 6) constrains with confidence the location, thinness (~7 kin) and extent of each TMU. Considering metamorphic and planar fabric evolution for the three units, the dominant metamorphic imprint does not coincide with Tmax-PTmaxin each P-T-d-t path because it always corresponds to that of the most pervasive fabric, for which the degree of grain-scale reorganization of the planar fabric always overstepped stage 4 (Bell & Rubenach 1983). This is highlighted in Figure 6, where the dominant metamorphic imprints recorded in the three units are shown in the P-T-d-t loops and are represented with different patterns on the interpretative structural map, together with the boundaries of the DCZ, DOZ and MMZ TMUs.

240

M. IOLE SPALLA ETAL.

Circled D1, D2 and D3 correspond to P - T conditions for assemblages formed during each deformation stage. Patterns attributed in the P - T diagrams to different tectono-metamorphic imprints, developed during the first, second and third generation of structures, correspond on the map to volumes in which each P - T re-equilibration stage is dominant. The mismatch between the dominant metamorphic imprint and Tma• is clear in DOZ, where different dominant metamorphic imprints occupy adjacent portions of this unit, each corresponding to sectors characterized by different stages of D2 fabric evolution, or where $3 dominates. Taking into consideration the D2 fabric heterogeneity, where the $2 planar fabric is less evolved (up to stage 3 and 4), the syn-D1 intermediate pressure amphibolitefacies minerals are dominant in volume; whereas if the $2 planar fabric is more evolved (up to stages 5-6), or if the foliation is mylonitic, the HT-LP metamorphic mineral assemblages prevail and the dominant metamorphic imprint coincides with the Tmax-PTmax experienced by the DOZ. In the MMZ rocks (Fig. 6), the dominant D2 greenschist metamorphic imprint does not coincide with Tmax-Prmax that was recorded under the conditions characterizing the D1 tectono-metamorphic stage. The same holds for DCZ, where the most pervasive syn-D3 metamorphic imprint does not appear to correspond with Tmax-Prmax conditions of the P-T-d-t loop, which instead match the earlier D2 tectonometamorphic stage.

In the Upper Valtellina-Upper Val Camonica region, three lithostratigraphic units were distinguished on the ground of lithostratigraphic settings and dominant metanaorphic imprints, suggesting different depths of crustal derivation (Ragni & Bonsignore 1966; Bonsignore et al. 1971): low- to medium-grade metapelites and metaintrusives of 'Scisti di Pietra Rossa Series' (SPRS coinciding with LCN), medium-grade metapelites of 'Scisti della Cima Rovaia Series' (SCRS part of TS), and high-grade gneisses of 'Scisti del Tonale Series' (STS part of TS). In these units (Fig. 7), Permian intrusives (granitoids, diorites and minor gabbroids) commonly occur variably reworked by Alpine tectonics during the Cretaceous (Del Moro et al. 1981; Tribuzio et al. 1999 and references therein). Alpine polyphase deformation and metamorphic transformations are heterogeneously recorded by Permian intrusives and their host rocks. The P-T-d-t path resulting from the polycyclic structural and metamorphic evolution is characterized by high T/depth ratio during the pre-Alpine exhumation (>_50 ~ km -1) and by low T/depth ratio during the Alpine burial (c. 17 ~ -1) and exhumation. The Alpine tectono-metamorphic evolution is fully shared either by the LCN and TS basement units (Gazzola et al. 2000; Zucali 2001), making existing lithostratigraphic subdivisions trivial with respect to reconstructions of tectonothermal records. The two units may therefore be grouped into a single Alpine TMU: the Languard-Tonale unit. D e f o r m a t i o n - m e t a m o r p h i s m interactions

Example of the polycyclic basement of the Languard-Campo Nappe G e o l o g i c a l setting

In the Austroalpine units of Central Alps, the Alpine HP-LT metamorphic imprints, widespread in western and eastern Alps, are rare (e.g. Dal Piaz et al. 1972; Compagnoni et al. 1977; Vogler & Voll 1981; Hoinkes et al. 1991; Thoeni & Jagoutz 1993). Along the upper Val CamonicaValtellina ridge, the Upper Austroalpine units include the Languard-Campo Nappe (LCN) and the Tonale Series (TS), both constituted of polymetamorphic rocks (Figs 3 and 7). The LCN and TS display a steeply dipping attitude immediately north of the Insubric line (Southern Steep Belt; Schmid et al. 1996), and LCN is the uppermost unit of the Central Alps nappe pile. In the past, LCN has been distinguished from TS on lithological ground (Venzo et al. 1971; Schmid et al. 1996 and references therein).

Permian intrusives, used to distinguish Alpine from pre-Alpine structures and metamorphic imprints, and areal extent of strain states corresponding to undeformed up to mylonitic textures (fabric gradients), are shown in Figure 7. These rocks are interlayered as bodies of various sizes and are deformed together with the LCN and TS gneisses. In the country rocks, two groups of structures (D1 and D2) formed under medium- to high-grade metamorphic conditions before the Permian; three groups of synmetamorphic structures (D3, D4 and D5) developed in Permian intrusives and host rocks during the Alpine orogenic cycle under epidoteamphibolite (D3) and greenschist facies conditions (D4 and D5). P - T estimates of each structural and metamorphic stage are reported in Table 1. Chemical compositions of longliving minerals, as amphiboles in metadiorites, vary in the different microstructural sites and are re-homogenized as a function of the degree

TECTONO-THERMAL MEMORY OF ROCKS

241

Fig. 7. Structural-petrographic map of the Languard-Tonale TMU in the Mortirolo pass area simplifiedafter Gosso et al.

2004.

of the fabric evolution (from undeformed igneous to mylonitic textures; Fig. 8). Pre-Permian S1 fabrics, preserved in lowstrain D2 domains (up to 10m scale), are spaced or continuous and differentiated foliations, always evolved up to stages 5 and 6 (Bell & Rubenach 1983). In the rare domains where S1 and the associated assemblages dominate, $2 planar fabrics evolved up to stages 3 and 4 (Bell & Rubenach 1983), corresponding to a crenulation and a differentiated crenulation cleavage, respectively. Generally, $2 is a differentiated foliation, continuous or not, without relics of previous planar fabric in the microlithons (stages 5 or 6). In this case, the HT sillimanite-biotite-bearing assemblage is widespread (HD2 = high grain-scale D2 deformational domains), as represented in the map of Figure 9. Here, areas dominated by a new synmetamorphic fabric are shown together with the degree of metamorphic transformation (deformational vs. metamorphic imprint). As for the Sesia Lanzo basement (Fig. 5), a direct appraisal of adjacent rock volumes characterized by a different memory of the tectono-metamorphic evolution is facilitated.

In metapelites, the earlier Alpine fabric $3 is a differentiated foliation with evolutionary stages from crenulation cleavage (stage 4) up to continuous foliation (stage 6) in the central sector of the map, whereas D3 is expressed as a simple crenulation (stage 3) in the southern and northern parts. In metaintrusives, $3 is a continuous foliation with local mylonitic characters (stages 5 and 6; Fig. 8). In volumes with most evolved $3 (in Fig. 9, grain-scale D3 high deformation domains = HD3 corresponding to mylonitic textures of Fig. 8), the HP Alpine metamorphic imprint is dominant. In addition, strain partitioning during D3 strongly influenced the fabric evolution during successive deformation stages. D4 structures consist of open to tight folds, only locally (Monte Pagano and Passo Foppa in Fig. 7) associated with a differentiated axial plane foliation ($4) of continuous type (stages 5 and 6; HD4 in Fig. 9). Strain heterogeneities related to the extremely localized D5 structures are not represented on Figure 8. In volumes poorly affected by Alpine deformations (D1, LD2 and HD2 in Fig. 9), Alpine minerals developed as finegrained aggregates, pseudomorphic upon preAlpine minerals, or forming reaction rims around

242

M. IOLE SPALLA ET AL.

them, but the degree of metamorphic transformations is incomplete with respect to the pervasively foliated volumes (HD3 and HD4 in Fig. 9). Also in these rocks, as in those of the Sesia-Lanzo Zone, the preservation of protolith textures where new metamorphic assemblages consist of fine-grained aggregates replacing larger mineral grains, suggests that strain was not localized by grain size reduction (reaction softening). Summary and discussion

In this portion of the Austroalpine basement, a positive correlation occurs between the degree

of fabric evolution and the progress of metamorphic transformation (Fig. 8), in agreement with observations from the Western Austroalpine and the Central Southalpine. If the distribution of high and low grain-scale Alpine deformation domains (LD3, HD3, LD4 and HD4) is compared with the boundaries previously proposed for the three lithostratigraphic units (Fig. 7), these three units coincide with zones of different intensity of Alpine deformation. Actually, SPRS and SCRS correspond to zones where the dominant fabrics are pre-Alpine (D1, LD2 and HD2) or where low grain-scale Alpine deformation domains occur (LD3 and LD4), and

Fig. 8. The influence of fabric evolution on metamorphic reaction rate is semiquantitativelyconstrained by a comparison of compositional evolution in amphiboles from metadiorites deformed during D3 (a-t) together with quantitative fabric data (QuantitativeTextures Analyses - QTA; g, h, i). In coronitic metadiorites exclusively (a, d), amphibole compositions cover the full range from Permian igneous high-Ti amphibole to Alpine metamorphic tschermakitic amphibole, whereas in mylonitic diorites (e, f) no Ti-rich amphibole compositions are preserved (Gazzola et al. 2000). Plane polarized light photomicrographsshow: (a) undeformed metadiorites where an Alpine light-colouredamphibolegrows at the rims of igneous amphibole (coronitic texture); (b) normally foliated metadiorites where pre-Alpine dark coloured amphibole is preserved in cores of amphiboleporphyroclasts re-oriented parallel to $3; (c) mylonitic metadiorites, where no pre-Alpine minerals are preserved. QTA of amphiboles (g, h, i) involved in synD3 metre-size shear zones: from undeformed to mylonitic domains reciprocal directions tend to be parallel to the $3 foliation plane. The symmetry among lattice orientations and $3 foliation is monoclinic in normally foliated domains, while it becomes more close to orthorhombic in mylonitic domains; similar relationships were shown for non-coaxial deformation of quartz and calcite (Lister & Hobbs 1980; Lister 1981; Gapais & Cobbold 1987; Law 1990; for references on QTA using neutron diffraction see Zucali et al. 2002a).

TECTONO-THERMAL MEMORY OF ROCKS

243

Fig. 9. Tectono-metamorphic outline of Languard-Tonale TMU in Mortirolo area, redrawn after Zucali (2001). Metamorphic history (P-T-d-t path; Gazzola et al. 2000) is shown together with the map of deformation imprint, to relate fabric evolution to degree of metamorphic transformation.Table of metamorphic imprint versus dominant fabric estimates the relations between metamorphic imprint (D 1, D2, D3, D4) and degree of synmetamorphicdominant fabric development ($2, $2, $3, $4); total metamorphic imprint represents the accumulation of each metamorphic imprint component. Areas with different dominant metamorphic imprints are distinguished and represented with corresponding tectono-metamorphic stages on the P-T-d-t path; the dominant metamorphic imprint of different crustal volumes always corresponds to the one marking the most pervasive fabric.

where Alpine metamorphic transformations are mainly coronitic. Even in this case, the dominant metamorphic imprint always corresponds to that of the most pervasive fabric, where the degree of grain-scale reorganization exceeded stage 4. This is well represented in Figure 8, where different metamorphic imprints, dominant in adjacent portions of this single Alpine TMU, characterize sectors in which different evolutionary stages of Alpine or pre-Alpine fabrics prevail. The map of heterogeneity of deformation history, as in the case of Sesia-Lanzo Zone (Fig. 9), shows that the reconstruction of deformation versus metamorphism relationships requires a critical mapping extent before attributing a regional significance to structure correlations that are strongly influenced by

deformation heterogeneity. The different sequences of prevailing deformation stages indicate a contrasted structural and metamorphic rock memory in adjacent portions of a single Alpine TMU.

Conclusions The analysis of interactions between fabric and reaction gradients in three localities from Western and Central Alps demonstrates that their positive relationships are recurrent and not simply due to strain localization induced by reaction softening. This is suggested by the overgrowth of fine grain size Alpine minerals on pre-Alpine large grains in poorly deformed

244

M. IOLE SPALLA ET AL.

rocks from Austroalpine units; these rocks do retain the longest tectono-thermal memory. As shown on the map of deformational and metamorphic imprints, the heterogeneity of each deformation imprint determines the minimal critical size of the structural-petrographic map that permits regional validation of the P-T-d-t evolutions; regional validity is reached when the map includes the complete mosaic of heterogeneities of the deformation history. The analysis of the tectono-thermal rock memory allowed us to define TMUs that are very different in shape from units drawn on the basis of dominant metamorphic imprints or of pure lithostratigraphic settings for the following reasons (1)

(2)

(3)

The dominant metamorphic imprint does not coincide with Zmax-PTmax of each P-T-d-t loop; actually, the dominant metamorphic imprint is the one expressed by the mineralogical support of the most pervasive fabric, provided the degree of grainscale reorganization overstepped stage 4 of decrenulation. The regional distribution of dominant metamorphic imprints does not necessarily correspond to the 'metamorphic field gradient', but is strongly influenced by heterogeneities of deformation. Therefore, it can be used to determine TMU sizes only if the areal distribution of superposed synmetamorphic fabrics has been analysed with a map fully displaying heterogeneities of the deformation history. Indeed, in the examined cases within a single TMU, different metamorphic imprints may dominate (e.g. EMC in Sesia-Lanzo Zone, Languard Tonale unit or DOZ in Southalpine domain), or an equivalent dominant metamorphic imprint can occur in different TMUs (e.g. MMZ and DCZ in Southalpine domain). The TMU contours, traced by P-T-d-t paths, do not coincide with lithostratigraphic units boundaries; in fact the lithostratigraphic subdivisions proposed in the Central Austroalpine basement were shown to be trivial, because they simply correspond to zones characterized by different intensities of Alpine deformation. Actually, following the exposed criteria, the three lithostratigraphic units forming the two complexes LCN and TS constitute a single TMU (Languard-Tonale unit) of Alpine age.

In nappe systems piled up during subductioncollision deep-seated processes, the size of the

tectonic units that contributed to thickening of the metamorphic crust may be more confidently constrained with an integrated structural and petrographic approach. This method is efficient for different thermal regimes, from HP-LT to HT-LP, as testified by the study of Alpine-type metamorphic basements. Suggestions and critical readings by P. Agard, M. Ballrvre and D. Gapais greatly improved the text. Funding was supplied by MURST 'Cofin' and CNR-IDPA-Milano. A. Rizzi, C. Malinverno and G. Chiodi provided EDS chemical analyses, thin sections and photomicrographs, respectively. M. Zucali thanks Bachir Ouladdiaf and Institute Laue Langevin (ILL, Grenoble) for supporting neutron diffraction analyses performed during his PhD thesis.

References AUSTRHEIM, H. 1990. The granulite-eclogite transition, a comparison of experimental work and a natural occurrence in the Bergen Arcs, western Norway. Lithos, 25, 163-169. BELL, T. H. & RUBENACH, M. J. 1983. Sequential porphyroblast growth and crenulation cleavage development during progressive deformation. Tectonophysics, 92, 171 - 194. BERTOTTI, G., SILETTO, G. B. & SPALLA,M. I. 1993. Deformation and metamorphism associated with crustal rifting, Permian to Liassic evolution of the Lake Lugano-Lake Como area (Southern Alps). Tectonophysics, 226, 271-284. BOCCHIO,R., CRESP1,R., LIBORIO,G. & MOTTANA,A. 1980. Variazioni composizionali delle miche chiare nel metamorfismo progrado degli scisti sudalpini dell'alto lago di Como. Memorie di Scienze Geologiche, Padova, 34, 153-176. BOCKEMUEHL,V. C. & PFISTER, H. 1985. Geologie der Serottini-Intrusion. Schweizerische Mineralogische und Petrographische Mitteilungen, 65, 79-94. BONSIGNORE, G. CASATI,P. et al. 1971. Note Illustrative della Carta Geologica d'Italia alla scala 1,100.000, Fogli 7 e 18, Pizzo Bernina e Sondrio. Nuova Tecnica Grafica - Servizio Geologico Italiano. BROWN, M. 2001. From microscope to mountain belt, 150 years of petrology and its contribution to understanding geodynamics, particularly the tectonics of orogens. Journal of Geodynamics, 32, 115-164. CARON J.M. ~ POLINO R. et al. 1984. Ou sont les sutures majeures dans les Alpes internes? (transversale Brianqon-Torino). Memorie Societgt Geologica Italiana, 29, 71-78. CASTELLI, D. 1991. Eclogitic metamorphism in carbonate rocks, the example of impure marbles from the Sesia-Lanzo Zone, Italian Western Alps. Journal of Metamorphic Geology, 9, 61-77. CLOOS, M. 1993. Lithospheric buoyancy and collisional orogenesis, subduction of oceanic plateaus, continental margins, island arcs, spreading ridges

TECTONO-THERMAL MEMORY OF ROCKS and seamounts. Geological Society of America Bulletin, 105, 715-737. COMPAGNONI, R., DAL PIAZ, G. V., HUNZlKER, J. C., Gosso, G., LOMBARDO, B. & WILLIAMS, P. F. 1977. The Sesia-Lanzo Zone, a slice of continental crust, with alpine HP-LT assemblages in the Western Italian Alps. Rendiconti Societgt ltaliana Mineralogia e Petrologia, 33, 281-334. CONNORS, K. A. & LISTER, G. S. 1995. Polyphase deformation in the western Mount Isa Inlier, Australia; episodic or continuous deformation? Journal of Structural Geology, 17, 305-328. DAL PIAZ, G. V. 1993. Evolution of Austroalpine and Upper Penninic basement in the Northwestern Alps from Variscan convergence to post-Variscan extension. In: YON RAUMER, J. & NEUBAUER, F. (eds) Pre-Mesozoic Geology in the Alps. Springer, Berlin, 325-342. DAL PIAZ, G. V., HUNZIKER, J. C. & MARTINOTTI, G. 1972. La Zona Sesia - Lanzo e l'evoluzione tettonico-metamorfica delle Alpi Nordoccidentali interne. Memorie Societ~ Geologica Italiana, 11, 433 -460. DAL PIAZ, G. V., VENTURELLI, G. & SCOLARI, A. 1979. Calc-alkaline to ultrapotassic post-collisional volcanic activity in the internal northwestern Alps. Memorie di Scienze Geologiche, Padova, 32, 4-15. DE SITTER, L. U. & DE SITTER-KOOMANS, C. M. 1949. Geology of the Bergamasc Alps, Lombardia, Italy. Leidse Geologische Mededelingen, 14, 1-257. DEE MORO, A., NOTARPIETRO, A. & POTENZA, R. 1981. Revisione del significato strutturale delle masse intrusive minori dell'alta Valtellina, risultati preliminari. Rendiconti Societh Italiana Mineralogia e Petrologia, 38, 89-96. DI PAOLA, S. & SPALLA, M. I. 2000. Contrasting tectonic records in pre-Alpine metabasites of the Southern Alps (Lake Como, Italy). Journal of Geodynamics, 30, 167-189. DI PAOLA, S., SPALLA, M. I. & GOSSO, G. 2001. New structural mapping and metamorphic evolution of the Domaso-Cortaf6 Zone (Southern Alps - Lake Como). Memorie di Scienze Geologiche, Padova, 53, 1-14. DI VINCENZO, G. & PALMERI, R. 2001. An 40Ar-39Ar investigation of high-pressure metamorphism and the retrogressive history of mafic eclogites from the Lanterman Range (Antarctica), evidence against a simple temperature control on Argon transport in amphibole. Contribution to Mineralogy and Petrology, 141, 15-35. DIELLA, V., SPALLA, M. I. & TUNESI, A. 1992. Contrasted thermo-mechanical evolutions in the Southalpine metamorphic basement of the Orobic Alps (Central Alps, Italy). Journal of Metamorphic Geology, 10, 203-219. EL TAHLAWI, M. R. 1965. Geologie und Petrographie des nordoestlichen Comerseegebietes (Provinz Como, Italien). PhD thesis, ETH Ziirich. ENGLAND, P. C. & RICHARDSON, S. W. 1977. The influence of erosion upon the mineral facies of rocks from different metamorphic environments.

245

Journal of the Geological Society, London, 134, 201-213. ENGLAND, P. C. & THOMPSON, A. B. 1984. PressureTemperature-Time paths of regional metamorphismI. Heat transfer during the evolution of regions of thickened continental crust. Journal of Petrology, 25, 894-928. ERNST, W. G. 2001. Subduction, ultrahigh-pressure metamorphism, and regurgitation of buoyant crustal slices - implications for arcs and continental growth. Physics of the Earth and Planetary Interiors, 127, 253-275. FUMASOLI, M. W. 1974. Geologie des Gebietes noerdlich und suedlich der Jorio-Tonale-Linie im Westen yon Gravedona. PhD thesis, E.T.H. Ztirich. GAPAIS, D. & COBBOLD, P. R. 1987. Slip system domains; 2, Kinematic aspects of fabric development in polycrystalline aggregates. Tectonophysics, 138, 289-309. GARZANTI, E. & SCIUNNACH,D. 1997. Permo-Triassic evolution - constraints from the sedimentary records. Ofioliti, 22, 147-152. GAZZOLA, D., Gosso, G., PULCRANO, E. & SPALLA, M. I. 2000. Eo-Alpine HP metamorphism in the Permian intrusives from the steep belt of the central Alps (Languard-Campo nappe and Tonale Series). Geodinamica Acta, 13, 149-167. Gosso, G., BRIZlO, D. et al. 1983. Due cinematiche possibili per la coppia di falde Brianzonese ligure - Flysch a Elmintoidi. Memorie Societh Geologica Italiana, 26, 463-472. Gosso, G., SILETTO, G. B. & SPALLA, M. I. 1997. H-T/L-P metamorphism and structures in the South-Alpine basement near Lake Como, Orobic Alps; intracontinental imprints of the PermoTriassic rifting. Ofioliti, 22, 133-145. Gosso, G., SALVI, F., SPALLA, M. I. & ZUCAL1, M. 2004. Map of deformation partitioning in the polydeformed and polymetamorphic Austroalpine basement of the Central Alps (Upper Valtellina and Upper Valcamonica, Italy). In: PASQUAI~E,G., VENTURINI, C., (eds) Mapping Geology in Italy, APAT, Roma, 291-300. HOBBS, B. E., MEANS, W. D. & WILLIAMS, P. F. 1976. An Outline of Structural Geology. Wiley, New York. HOLDAWAY,M. J. 1971. Stability of andalusite and the aluminium silicate phase diagram. American Journal of Science, 271, 97-131. HOINKES, G., KOSTNER, A. & THOENI, M. 1991. Petrologic constraints for eoalpine eclo.gite facies metamorphism in the Austroalpine Otztal Basement. Mineralogy and Petrology, 43, 237-254. HUNZIKER, J. C., DESMONS,J. & HURFORD,A. J. 1992. Thirty-two years of geochronological work in the Central and Western Alps, a review on seven maps. Mdmoire de Gdologie, Lausanne, 13, 1-59. INGER, S., RAMSBOTHAM, W., CLIFF, R. A. & REX, D. C. 1996. Metamorphic evolution of the Sesia-Lanzo Zone, Western Alps, time constraints from multisystem geochronology. Contribution to Mineralogy and Petrology, 126, 152-168. JOHNSON, S. E. & DUNCAN, A. C. 1992. Fault identification in complexly deformed schist terrains,

246

M. IOLE SPALLA ET AL.

examples from the USA and Australia. Tectonophysics, 216, 291-308. JOHNSON, S. E. & VERNON, R. H. 1995. Inferring the timing of porphyroblast growth in the absence of continuity between inclusion trails and matrix foliations, can it be reliably done? Journal of Structural Geology, 17, 1203-1206. KOONS, P. O. 1986. Relative geobarometry from high-pressure rocks of quartzofeldspathic composition from the Sesia Zone, Western Alps, Italy. Contribution to Mineralogy and Petrology, 93, 322-334. KRETZ, R. 1994. Metamorphic Crystallisation. Wiley, Chichester. LALLEMAND, S. 1999. La Subduction Ocdanique. Gordon and Breach Science, Amsterdam. LARDEAUX, J. M., GOSSO, G., KIENAST, J. R. & LOMBARDO, B. 1982. Relations entre le mrtamorphisme et la drformation dans la zone SrsiaLanzo (Alpes Occidentales) et le problrme de l'~clogitisation de la crofite continentale. Bulletin de la Societd Gdologique de France, 24, 793-800. LARDEAUX, J. M. & SPALLA, M. I. 1991. From granulites to eclogites in the Sesia zone (Italian Western Alps), a record of the opening and closure of the Piedmont ocean. Journal of Metamorphic Geology, 9, 35-59. LAW, R. D. 1990. Crystallographic Fabrics, a Selective Review of their Application to Research in Structural Geology. In: KNIPE, R. J. & RUTTER, E. H. (eds) Deformation Mechnaisms, Rheology and Tectonics. Geological Society, London, Special Publications, 54, 335-352. LISTER, G. S. 1981. The effect of the basal-prism mechanism switch on fabric development during plastic deformation of quartzite. Journal of Structural Geology, 3, 67-75. LISTER, G. S. & HOBBS, B. E. 1980. The simulation of fabric development during plastic deformation and its application to quartzite; the influence of deformation history. Journal of Structural Geology, 2, 355 -370. MORK, M. B. 1985. A gabbro to eclogite transition on Flemsoy, Suamore, Western Norway. Chemical Geology, 50, 283-310. MOTTANA, A., NICOLETTI, M., PETRUCCIANI, G., LIBORIO, G., DE CAPITANI, L., BOCCHIO, R. 1985. Pre-Alpine and Alpine evolution of the Southalpine basement of the Orobic Alps. Geologische Rundschau, 74, 353-366. MYERS, J. S. 1978. Formation of banded gneisses by deformation of igneous rocks. Precambrian Research, 6, 43-64. OBERHAENSLI, R., HUNZIKER, J. C., MARTINOTTI, G. & STERN, W. B. 1985. Geochemistry, geochronology and petrology of Monte Mucrone, an example of Eo-Alpine eclogitisation of Permian granitoids in the Sesia-Lanzo Zone, Western Alps, Italy. Chemical Geology, 52, 165-184. PARK, R. G. 1969. Structural correlations in metamorphic belts. Tectonophysics, 7(4), 323-338. PASSCHIER, C. W. & TROUW, R. A. J. 1996. Microtectonics. Springer Verlag, Berlin.

PASSCHIER, C. W., MYERS, J. S. & KRONER, A. 1990. Field Geology of High-Grade Gneiss Terrains. Springer Verlag, Berlin. POGNANTE, U. 1985. Coronitic reactions and ductile shear zones in an eclogitized ophiolite metagabbro, Western Alps, northern Italy. Chemical Geology, 50, 99-109. POGNANTE, U. 1991. Petrological constraints on the eclogite- and blueschist-facies metamorphism and P-T-t paths in the Western Alps. Journal of Metamorphic Geology, 9, 5-17. POGNANTE, U., TALARICO, F. & BENNA, P. 1987. Incomplete blueschist re-crystallization in high grade metamorphics from the Sesia-Lanzo unit (Vasario Sparone subunit, Western Alps ophiolites): a case history of metastability. Lithos, 21, 129-142. RAGNI, U. & BONSIGNORE,G. 1966. Carta Geologica dell'alta Valtellina e dell'alta Val Camonica. Istituto Mineralogia Petrografia Geochimica e Giacimenti Minerari, Universith degli Studi di Milano. REBAY, G. & SPALLA, M. I. 2001. Emplacement at granulite facies conditions of the Sesia-Lanzo metagabbros, an early record of Permian rifting? Lithos, 58, 85-104. RUBATTO, D., GEBAUER,D. & COMPAGNONI,R. 1999. Dating of eclogit-facies zircons; the age of Alpine metamorphism in Sesia-Lanzo Zone (Western Alps). Earth and Planetary Science Letters, 167, 141-158. RUBIE, D. C. 1990. Role of kinetics in the formation and preservation of eclogites. In: CARSWELL,D. A. (ed) Eclogite Facies Rocks. Blackie, Glasgow, 111 - 140. SANDERS, C. A. E., BERTOTTI, G., TOMMASINI, S., DAVIES, G. R. & WIJBRANS, J. R. 1996. Triassic pegmatites in the Mesozoic middle crust of the Southern Alps (Italy), fluid inclusions, radiometric dating and tectonic implications. Eclogae Geologicae Helvetiae, 89, 505-525. SCHMID, S. M., BERGER, A. et al. 1996. The Bergell Pluton (Southern Switzerland-Northern Italy), overview accompanying a geological-tectonic map of the intrusion and surrounding country rocks. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 329-355. SCHUMACHER, M. E. & LAUBSCHER, H, P. 1996. 3D crustal architecture of the Alps-Apennine join - a new view on seismic data. Tectonophysics, 260, 349-363. SPALLA, M. I. 1993. Microstructural control on the P-T path construction in the metapelites from the Austroalpine crust (Texel Gruppe, Eastern Alps). Schweizerische Mineralogische und Petrographische Mitteilungen, 73, 259-275. SPALLA, M. I., LARDEAUX,J. M., DAL PIAZ, G. V. & GOSSO, G. 1991. Metamorphisme et tectonique a la marge externe de la zone Sesia-Lanzo (Alpes occidentales). Memorie di Scienze Geologiche, Padova, 43, 361-369. SPALLA, M. I., MESSIGA, B. & Gosso, G. 1995. LT-alpine overprint on the HT-rifting related metamorphism in the steep belt of the LanguardCampo nappe. The Cima Rovaia and Scisti del Tonale units represent two different extents of

TECTONO-THERMAL MEMORY OF ROCKS alpine re-equilibration. International Ophiolite Symposium, Pavia, Abstract Volume, 148. SPALLA, M. I., LARDEAUX,J. M., DAL PIAZ, G. V., Gosso, G. & MESSIGA, B. 1996. Tectonic significance of alpine eclogites. Journal of Geodynamics, 21, 257-285. SPAEEA, M. I., SIEETTO, G. B., OI PAOLA, S. & Gosso, G. 2000. The role of structural and metamorphic memory in the distinction of tectonometamorphic units, the basement of the Como Lake in the Southern Alps. Journal of Geodynamics, 30, 191-204. SPAELA,M. I., OI PAOEA,S., GOSSO,G., SIEETTO,G. B. & BISTACCHI, A. 2002. Mapping tectono-metamorphic histories in the Lake Como Basement (Southern Alps, Italy). Memorie di Scienze Geologiche, Padova, 54, 1-25. SPEAR, F. S., SEEVERSTONE, J., HICKMOTT, D., CROWEEY, P. & HODGES, K. V. 1984. P-T paths from garnet zoning. A new technique for deciphering tectonic processes in metamorphic terrains. Geology, 12, 87-90. STOECKHERT, B., JAEGER,E. & VOLE, G. 1986. K-Ar age determinations on phengites from the internal part of the Sesia zone, lower Aosta Valley (Western Alps, Italy). Contribution to Mineralogy and Petrology, 92, 456-470. ST~3NITZ, H. 1989. Partitioning of metamorphism

and deformation in the boundary region of the 'Seconda Zona Diorito-Kinzigitica', Sesia Zone, Western Alps. PhD thesis, ETH, ZiJdch. TAO, W. C. & O'CONNEE, R. J. 1992. Ablative subduction, a two-sided alternative to the conventional subduction model. Journal of Geophysical Research, 97, 8877-8904. THOENI, M. & JAGOUTZ,E. 1993. Isotopic constraints for eo-Alpine high-P metamorphism in the Austroalpine nappes of the Eastern Alps, bearing on Alpine Orogenesis. Schweizerische Mineralogische und Petrographische Mitteilungen, 73, 177-189. TRIBUZIO, R., THIRLWAEL, M. F. & MESS1GA, B. 1999. Petrology, mineral and isotope geochemistry of the Sondalo gabbroic complex (Central Alps, Northern Italy), implications for the origin of post-Variscan magmatism. Contribution to Mineralogy and Petrology, 136, 48-62. TURNER, F. J. & WEISS, L. E. 1963. Structural Analysis of Metamorphic Tectonites. MacGrawHill, New York. VAI, G. B., BORIANI,A., RIVAEENTI,G. & SASSI, F. P. 1981. Catena Ercinica e Paleozoico nelle Alpi Meridionali. In: Cento Anni di Geologia Italiana, Societ~ Geologica Italiana, Jubilee Volume, 133-154.

247

VAN ROERMUND, H., LISTER, G. & WILLIAMS, P. F. 1979. Progressive development of quartz fabrics in a shear zone from Monte Mucrone, SesiaLanzo Zone, Italian Alps. Journal of Structural Geology, 1, 43-52. VENZO, S., CRESPI, R., SCHIAVINATO, G. & FAGNANI, G. 1971. Carta geologico-petrografica delle Alpi Insubriche Valtellinesi tra la Val Masino e la Val Malenco (Sondrio). Memorie Societh Italiana di Scienze Naturali, Milano, 19. VOGEER, S. • VOLE, G. 1981. Deformation and metamorphism at the south-margin of the Alps, east of Bellinzona. Geologische Rundschau, 70, 1232-1262. VUICHARD, J. P. t~ BAELI~VRE, M. 1988. Gametchloritoid equilibria in eclogitic pelitic rocks from the Sesia Zone (Western Alps), their bearing on phase realtions in high pressure metapelites. Journal of Metamorphic Geology, 6, 135-157. WILLIAMS, P. F. 1971. Structural analysis of Bermagui area, New South Wales. Geological Society of Australia Journal, 18, 215-228. WILLIAMS, P. F. 1985. Multiply deformed terrains problems of correlation. Journal of Structural Geology, 7, 269-280. ZUCAEI, M. 2001. La correlazione nei terreni meta-

morfici, due esempi dall'Austroalpino occidentale (Zona Sesia-Lanzo) e centrale (Falda LanguardCampo/Serie del Tonale). PhD thesis, Universith degli Studi di Milano. ZUCALI, M. 2002. Foliation map of the 'Eclogitic Micaschists Complex' (M. Mucrone-M. MarsMombarone, Sesia-Lanzo Zone, Italy). Memorie di Scienze Geologiche, Padova, 54, 87-100. ZUCAEI, M., CHATEIGNER, D., DUGNANI, M., LUTTEROTTI, L. t~ OULADDIAF,B. 2002a. Quantitative Texture Analysis of Naturally Deformed Hornblendite Under Eclogite Facies Conditions (Sesia-Lanzo Zone, Western Alps), Comparison Between X-ray and Neutron Diffraction Analysis. In: DE MEER, S., DRURY, M. R., DE BRESSER, J. n. P. & PENNOCK,G. M. (eds) Deformation Mech-

anisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 239-253. ZUCAEI, M., SPAEEA,M. I. & Gosso, G. 2002b. Fabric evolution and reaction rate as correlation tool, the example of the Eclogitic Micaschists complex in the Sesia-Lanzo Zone (Monte Mucrone - Monte Mars, Western Alps Italy). Schweizerische Mineralogische und Petrographische Mitteilungen, 82, 429-454.

Decoupling and its relation to strain partitioning in continental lithosphere: insight from the Periadriatic fault system (European Alps) M. R. H A N D Y , J. BABIST, R. W A G N E R , C. R O S E N B E R G & M. K O N R A D

Geowissenschaften, Freie Universitiit Berlin, D-12249 Germany (e-mail: [email protected]) Abstract: The Periadriatic fault system (PFS) is an array of late orogenic faults (35-15 Ma) in the retro-wedge of the Alpine orogen that accommodated dextral transpression during oblique indentation by the southern Alpine crust. Decoupling along the leading edges of the southern Alpine indenter occurred where inherited lithological and rheological contrasts were accentuated by lateral thermal gradients during emplacement of the warm orogenic retro-wedge next to the cold indenter. In contrast, decoupling within the core and retro-wedge of the orogen occurred in a network of folds and mylonitic faults. In the Eastern Alps, this network comprises conjugate sets of upright, constrictional folds, strike-slip faults and low-angle normal faults that accommodated nearly coaxial NNESSW shortening and E - W extensional exhumation of the Tauern thermal dome. The dextral shear component of oblique convergence was taken up by a discrete, brittle fault parallel to the indenter surface. In the Central and Western Alps, a steep mylonitic backthrust, upright folds, and low-angle normal faults effected transpressional exhumation of the Lepontine thermal dome. Mylonitic thrusting and dextral strike-slip shearing along the steep indenter surface are transitional along strike to low-angle normal faults that accommodated extension at the western termination of the PFS. The areal distribution of poles to mylonitic foliation and stretching lineation of these networked structures is related to the local shape and orientation of the southern Alpine indenter surface, supporting the interpretation of this surface as the macroscopic shearing plane for all mylonitic segments of the PFS. We propose that mylonitic faults nucleate as viscous instabilities induced by cooling, or more often, by folding and progressive rotation of pre-existing foliations into orientations that are optimal for simple shearing parallel to the eigenvectors of flow. The mechanical anisotropy of the viscous continental crust makes it a preferred site of decoupling and weakening. Networking of folds and mylonitic fault zones allow the viscous crust to maintain strain compatibility between the stronger brittle crust and upper mantle, while transmitting plate forces through the lithosphere. Decoupling within the continental lithosphere is therefore governed by the symmetry and kinematics of strain partitioning at, and below, the brittle-to-viscous transition.

Introduction The structure of continental fault zones at strikeslip or oblique-slip plate boundaries varies with depth, from discrete faults and folds in the uppermost 5 km (e.g. Sylvester 1988), through anastomozing mylonitic shear zones in the intermediate to lower crust (Sibson 1977) to large zones of ductile, mylonitic shear in the upper mantle (Teyssier & Tikoff 1998 and references therein). This vertical change in structural style is associated with a change in kinematics: whereas displacement near the surface is partitioned into d i p - and strike-slip zones that bound crustal blocks, in the lithospheric mantle it is apparently accommodated simply within a

broad shear zone (Vauchez & Tommasi 2003). This necessitates a zone of accommodation somewhere in between, usually taken to be the middle to lower crust. Recent debate has centered on the nature and mechanisms of this accommodation (e.g. Tikoff et al. 2002). On the one hand, intracrustal accommodation appears to involve decoupling within subhorizontal detachment layers, a view supported by the geometry of fault systems imaged in reflection seismic profiles across orogens (e.g. Schmid & Kissling 2000) and extended continental lithosphere (e.g. Davis & Lister 1988; Boillot et al. 1995). These show that faults root at 10-15 km depth within the crust, as well as at, or just above, the MOHO. The depths for

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 249-276. 0305-8719/05/$15.00 9 The Geolo2ical Society of London 2005.

250

M.R. HANDY ET AL.

these inferred decoupling zones are corroborated by strength-depth profiles constructed from experimental flow laws for quartz-rich crustal rocks and olivine-rich mantle rocks (e.g. Brace & Kohlstedt 1980; Ranalli & Murphy 1987). On the other hand, intracrustal strain accommodation must also involve vertical coupling in order to explain coincident displacement fields for the upper crust and upper mantle, as inferred from geodetic and teleseismic observations in tectonically active areas (e.g. Holt 2000). Thus, one is left with the paradox that both strain discontinuity (detachment) and continuity (attachment) are required, within the crust and across the crust-mantle boundary. Molnar (1992) argued, partly by reference to the analogue experiments of Richard & Cobbold (1989), that subhorizontal viscous layer(s) at the base of the brittle upper crust maintain vertical strain continuity while facilitating strain partitioning on brittle faults above. They do this because isotropic, viscous media apparently do not transmit shear stresses from below, causing the principal stress axes to be nearly orthogonal to the horizontal base of the overlying crustal blocks. Strain continuum models of Fossen & Tikoff (1998) and Teyssier et al. (2002) predict that intracrustal accommodation occurs within zones of continuous, three-dimensional strain just beneath the brittle crust. In contrast to strain continuum models, rheological models incorporating timedependent brittle behaviour of the upper crust (Tse & Rice 1986; Li & Rice 1987) predict that, while viscous creep and intracrustal attachment prevail at depths below about 15 km, vertical attachment above this level is punctuated by episodic, coseismic detachment. All of these models shed valuable light on some, but not all, aspects of strain accommodation at rheological transitions in the lithosphere. The vexing question remains whether and/or to what extent strain compatibility in the lithosphere is maintained by continuous three-dimensional flow, by episodic detachment, or by some combination of both. Resolving this debate is obviously difficult given our inability to observe the deep levels of active fault zones. Geophysical methods, although helpful in imaging some structural properties 'in situ' (e.g. Mooney & Ginzburg 1986), provide little information on the evolution of structures and kinematics at depth on geological time-scales. An alternative is therefore to examine structures in exhumed, inactive fault systems whose tectonothermal history is sufficiently well known to permit inferences on their dynamic evolution.

The term 'fault' is used in this paper for any continuous surface or zone across which there has been displacement, regardless of whether it is a discrete, cataclasitic zone or a ductile, mylonitic shear zone. We therefore use the modifiers 'brittle' and 'mylonitic' to characterize the type of fault. A 'fault system' comprises several faults that are kinematically and temporally related on the crustal scale. The Periadriatic fault system in the European Alps (Fig. 1) is a natural laboratory for testing hypotheses on the depth-dependence of fault kinematics and dynamics. This exhumed fault system (henceforth abbreviated PFS) trends E - W over a distance of approximately 700 km from northwestern Italy to northern Slovenia. It was active in Tertiary time at the front of an orogenic indenter corresponding to the southern Alps in map view (Fig. 1). Differential exhumation and erosion have exposed sections across this currently inactive fault system at palaeostructural levels ranging from near-surface faulting down to mylonitic shearing at 2 5 - 3 0 k m depth. The PFS is the site of several Tertiary granitoid bodies, which are useful markers for constraining the age and kinematics of displacement. In addition, the PFS is transected by four geophysical profiles (TRANSALP, NFP20-E, NFP20-W, ECORS-CROP in Fig. 1), which provide crucial information on its relationship to the deep structure of the Alps. This paper focuses on three segments of the Periadriatic fault system (boxed areas in Fig. 1) where excellent exposure combined with deep seismic images allow investigation of how strain partitioning during transpression was related to crustal structure. After a brief description of criteria for identifying decoupling, we characterize the structure and kinematics of mylonitic faulting along these segments. These segments are then considered in the broader context of the kinematics and timing of mylonitic faulting and folding along the entire PFS in the Alps, with particular attention paid to the relationship of exhumation and strike-slip faulting to the shape and motion of the southern Alpine indenter. This allows the identification of first-order strain localization patterns that, we argue in the ensuing section, must have been related to decoupling in different levels of the continental crust. This forms the basis for a discussion of how networked folds and shear zones reduce crustal strength while transferring plate-scale forces through the continental lithosphere. The paper concludes with a generic model for partitioned, localized strain and decoupling in orogenic continental crust.

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

251

~o~

~ ~~ ~ 0 ,-~ ,-~ ~ ~.~ ~ ~ .,., ~

o

~ ,-~

~

~.~~ . ~

o

~...

. ~

~ " ~ ,..~ ~

.~

o

-9= ~~ ~ ~

~

o

~:~_-~ o

""~

252

M.R. HANDY ET AL.

Recognizing mechanical decoupling in rocks 'Coupling' within or between tectonic plates is often implicity equated with the transfer of tectonic stresses, especially shear stress, across active faults. In contrast, 'decoupling' is taken to signify the reduction, or even temporary loss, of shear strength on such surfaces, allowing strain to localize and strain rate to increase. Structural geologists use finite strain gradients or strain discontinuities in rocks to locate faults. Such discontinuities are interpreted in a purely kinematic sense to be the sites of detachment. For detachment to occur, the fault rock must have weakened with respect to the surrounding rock, so the kinematic concept of detachment is related to the mechanical notion of decoupling. Recent model-based attempts to use curved foliation patterns and uniform shear-sense indicators across mylonitic shear zones as criteria for intracrustal 'attachment' and mechanical coupling (Tikoff et al. 2002) are ambiguous because these same structures are ubiquitous in exhumed detachment zones, where offset markers indicate that decoupling must have occurred. Furthermore, structures and finite strain gradients that appear to be continuous at one scale of observation can be discontinuous at other scales. Therefore, we opted for a combination of two criteria to identify sites of former decoupling in the exhumed Periadriatic fault system: (1) the size of the fault and, where visible, its displacement compared to the sharpness of the strain gradient across it. Faults with widths of > 1 0 2 m and/or with comparably large displacements were taken as strong evidence for crustal-scale decoupling. Previously published radiometric ages of fault-related exhumation, or better, ages of syntectonic mineralization within fault rocks, were used to establish whether strain gradients at the margins of a fault system formed during the same time interval as the offset of the markers across this system; (2) the type of fault rock (cataclasite or mylonite). By comparing natural microstructures with microstructures in rock-deformation experiments, we identified the deformation mechanisms and inferred the rheology of the fault rock (criteria in Schmid & Handy 1991). For example, the presence of mylonite or foliated cataclasite is diagnostic of aseismic deformation, whereas pseudotachylite is an indicator of frictional melting and coseismic slip (Sibson 1975; Wenk 1978; Maddock 1983). The mutual overprinting of pseudotachylite and mylonite is interpreted in terms of episodic, coseismic slip that alternated

with longer periods of aseismic viscous creep (Hobbs et al. 1986; Magloughlin & Spray 1992; McNulty 1995). Thus, we not only identified faults associated with intracrustal decoupling, but also characterized the short-term mechanical stability of the rocks within these faults.

The Periadriatic fault system Overview

The PFS comprises many differently aged mylonitic and cataclastic faults within the Tertiary Alpine edifice (Fig. 1). For the purposes of this paper, we restrict our analysis to Oligocene to Miocene parts of the PFS, especially the so-called Insubric mylonite belt, shown with thick black lines in Figure 1. Following Argand (1916) and later Schmid et al. (1989), we reserve the term 'Insubric' for Oligo-Miocene deformation in the Alps centered on the PFS. At that time, the Insubric mylonites delimited the warmer, exhuming retro-wedge of the Alpine orogen to the north from colder, southern Alpine units (Schmid et al. 1996; Escher & Beaumont 1997). The latter units had cooled to below 300 ~ already in Jurassic times (Hunziker 1974; Handy & Zingg 1991), making them hard and brittle during Tertiary Alpine collision (Schmid et al. 1989). The Insubric mylonite belt is overprinted by brittle faults, including the sinistral Giudicarie fault (Gu in Fig. 1). The Insubric segments of the PFS described below occupy parts of the Alpine orogen with quite different characteristics. Whereas the eastern segment (Figs 2 to 5) transects the broadest part of the orogen in the TRANSALP section, the central segment (Fig. 6) occupies the transition from the narrowest part of the orogen along the NFP20-E section to the western termination of the PFS (Figs 7 to 9) along the NFP20-W and ECORS-CROP sections. We will return to these differences in the context of strain partitioning, but first describe the structure of these segments in turn. Readers who are in a hurry can skip the new data in the next three sections and advance directly to the regional synthesis for the Oligo-Miocene PFS (Figs 10, 11) at the end of this chapter.

The eastern s e g m e n t

In map view (Fig. 2), the PFS comprises several faults within a slender wedge of the Austroalpine basement just south of the Tauern Window: (1) the Cima Dura fault (CD), a steeply dipping band of

Structure and kinematics.

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

253

Fig. 2. The Periadriaticfault system at the southwestern end of the Tauern window: Structural map. Equal area projections are lower hemispheres.CD, Cima Dura fault; DAV, Defereggen-Antholz-Vals fault; KV, Kalkstein-Valarga fault; Pu, Pustertal fault; Ta, Tauern basement; Ri, Rieserfernerpluton; Re, Rensen pluton. greenschist-facies mylonite; (2) the Defereggen-Antholz-Vals fault (DAV, Borsi et al. 1973), likewise a steeply dipping band of greenschist-facies mylonite up to several hundred metres thick that forms the southern limit of Alpine metamorphism (Hoinkes et al. 1999); (3) the Kalkstein-Valarga fault (KV), a poorly exposed, discontinuous line of cataclasites; and (4) the Pustertal fault (Pu), a line of cataclasites in basement rocks and low-temperature mylonites in Mesozoic carbonates separating the Austroalpine and southern Alpine basements. Insubric deformation engendered the main foliation ($2) north of the DAV fault and coincided with activity of the DAV and CD faults, as well as with the intrusion of the Rieserferner pluton (Wagner 2004). The depth of Insubric faulting exposed in this area is about 10 km based on pressure estimates in and around this syntectonic pluton (290-350 MPa, mineral equilibria in the contact aureole, Cesare 1994; 220-510 MPa; Al-in-hornblende geobarometry in massive tonalite, Wagner 2005). Insubric (F2) folds tighten and become isoclinal as the basement wedge narrows to the west. Their axes are subhorizontal, subparallel to a stretching lineation, L2 (Fig. 2). The main body of the Rieserferner pluton occupies

the core of a tight, upright F2 antiform just to the north of the DAV fault (Fig. 3). These Insubric structures are offset by conjugate, mylonitic to cataclastic shear zones that were active under retrograde, greenschist-facies conditions (dashed-dotted lines in Fig. 2). Several lines of evidence indicate that the Insubric strain field in the Austroalpine basement involved bulk vertical flattening with a strong component of horizontal, E N E - W S W extension. First, the stretching lineation, L2, is subhorizontal and E N E - W S W trending on $2 surfaces, which are generally steeply dipping (Fig. 2). Only in the vicinity of the Rieserferner pluton, where the $2 surfaces are concordant with the pluton's sides and roof (Fig. 3), does the strain field appear to be locally constrictional. Secondly, the largest D2 shear zones are both sinistral (DAV, Kleinschrodt 1987) and dextral (CD, Wagner 2004). This also applies to the somewhat younger conjugate, ductile-brittle faults, which show opposite drag senses of $2 in map view (Fig. 2). Thirdly, the kinematic reconstruction of Frisch et al. (1998, fig. 2a) indicates that 12 km of N - S shortening of the Austroalpine basement wedge was associated with an E - W elongation of 53 km. Vertical motion during Insubric (D2) exhumation of

254

M.R. HANDY ETAL.

Fig. 3. Cross-sectionof the southwestern part of the Tauern window, including the Rieserfernerpluton and DAV fault (modified from Senarclens-Grancy 1965); structures projected into profile section marked by the trace in Figure 2. Symbols as in Figure 2. the Austroalpine basement was at most l0 km according to the radiometric age constraints discussed below. The axes of N - S shortening and E - W stretching are, respectively, normal and parallel to the E - W striking $2 schistosity. Thus, the principle axes of the bulk finite strain ellipsoid during Insubric faulting were oriented E - W (+el), vertical (+e2) and N - S (-e3), with elongation taken to be positive. Insubric shearing in the Austroalpine basement wedge was therefore highly coaxial with a non-coaxial component of sinistral shear, mostly accommodated by the DAV fault. A similar pattern of strain partitioning, but involving dextral transpression, occurs along the M611tal fault at the southeast end of the Tauern window (Fig. 1), as discussed below. Upper greenschist-facies Insubric mylonites along the contact of Austroalpine and LiguroPiemont units in Fig. 2a dip subvertically and have a subhorizontal stretching lineation, indicative of E N E - W S W elongation. A similar stretching direction is inferred from subhorizontal L-tectonites in upright, tight to isoclinal folds in the western Tauern window (Lammerer & Weger 1996). Geochronological evidence discussed in the next section suggests that the southern border of the Tauern window was also the site of significant N-side-up exhumation in Miocene time. Timing o f strike-slip faulting and exhumation.

Age intervals of faulting and exhumation are reconstructed in Figures 4 and 5 from the available geochronological data and from knowledge of the peak temperatures and pressures attained in the fault-bounded crustal blocks. Because metamorphic temperatures in the Austroalpine basement never exceeded about 400 ~ in Tertiary time, and the R b - S r system in white mica closes to diffusion at about 500 ~ (e.g. von Blanckenburg et al. 1989), 3 3 - 3 0 Ma R b - S r white-mica ages on $2 (Mfiller et al. 2001) (stars in Fig. 4) are interpreted as maximum,

formational ages for the onset of strike-slip motion along the DAV fault and CD. The R b Sr biotite and K - A r white-mica cooling age pattern in Figure 4 is interpreted to reflect both fault-related exhumation of basement units north of the DAV fault and the thermal anomaly around the Rieserferner pluton. Most strike-slip activity of the DAV and CD faults occurred at 33-28 Ma (Mfiller et al. 2001; Wagner 2005). Figure 4 shows that whereas the DAV fault truncates the mica cooling age pattern, the CD fault does not offset the pattern of 1 4 - 2 0 M a and 2 0 - 2 5 M a cooling age intervals. This is interpreted to indicate that mylonitic strike-slip faulting lasted until no later than 20Ma. During and/or after that time, basement units north of and along the DAV fault rose through the 300 ~ isotherm. Pseudotachylite locally overprinted by mylonites of the DAV fault (Mancktelow et al. 2001; Mfiller et al. 2001) is interpreted as evidence for local, coseismic decoupling during strike-slip mylonitization. Conjugate, mylonitic-to-cataclastic shear zones that displace the CD and the DAV faults also offset the Rieserferner pluton, but do not seem to affect the 20 Ma mica cooling age isochron (Fig. 4). Most dextral strike-slip movement was taken up along along the brittle Pustertal fault between 20 and 30 Ma ago. This age range is constrained by the coincidence of this fault with 2 8 - 2 9 M a granitoids (Scharbart 1975) and by a 2 2 - 2 0 Ma 4~ crystallization age of pseudotachylite along the fault plane (Mtiller et al. 2001). The Pustertal fault has little, if any, vertical throw (Most 2003). Dextral strike-slip displacement is argued below to amount to some 100kin based on estimates of displacement along the Tonale segment of the Insubric mylonites to the east (see below and Stipp et al. 2004). Figure 5 shows that crustal exhumation migrated from south to north in front of the

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

255

Fig. 4. Age interval contours for the area in Fig. 2 constructed from Rb-Sr biotite and K-Ar white-mica ages of Borsi et al. (1978b) and mica cooling ages contoured in the Tectonometamorphic Age Map of the Alps (Handy & Oberh~insli 2004). Rb-Sr white mica ages are from Mtiller et al. (2000) and Wagner (2004). CD, Cima Dura faults; DAV,

Defereggen-Antholz-Vals fault; KV, Kalkstein-Valarga fault; Pu, Pustertal fault.

southern Alpine indenter. For the time interval of 100 to 30Ma, the KV and DAV faults accommodated, respectively, about 2 - 3 km and 25 km of vertical exhumation. The DAV fault is therefore interpreted to coincide with the site of an older S-dipping, low-angle normal fault that juxtaposed high-pressure amphibolitefacies assemblages in the north (700MPa; St6ckhert 1984) with Alpine-unmetamorphosed rocks in the south during Early Tertiary (Mtiller et al. 2001) or even Late Cretaceous time (Borsi et al. 1973; St6ckhert 1984; discussion in Mancktelow et al. 2001). Oligo-Miocene, N - S shortening during Periadriatic faulting then rotated this extensional fault into its current steep orientation as a sinistral, strike-slip fault. Most exhumation related to Insubric deformation (c. 47km) is concentrated along the southern border of the Tauern window and occurred between 30 and 5 Ma, with only a modest amount of Insubric exhumation along the DAV fault (< 10 km) at 20 Ma, certainly no later than 10 Ma (Fig. 5 and discussion above). The Austroalpine crust north of and along the DAV fault passed through the 280~ isotherm for the viscous-to-brittle transition in quartz (St6ckhert et al. 1999) some 24 to 20 Ma according to zircon fission track ages (Most 2003, and references therein). This late, N-block-up exhumation along the DAV fault involved brittle

reactivation of the greenschist-facies mylonites. Apatite fission track ages in this area indicate no significant differential exhumation, or indeed any fault activity, along this part of the Periadriatic fault system after 10 Ma (StiSckhert et al. 1999). Taken together, these age patterns reveal that most E - W stretching of the orogenic crust in front of the southern Alpine indenter began before exhumation through the viscous-tobrittle transition in quartz-rich rocks. The amount of exhumation decreased from west to east in the area of Figure 4 (Borsi et al. 1978a; Steenken et al. 2002), a trend that is possibly related to the westward decrease of orogennormal shortening (Frisch et al. 1998) associated with overall lateral, eastward extrusion of the Eastern Alps (Ratschbacher et al. 1991b). Most E - W stretching was coeval with the onset of exhumation and rapid cooling of basement rocks in the western Tauern window. E - W stretching and N - S shortening of the Austroalpine basement associated with conjugate, viscous-brittle faulting ceased by 14 Ma at the latest, somewhat before the end of N-block-up exhumation along the DAV fault at 10 Ma. The central segment

The structure of the Tonale segment of the PFS transected by the NFP20-E section (Figs 1 and 6)

256

M.R. HANDY E T AL.

Fig. 5. Age and depth of exhumationacross the Periadriatic fault system along the N-S oriented profile trace in Figure 4. Peak pressures used to estimate maximumdepth of burial are from Strckhert (1984) and Zimmermannet al. (1994). Depths calculated by assuming a geotherm of 30 ~ km-~ and an average crustal density of 2.8 g cm-3. Referencesfor radiometric ages: Cesare (1992, 1994), Most (2003), Steenken et al. (2002), Strckhert (1984), Strckhert et al. (1999), and Zimmermannet al. (1994). is relatively straightforward. There, the Insubric mylonite belt comprises a 1 km wide band of retrograde greenschist-facies mylonite that grades northwards into amphibolite-facies mylonite and post-nappe folds of the so-called 'southern steep belt' (Milnes 1974). This steeply N-dipping belt served as a conduit for the syntectonic ascent of tonalitic melt that fed the Oligocene Bergell pluton (Fig. 6; Rosenberg et al. 1995). South of this pluton, the Insubric mylonites accommodated dextral shear parallel to a subhorizontal stretching lineation (Wiedenbeck 1986; Fisch 1989), whereas within the pluton and to the north, an additional component

of N-side-up reverse faulting ('backthrusting' in Alpine parlance) and exhumation occurred parallel to a steeply N-plunging stretching lineation (Berger et al. 1996). The dip-slip and strike-slip lineations are preserved in separate, though broadly coeval, northern and southern parts of the Insubric mylonite belt (Schmid et al. 1987). The mylonite belt was active from 35 Ma to 20-15 Ma, as constrained by crosscutting relations with Tertiary intrusives (Bergell pluton-Rosenberg et al. 1995, Oberli et al. 2004; Adamello pluton-Stipp et al. 2004), the successive closure of mineral isotopic systems in the Lepontine dome (Hurford 1986; Villa & von Blankenburg 1992) and the age and uplift rate of Bergell granitoid components in the southern Alpine molasse (Wagner et al. 1977; Gunzenhauser 1985; Giger & Hurford 1989). Continued dextral strike-slip motion involved cataclasis along the brittle Tonale fault (Fig. 6) and its continuation to the west, the Centovalli fault (marked Ce in Fig. 7). The exhumation accommodated by the Insubric mylonite belt ranges from 5 km in the Adamello region (Stipp et al. 2004) to a maximum value of 2 0 - 2 5 k m at the eastern margin of the Lepontine dome transected by the NFP20-E profile (Fig. 1). Most, if not all, of this exhumation occurred after 28 Ma, when temperatures reached the solidus of the Bergell tonalite at a depth of about 25 km (Oberli et al. 2004). Exhumation and cooling rates were highest from 23 to 19 Ma (2.2 mm/a) before slowing to current rates after 19 Ma (0.4 mm/a, Hurford 1986). North of the Insubric mylonite belt, stretching parallel to east-west trending fold axes and mineral lineation affected the folded base of the Bergell pluton during and after intrusion (Davidson et al. 1996). Orogen-parallel, E - W extensional exhumation of the Tertiary nappe pile from beneath the Late Cretaceous, Austroalpine nappes was accommodated initially by the Oligocene Turba extensional fault (Nievergelt et al. 1996) and then by the 25-18 Ma Forcola extensional fault (Meyre et al. 1998) (Fig. 6b). In the western part of the southern steep belt, the age interval of rapid cooling (18-15 Ma) is somewhat younger than in the eastern part (30-25Ma, Hurford 1986), suggesting that rapid exhumation migrated from east to west along the E - W striking part of the Insubric mylonite belt (Schmid et al. 1989). Extensional exhumation of the western Lepontine dome initiated as early as 35-30 Ma with the onset of dextral transpressional shearing of the Early Tertiary nappe pile (Steck & Hunziker 1994) and continued from 25 to 11 Ma during topSW, orogen-parallel mylonitic faulting localized

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

257

Fig. 6. The central segment of the Periadriatic fault system: (a) geologic map; (b) cross-section with up-plunge projection into the N-S profile line marked A-A' in Figure 6a (modified from Davidson et al. 1996). Dotted line indicates Tonal brittle fault overprinting the Insubric mylonite belt.

along the Simplon extensional fault (Si in Figs 1 and 7; Mancktelow 1992). The w e s t e r n s e g m e n t Structure and kinematics. At the western end of the PFS, the Insubric mylonite belt (locally named the Canavese Line or Canavese fault, Ca in Fig. 1) narrows progressively from a width of about 1 km to only several metres as it

follows the arcuate border between the basement rocks of the Sesia Zone and the southern Alpine indenter. The surface exposure of the indenter in this area is the Ivrea Zone, whose pre-Alpine mafic and ultramafic rocks can be traced downward in the ECOR-CROP and NFP20-W transects to a SE-dipping wedge of upper mantle rock, the Ivrea Body, that is deeply embedded within the retro-wedge of the Alpine orogen (Nicolas et al. 1991; Schmid & Kissling

258

M.R. HANDY ETAL.

2000). The Sesia Zone, located on the western side of the Insubric mylonite belt (Se in Fig. 1), is part of the Tertiary Alpine orogenic wedge that includes the Tethyan oceanic units and underlying European basement units (Fig. 7). The Canavese fault dips moderately to steeply to the NW, and comprises two strands of mylonite (Fig. 7): one in the west that accommodated top-SE to - E backthrusting (plots 5, 7), and a smaller strand in the east with oblique-normal dextral shear parallel to stretching lineations (plot 8). These strands correspond, respectively, to mylonitic belts 2 and 1 of Schmid et al. (1987, 1989). In the vicinity of the Tertiary Biella pluton (Fig. 7), the western strand bends to the west and enters the Sesia Zone, where it apparently follows the northern limit of eclogitebearing rocks. The eastern strand grades to the SW into cataclasites, which are themselves overprinted by cataclasites of the Cremosina fault (Cr in Fig. 7). Yet further to the SW, outside of the map area in Figure 7, the Sesia-Ivrea contact is marked by Late Cretaceous to Early Tertiary, greenschist-facies mylonite (Zingg & Hunziker 1990) and one, possibly two generations of Oligocene or younger cataclasite (Schmid et al. 1989). These mylonites are beyond the scope of this paper, as they are related to pre-Insubric extensional exhumation of high-pressure rocks in the Sesia Zone. As the Canavese fault narrows to the SW, several retrograde greenschist-facies, mylonitic zones fan out westward from the Sesia-Ivrea border and overprint older mylonitic foliations within the Alpine orogenic retro-wedge (Fig. 7). These Insubric splays are steeply dipping, fold and shear zones of up to several hundred metres thickness (Fig. 8) whose mylonitic foliation contains a mostly subhorizontal stretching lineation. The shear sense parallel to this lineation is predominantly dextral, although sinistral shear sense is found within some zones (Fig. 7). In one of these splays, the dextral shear sense is indicated by the left-stepping direction of mylonitic shear zones that merge along strike with en-6chelon, upright, tight to isoclinal, non-cylindrical folds with subhorizontal to moderately plunging axes (Fig. 7, plots 2 and 6; Insubric fold zone in Fig. 8). In other locations, the shear zones skirt narrow synforms that contain klippen of pre-Alpine rocks belonging to the Upper Unit of the Sesia Zone (Figs 7 and 8, 'Seconda Zona Dioritica Kinzigitica' or II ZDK of Gosso et al. 1979). Along the sides of some of these klippen, the mylonitic foliation dips moderately and the stretching lineation rotates into a direction consistent with

shear upward and outwards from the synformal cores (e.g. plot 6 in Fig. 7). The southern steep belt, including prominent S- to SE-vergent post-nappe folds like the Vanzone antiform (Va in Fig. 7), gradually disappears from NE to SW as it trends away from the arcuate Sesia-Ivrea border and acquires a progressively shallower southeastward dip in the western part of Figure 7. A dextral band of retrograde, amphibolite- to greenschist-facies mylonite follows the contact of LiguroPiemont and Brianconnais units marking the narrow southern limb of the Vanzone antiform (plots 3 and 4 in Fig. 7). In the axial trace of this antiform at the SW end of the Monte Rosa nappe, the mylonitic foliation dips moderately to the SSW (Fig. 7) with a SW-dipping stretching lineation (Steck & Hunziker 1994) This indicates N E - S W extension parallel both to the SW-plunging axis of the Vanzone antiform and to N E - S W trending L-tectonite fabrics in the core of this antiform (Reinhardt 1966). Going to the NE along the steep belt in Figure 7, this same mylonitic band steepens and partly overprints the retrograde greenschist-facies mylonitic foliation of the Gressoney fault (Fig. 8). The Gressoney fault (plot 1 in Fig. 7) has been interpreted as a mid- to late Eocene, lowangle normal fault that accommodated top-SE shear (Wheeler & Butler 1993) prior to Insubric deformation. Age of faulting and exhumation. The coincidence of Insubric mylonitic splays emanating westward from the southern Alpine indenter with the disappearance to the southwest of both the Insubric mylonite belt and the southern steep belt suggests a kinematic link between Tertiary shortening, exhumation and N E SW-directed, orogen-parallel extension in the Western Alps. Establishing this link requires accurate dating of the minerals defining the mylonitic foliations in Figure 7. The tectono-metamorphic age map in Figure 9 shows the distribution of selected R b - S r and K - A r white-mica ages in the literature for areas where the main foliation has a clear relative age (based on cross-cutting relationships) and an unambiguous shear sense (based on kinematic indicators). Only samples for which there was no discernable evidence for inherited effects of pre-Alpine or high-pressure metamorphism are plotted in this figure. Where dated samples were lacking, we extrapolated and interpolated age results parallel to the strike of foliations that were structurally and kinematically continuous. The resulting age map allows one to distinguish mylonites of the

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

259

o "~

e-.

.=.=~

.o;~o= o

~ o

~r

. ,...w

~'=~

~r,,w ~ ~

260

M.R. HANDY ET AL.

Fig. 8. Cross-sectionparallel to the trace in Figure 7.

southern steep belt and the Canavese fault from Early Tertiary, extensional faults (Ometto and Gressoney faults in Fig. 9). The Ometto and Gressoney extensional faults contributed to the exhumation of Late Cretaceous and Early Tertiary eclogites, respectively, in the Sesia Zone and in the underlying ophiolitic and European basement units. Dextral shearing under amphibolite- to greenschist-facies conditions in the southem steep belt is inferred to syn-date Vanzone backfolding and related NE-SW, axial extension, based on the structural relations above. Steck & Hunziker (1994) argue that dextral sheafing occurred some 35-25 Ma. The upper age limit for this mylonitization is given by 25-29 Ma U-Pb, R b - S r and 4~ ages of syn- to post-mylonitic dykes (locations in Fig. 9; Romer et al. 1996; SchS.rer et al. 1996). Vanzone backfolding affected the pattern of 38 Ma R b - S r white mica and 2 4 - 2 0 M a R b - S r biotite cooling ages (age contours in Fig. 9; Steck & Hunziker 1994), but not the pattern of zircon fission track ages at 15 Ma, the time when temperatures fell below 225-240 ~ (Hunziker et al. 1992). Therefore, this backfolding probably occurred between 35 and 20 Ma. W-side-up backthrusting and dextral strikeslip in greenschist-facies, Insubric mylonites of the Canavese fault are dated at 26-19 Ma by K - A r formational white-mica ages (Zingg & Hunziker 1990). This age range may extend back to 30 Ma based on the occurrence of

synmylonitic dykes with close mineralogical and geochemical affinity to the Bergell granitoids (Reinhardt 1966; Schmid et al. 1987). Within the limits of age resolution, the Insubric mylonitic splays in the northeastern part of the Sesia Zone fall in the same age range; certainly, these splays post-date the cessation of Gressoney extensional faulting by 36 Ma (Reddy et al. 1999 and references therein), but precede the transition from viscous to brittle deformation in quartz-rich rocks before cooling to the 225240 ~ closure temperature for fission tracks in zircon at about 15 Ma (Hunziker et al. 1992). The difference in closing temperatures of the R b - S r white mica and FT zircon systems indicates that the temperature in the NE Sesia Zone and adjacent Penninic units decreased by some 200-300 ~ between 35 and 15 Ma. Thus, exhumation due to Insubric folding and backthrusting in this area amounted to about 7 - 1 4 km for assumed geothermal gradients of 20-30 ~ Insubric exhumation is much more modest (<3 km) in the vicinity of the Biella pluton (Fig. 7) and to the southwest, based on the occurrence of 45-30 Ma zircon FT ages and 15-8 Ma apatite FT ages in that area (Hurford et al. 1991). This is corroborated by several observations suggesting that the southwestern part of the Sesia Zone was at or near the surface already prior to, or during, Oligocene time: (1) the Oligocene Biella pluton is a shallow intrusion (<5 km, Rosenberg 2004) and truncates the foliation of rocks with Late Cretaceous, R b - S r white

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

261

Fig. 9. Tectonometamorphicage map of the westemend of the Periadriaticfault systembased on geochronologicaldata of Duch~neet al. (1997), Hunziker(1974), Hunzikere t al. (1992), Inger e t al. (1996), Reddy et al. (1996, 2003), Romer et al. (1996), Rubatto et al. (1999),Ruffetet aI. (1997), Sch~er et al. (1996), Wemmer(1991), Zingg& Hunziker(1990). Contours of Rb-Sr cooling ages from Steck & Hunziker(1994). Outlineof tectonic units corresponds to Figure 7.

mica (Fig. 6) and Early Tertiary, R b - S r biotite cooling ages (Hunziker et al. 1992); (2) the Oligocene volcanoclastic cover of the Sesia Zone includes boulders of exhumed, Late Cretaceous (Duch~ne et al. 1997; Rubatto et al. 1999) eclogite (Bianchi & Dal Piaz 1963). Subsequent rotation of the Sesia-Ivrea contact into its current steep attitude along the Boccioleto antiform (Bo in Fig. 7; Steck & Hunziker 1994) cannot have involved significant postOligocene vertical displacement, because Oligocene igneous rocks in the internal part of the Sesia Zone (e.g. Biella pluton, volcanoclastics, Fig. 7) are exposed at about the same altitude as Tertiary intrusives in the Ivrea Zone (Miagliano tonalite; Carraro & Ferrara 1968).

Synthesis of structural, kinematic and age r e l a t i o n s o f I n s u b r i c f a u l t i n g a l o n g the P F S

The structure, kinematics and age ranges of mylonitic faults making up the PFS are

summarized in Figures 10 and 11 from results and literature cited above and in the figure captions. In Figure 10, only mylonitic foliations and stretching lineations from Oligo-Miocene segments of the PFS are included in the projections. The distribution of these fabrics with respect to the orientation of the indenter surface lends insight into how strain was partitioned on the orogenic scale in front of the southern Alpine indenter. The black arrows indicate the Oligo-Miocene motion of the southern Alpine indenter relative to a stable Europe, as determined by palaeomagnetic studies (e.g. Dercourt et al. 1986; Dewey et al. 1989). Post-nappe folding and NE-SW-directed axial extension in the Western Alps (Fig. 10) was broadly coeval with N-side-up exhumation and dextral strike-slip along the PFS at the east-west trending, northern surface of the southern Alpine indenter in the Central Alps (Fig. 11). At the NW corner of the indenter, mylonite belts that accommodated dextral strike-slip splay westwards, away from the

262

M.R. HANDY E T AL.

angular indenter surface and into the arc of the Westem Alps (Fig. 10). Strike-slip and dip-slip motion along and in front of the indenter were strongly partitioned, as indicated by the point-maxima stretching lineation patterns in Insubric mylonites of the westem Canavese segment and in the northeastern part of the Sesia Zone (Fig. 10). The great-circle lineation pattern along the eastern Canavese segment (Fig. 10) is interpreted to reflect changes in this partitioning during dextral, transpressional flow around the arcuate tip of the indenter. The Tonale and Pustertal segments of the PFS are estimated to have accommodated about 100km of dextral strike-slip motion parallel to subhorizontal stretching lineations (To in Fig. 10; Schmid & Kissling 2000). Total values of dextral displacement range from l l 0 k m (Lacassin 1989) to 150km (Laubscher 1991), but post-Oligocene movement probably amounted to no more than 30km (MUller et al. 2001). Most, if not all, of this limited, postOligocene movement was probably accommodated by the aforementioned brittle faults that overprint the Insubric mylonite belt. The PFS, including the Insubric mylonites, also accommodated a poorly constrained amount of Tertiary N - S shortening to the north of the southern Alpine indenter. Regarded perpendicular to the current E - W trend of the PFS, this shortening varies along strike from a maximum postOligocene value of about 113 km in the TRANSALP transect (Frisch et al. 1998) to a post-Eocene estimate of 63-71 km parallel to the NFP20-E transect (Schmid & Kissling 2000). Along the Tonale segment, part of this N - S shortening was taken up by S-directed backthrusting parallel to steep, N-plunging stretching lineations in the Insubric mylonite belt (Fig. 10). Orogen-parallel extension along low-angle normal faults beginning in Oligocene time exhumed progressively deeper units of the folded, Early Tertiary nappe pile. The Simplon and Forcola faults are largely responsible for oblique, N E - S W directed extensional unroofing of the western Lepontine dome by late Miocene time (lineation poles and open arrows in Fig. 10, ages in Fig. 11). The Turba fault at the eastern end of the Lepontine dome is the structurally highest of these extensional faults and was active already in early Oligocene time (Fig. 11). Along the Tonale segment of the Insubric mylonite belt, dextral strike-slip motion outlasted Nside-up exhumation and continued under brittle conditions into Mio-Pliocene time (Fig. 11). The structural and age relations above support the idea that coeval N W - S E shortening and

NE-SW-directed orogen-parallel extension in the intemal basement units of the Western Alps absorbed a considerable amount of the previously cited 100 km of dextral strike-slip displacement during Oligocene-Miocene transpression along the Tonale segment of the PFS (arrows in Fig. 10). The magnitude of this N E - S W extension obviously varies with the age and amount of displacement along the different segments of the PFS, especially the Insubric mylonite belt, the Simplon fault and the Penninic frontal thrust (Fig. 10). If the estimated 30 km of post-Oligocene motion along the PFS cited by Miiller et al. (2001) were taken up by 25-11 Ma Simplon extensional faulting (3236 km displacement estimates of Grasemann & Mancktelow 1993; Wawrzyniec et aL 2001), then the remaining 70km of pre-Miocene strike-slip displacement along the PFS were probably accommodated by 35-20Ma, N W SE shortening and N E - S W orogen-parallel extension of the Monte Rosa basement unit that structurally overlies the Simplon extensional fault. This estimate is close to the maximum 80 km of N E - S W extension proposed by Steck and Hunziker (1994) for the entire Tertiary nappe edifice, including the Simplon fault. Greater amounts of post-Oligocene shortening along the Penninic frontal thrust in the Western Alps, for example, 60 km proposed by Schmid & Kissling (2000), lead to more modest estimates of basement shortening and related N E - S W extension at the western termination of the Insubric mylonite belt. In the Eastern Alps, sinistral shear of the orogenic crust north of the DAV fault at the southwestern margin of the Tauern window (Fig. 10) was transitional in space and time to doming of the western Tauern window (Fig. 11). This doming involved extension parallel to colinear ENE-WSW trending axes of km-scale, upright folds and colinear stretching lineations (Lammerer & Weger 1996), as indicated by the single-girdle pattern of the foliation poles and the point-maximum of the lineations (Fig. 10). These patterns are consistent with sinistral transpression with a strong component of ENE-WSW stretching. This was coeval with N-side-up exhumation, which largely ended along the DAV fault by 20 Ma but migrated northward into the western Tauern window (Fig. 11). Doming was then transitional to ENE-WSW extension parallel to the stretching lineation in mylonite of the overlying Brenner extensional fault (Figs 10 and 11). The Mrlltal fault at the southeastern margin of the Tauern window shows a similar, perhaps somewhat younger evolution to the DAV fault, but with

D E C O U P L I N G AND STRAIN PARTITIONING IN C O N T I N E N T A L L I T H O S P H E R E

263

o o

o ~.~.~

~o

~.o~

~~ i ~

~o

~o o ~ o ~

~

..~ = ~ < ~

~.~<

0 &-~= ~.~ ~~ ~ ~

~ ~-~

~.~ ~---~

"~.~

~'~

~

9~

~

~ .~ ~.~

.~,

~.~.~ ~

~.~ ~ ~ , ~chc~ o'n

264

M.R. HANDY Western

ET AL.

+ Central Alps

Eastern Alps

0

0 PLIO' & PLEI-

"

": " 9 ":':"

CENE

-ME --,

,

~.~

!::: :

i / : .....................~

i::: L::

TO

,

~PRI ua

Fig. 11. Age ranges of the Periadriatic fault segments. Stippled, brittle deformation; solid, mylonitic deformation; striped pattern, poor age constraints on deformation. Sources: Amphibolite- to greenschist-facies, dextral strike-slip in the southern steep belt (Steck & Hunziker 1994; Romer et al. 1996; SchSxeret al. 1996). Backfolding (this paper; Steck & Hunziker 1994; Hurford et al. 1991; Hunziker et al. 1992), strike-slip faulting in the northeastern Sesia Zone (this paper), Simplon extensional fault (Mancktelow 1992; Grasemann & Mancktelow 1993; Steck & Hunziker 1994), Canavese segment of the Insubric mylonite belt (this paper; Zingg & Hunziker 1990), Tonale segment of the Insubric mylonite belt (Berger et al. 1996; Schmid et al. 1989; Schmid et al. 1996; Stipp et al. 2004), Forcola extensional fault (Meyre et al. 1998), Turba extensional fault (Nievergelt et al. 1996; Schmid et al. 1996), Giudicarie fault (Mtiller et al. 2001; Stipp et al. 2004), Brenner extensional fault (Frisch et al. 2000; Ftigenschuh et al. 1997), Western Tauern folding (compilations of Ftigenschuh et al. 1997; Most 2003), Defereggen-Antholz-Vals fault (this paper; Borsi et al. 1978a, b; Mtiller et al. 2001), Pustertal fault (Mtiller et aL 2001; Stipp et al. 2004), M611tal fault (Kurz & Neubauer 1996; Reddy et al. 1993; Inger & Cliff 1994), Eastern Tauern folding (Cliff et al. 1985; Reddy et al. 1993), Katschberg extensional fault (Dunkl et al. 2003; Frisch et al. 2000); time-scale according to Harland et al. (1989).

the opposite shear sense in map view (Figs 10 and 11). It accommodated N-side-up exhumation combined with dextral strike-slip shear that offsets the contact between Austroalpine and Liguro-Piemont units (Fig. 1). The M611tal fault trends parallel to km-scale, upright folds (Kurz & Neubauer 1996) that effected doming of the eastern Tauern window at 2 6 - 1 2 Ma (Cliff e t al. 1985; Reddy e t al. 1993). Like the large folds north of the DAV fault, these folds have constrictional fabrics in their cores (Cliff e t a l . 1971) and define great-circle foliation girdles and subhorizontal point-maxima lineation patterns (Fig. 10). Such fabrics are diagnostic of N W - S E directed extension parallel to the fold axes, leading Kurz & Neubauer (1996) to interpret the M611tal fault as a stretching fault in the sense of Means (1989). The continued activity of the M611tal fault under

brittle conditions to recent times is attributed to Plio-Pleistocene dextral motion along the Karawanken fault (Ka in Fig. 1; Polinski & Eisbacher 1991). The Katschberg extensional fault at the eastern end of the Tauern window accommodated top-SE motion parallel to a stretching lineation (Fig. 10; Genser & Neubauer 1989), opposite to the top-WSW extension along the broadly coeval Brenner fault at the western end of the Tauern window. The similar structural evolution and conjugate geometry of the faults and related folds at opposite ends of the Tauern window suggests that they effected orogen-parallel extension and exhumation of the Early Tertiary nappe pile in the Tauern window (open arrows in Fig. 10). Estimates of this extension range from a minimum of about 37 km (displacement estimates of 20 km and 17 km, respectively, for the

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE Brenner and Katschberg extensional faults; Behrmann 1988; Selverstone 1988; Genser & Neubauer 1989) to a maximum of 160km (Frisch et al. 1998), depending on the structural markers used and assumptions made about the amount of erosion. Tertiary E - W stretching and exhumation of the Tauern thermal dome in the Eastern Alps has been attributed to simultaneous northward motion of the southern Alpine indenter and lithospheric thinning beneath the Pannonian basin during eastward, hinge roll-back subduction in the Carpathians (e.g. Royden & Burchfiel 1989; Ratschbacher et al. 1991a). As pointed out in the next section, N - S to N N E - S S W directed shortening and E - W stretching of the Eastern Alps during oblique indentation implies decoupling along the dextral Pustertal fault marking the front of the southern Alpine indenter (Pu in Fig. 10). The brittle Pustertal fault was active in OligoMiocene times before being offset by sinistral strike-slip motion along the Giudicarie fault (Fig. 11). The kinematic history of the Giudicarie fault is controversial, with some authors interpreting it as an inherited restraining bend in the PFS with only minor sinistral displacement (Mancktelow et al. 2001; Viola et al. 2001) and others arguing that it was originally E - W oriented and rotated into its current NNE-SSW trend while accommodating some 70 km of sinistral offset from 20 to 12Ma (Fig. 11; Schmid et al. 1996; Stipp et al. 2004). We favour the latter scenario, because the amount of Tertiary, E - W shortening in the vicinity of the Giudicarie fault is too small (_<30km) to accommodate the aforementioned 70 km of pre-Miocene, dextral strike-slip displacement along the Tonale segment of the Insubric mylonite belt. Also, sinistral motion of the Giudicarie fault is kinematically linked to Miocene, Sdirected thin-skinned folding and thrusting beneath the southern Alpine Molasse (Pieri & Groppi 1981; Laubscher 1996; Prosser 1998; Schtnborn 1999). Therefore, prior to 20 Ma, the Insubric mylonite belt in the Western and Central Alps and the Pustertal fault in the Eastern Alps were part of a continuous, E - W trending lineament along the northern edge of the southern Alpine indenter. Lateral stretching and eastward extrusion of the Eastern Alps in front of the southern Alpine indenter continued well into late Miocene time along a conjugate array of brittle, strike-slip faults (In, Setup, Mm, La, M6, Pu in Fig. 1; e.g. Ratschbacher et al. 199Ib). The synthesis above reveals that for all parts of the PFS, lateral extension and exhumation of basement nappes along and in front of the indenter surface involved constrictional folding

265

and mylonitic faulting. This occurred when pre-existing foliations rotated into a steep orientation, thereby facilitating strike-slip and dip-slip displacements. Folding and strike-slip faulting were transitional in space and time to low-angle normal faulting at the viscous-tobrittle transition. The orientations of these mylonitic faults and folds with respect to the indenter surface reflect significant differences in intracrustal decoupling patterns, as discussed in the next section.

Intracrustal decoupling along the Periadriatic fault system The correlation of strain gradients, displacements and ages of fault activity along the various mylonitic segments of the PFS in the previous section now allow us to identify decoupling zones as defined at the outset of this paper. These zones are related to the subsurface structure of the Alpine orogen, depicted in a block diagram of the Western Alps in Figure 12. This diagram was constructed by combining the geometry and kinematics of the PFS mapped at the surface with subsurface images of the lower crust and upper mantle along the ECORS-CROP and NFP20-W geophysical transects in Figure 1. It reveals that decoupling within the crust occurred along two types of heterogeneity: (1) inherited lithological and/or thermal discontinuities (coloured red), especially the surface of the southern Alpine orogenic indenter; and (2) strain-induced heterogeneities (coloured yellow) within the orogenic crust in front of the indenter. D e c o u p l i n g at the i n d e n t e r s u r f a c e

The most prominent decoupling zone was the interface between the cold, brittle pre-Alpine rocks of the orogenic indenter and the warm, amphibolite-facies rocks of the Early Tertiary nappe pile. This surface was the locus of exhumation and strike-slip motion within the retro-wedge of the Alpine orogen. The prevalence of mylonite along this surface indicates that decoupling involved largely aseismic, viscous creep. Pseudotachylite zones in the Insubric mylonite (e.g. DAV fault; Mancktelow et al. 2001) and in the adjacent southern Alpine rocks (Ivrea Zone; Zingg et al. 1990) are modest in extent (cm to m length) and are therefore interpreted to represent only short periods of embrittlement and limited coseismic slip along the indenter surface. The concentration of strain along the indenter surface is most pronounced along the Tonale and

266

M.R. HANDY

E T AL.

Fig. 12. Block diagram of mylonitic segments of the Periadriatic fault system in the Central and Western Alps as viewed towards the southeast. To simplify the view, basement units are stripped away. Decoupling localized along inherited, lithological and thermal heterogeneities (red surfaces) or along strain-induced, rotated anisotropies (yellow surfaces). Base of the lower crust adopted from subsurface interpretations of Schmid et al. (1996) and Schmid & Kissling (2000) for the NFP20-E, NFP20-W and ECOR-CROP transects (locations in Fig. 1). Canavese fault segments, where the Insubric mylonites follow the angular, leading edge of the Ivrea Body. None of the faults extend into the mantle, but root at, or somewhere above, the top of the lower crust (Fig. 12). The lower crust of both the lower (European) and upper (Apulian) plates consists of mafic rocks, based on its high seismic velocity and seismic reflectivity (Valasek & Mfiller 1997). Beneath the Central Alps, the Apulian lower crust forms an indented wedge that is replaced to the west by an imbricated wedge of lower European crust (Schmid & Kissling 2000). This wedgelike geometry suggests that the lower crust and upper mantle were very strong during OligoMiocene indentation and confirms the view previously expressed for the individual transects that the base of the viscous, quartz-rich crust was a first-order decoupling horizon in the Alps. Upper mantle structures like the Ivrea Body are by no means restricted to the Western Alps, having been imaged beneath the trace of several active transpressive faults, including the Alpine fault zone on the Tasmanian-Pacific plate boundary (e.g. Oliver & Coggon 1979),

the San Andreas fault on the Pacific-North American plate boundary (Teyssier & Tikoff 1998), and the E1 Pilar-Coche Fault Zone on the Caribbean-South American plate boundary (Teyssier e t al. 2002). The rheological contrasts associated with disparate MOHO depths across such active transpressional boundaries are accentuated by high, lateral thermal gradients, especially where the faults along these boundaries accommodate crustal exhumation or serve as conduits for the rapid ascent of melt. In the case of the PFS, progressive cooling and hardening of the viscously deforming hangingwall adjacent to the southern Alpine indenter is a possible explanation for the migration of exhumation into the retro-wedge of the Alpine orogen. No pseudotachylite has been found so far along Insubric mylonitic zones away from the surface of the southern Alpine indenter. The concentration of pseudotachlyite at the indenter surface may reflect the relative strain rate and/or width of these mylonitic fault zones when they were active under amphibolite- to greenschist-facies conditions. Chester (1995) argues that seismic instabilities are suppressed in wider zones

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE because local flow perturbations are damped, thereby favouring velocity-strengthening behaviour. This is especially true of mylonitic faults, where the dominant deformation mechanisms (dislocation creep, diffusion creep) are ratedependent and the rocks strengthen with increasing strain rate (Scholz 1990). At comparable displacement rates, narrower faults deform at higher strain rates, which favour a switch to velocityweakening, in some cases leading to seismic instability (Chester 1995). The actual width of the Insubric mylonite belt at any given time during its activity, especially during pseudotachylite formation, is not known. It was probably less than the present 1-2 km thickness, especially during the latter stages of exhumation when lateral thermal gradients across the fault zone led to a strong localization of strain in ultramylonites derived from relatively cold protoliths (>200 ~ of the southern Alpine indenter (Zingg et al. 1990). Decoupling within the orogenic crust The mylonitic faults in the orogenic crust in front of the southern Alpine indenter represent strain-induced mechanical heterogeneities (yellow surfaces in Fig. 12). Decoupling below the viscous-to-brittle transition was due to the progressive folding and rotation of older faults and foliations into orientations which became conveniently oriented for mylonitic shearing (Fig. 10). Both the Gressoney extensional fault in the Western Alps and the extensional predecessor to the DAV fault in the Eastern Alps were steepened, then reactivated as mylonitic strike-slip faults. The progressive reorientation and reactivation of such pre-existing anistropies represent the strain-dependent growth of new mechanical phases. Insight into the kinematic and dynamic significance of such anisotropies for decoupling within the orogenic crust is gained from the orientational distribution of foliation (Sm) poles and stretching lineations (Ls) for mylonitic faults of the PFS in Figure 13. The pole diagrams in this figure were constructed by reorienting the contoured foliation and stretching lineation poles for the faults in Figure 10 into a common coordinate system with the indenter surface along the PFS as a vertical reference plane. In Figure 13, this plane is the equatorial line of all the pole diagrams. The local orientation of the indenter surface obviously varies along strike of the PFS for each of the plots in Figure 13 (inset map), and was constructed by extrapolating foliation measurements at the surface down to subsurface geophysical images of the

267

indenter front along the transects in Figure 1 (see Schmid & Kissling 2000; TRANSALP Working Group 2002; for summaries of these images). The lines emanating from the periphery of the Ls pole plots show the average foliation trends and the thin black arrows indicate the motion parallel to the stretching lineations. Large, open arrows indicate the inferred, bulk extensional direction, as discussed below. Most of the pole patterns are very symmetrical and comprise great-circle maxima or point maxima that lie on great circles. In the Eastern Alps, these great circles define cross-girdle patterns and subhorizontal lineation maxima that reflect conjugate, non-coaxial shearing along the sinistral DAV and dextral Mrlltal faults (Fig. 13a), and constrictional folding in both the eastern and western parts of the Tauern window (Fig. 13b). Foliation and lineation maxima for the Brenner and Katschberg extensional faults (Fig. 13c) also outline a composite pattern that is nearly orthogonal to the indenter surface. Conjugate or duplex shearing on these faults and folds therefore extended the orogenic crust subparallel to the indenter surface in the Eastern Alps, as indicated by the open arrows in Figures 13a to c. In contrast, the foliation and lineation maxima along the Tonale segment of the PFS in the Central Alps (Fig. 13d) reflect partitioning of strike-slip and dip-slip motions on a single shearing plane subparallel to the indenter surface. The maxima associated with the Forcola and Simplon extensional faults (Fig. 13e) are oblique to the indenter surface, consistent with oblique, N E - S W directed extension during dextral transpression in the Central Alps. At the western end of the PFS, the Insubric mylonites (Fig. 13f) define a broad foliation maxima with two lineation maxima at a high mutual angle. This indicates strong partitioning of N E - S W lateral extension and SE-directed backthrusting, again subparallel to the indenter surface. The core of the Vanzone antiform defines a single-girdle foliation pattern with a point maxima indicative of coaxial stretching oblique to the indenter surface as defined by the front of the Ivrea geophysical body (Fig. 13g). The pole patterns described above reveal that oblique, NW motion of the southern Alpine indenter in Oligo-Miocene time was accommodated differently along strike of the Alpine orogen: NNE-directed shortening and E - W extension in the Eastern Alps contrasts with N W - S E shortening and N E - S W extension in the Western and Central Alps. A possible explanation is that the southern Alpine indenter was segmented, with separate eastern and western

268

M . R . HANDY ETAL.

Fig. 13. Symmetry and kinematics of mylonitic shearing along the Periadriatic Fault System. Equal area projections contain contoured poles to mylonitic foliation (Sin) and stretching lineation (Ls) with respect to the leading edge of the southern Alpine indenter oriented vertically east-west. Contours in projections taken from Figure 10. (a) DAV and M611tal mylonitic faults and related folds; (b) large uptight folds in the eastern and western parts of the Tauern window, Eastern Alps; (c) Brenner and Katschberg extensional faults, Eastern Alps; (d) Tonale segment of the Insubric mylonite belt near the Bergell pluton, Central Alps; (e) Forcola and Simplon extensional faults, Central Alps; (f) western Canavese segment of the Insubric mylonite belt, Insubric mylonites in the Sesia Zone; (g) core of the Vanzone antiform, Western Alps. Angles used for rotating the data in these plots with respect to the indenter surfaces are shown in the lower-tight diagram. Boxed abbreviations for the indenter surfaces: Iv, front of Ivrea geopyhsical body; Ap, front of Apulia crustal wedge; Pu, Pustertal fault and its projection to depth. Lines and arrows emanating from Ls pole plots show the average foliation trend and motion sense on the mylonitic fault surfaces. Open arrows indicate bulk extension direction (see text). Abbreviations for fault surfaces as in Figure 1.

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE blocks rotating clockwise along a mutual boundary defined by the sinistral Giudicarie fault. However, sinistral motion along this fault occurred after 20 Ma (Fig. 11), possibly in conjunction with a reversal of subduction direction beneath the Eastern Alps (Schmid et al. 2003, 2004). Furthermore, rigid block rotations within the main edifice of the Alps are only plausible for late middle to late Miocene time (Laubscher 1996) when the orogenic crust had cooled sufficiently to enable rigid, brittle behaviour of basement units currently exposed at the surface. In the absence of any evidence for segmentation and differential block rotations within the indenter in Oligocene time, the varied orientations of the bulk strain field during OligoMiocene Periadriatic faulting are attributed to decoupling around the variably oriented and shaped parts of the southern Alpine indenter. In the Eastern Alps, the straight, northern edge of the southern Alpine indenter was decoupled from the orogenic crust along the dextral Pustertal fault. The lateral, E - W displacement component of oblique convergence was therefore taken up by brittle faulting along the indenter surface, leaving conjugate faulting and folding to accommodate the normal, NNE-displacement component within the exhuming, orogenic crust in front of the indenter. In contrast, the orogenic crust in the Western Alps was constrained to flow up (to the SE) and around (to the NE) the angular, steeply NW-dipping edge of the geophysical Ivrea body (Fig. 12). The exhumational component of shear was concentrated at the indenter surface in the Insubric mylonite belt. Once the crust flowed around this restraining bend, it underwent N E - S W extension in the Central Alps, leading to the Miocene exhumation of the Lepontine thermal dome. The coeval activity of mylonitic faults along the PFS from 35 to 15 Ma together with the affinity of bulk stretching directions accommodated by these faults to the local orientation of the southern Alpine indenter surface suggest that the leading edge of this indenter was the macroscopic shearing plane for Insubric deformation. Furthermore, the fact that foliation and lineation pole patterns are highly symmetrical and define slip systems that are systematically disposed with respect to the indenter surface support the idea that the preexisting schistosities rotated into orientations which maximized the component of simple shearing within these slip systems. The kinematic significance of rotating anisotropies by folding and shearing in a general non-coaxial strain field is therefore that stable or semi-stable orientations of shear zones are attained for which the bulk stretching direction

269

corresponds to an eigenvector of flow, that is, a direction along which the rate of flow is the greatest and toward which all passive markers within a rock tend to rotate. Similar to our analysis above, Cobbold & Gapais (1986) compared networks of outcrop-scale shear zones to ideal fibre-type slip systems in which slip occurs along inextensible lines parallel to the shearing direction. At least three independent slip systems are needed to maintain strain compatibility during bulk plane-strain deformation, and the minimum number of systems increases to five for a three-dimensional strain field (von Mises 1928). Accordingly, the orientational distribution pattern of slip-planes and -lines with respect to the bulk shearing plane reflects the symmetry of the bulk strain field (Gapais & Cobbold 1987; Gapais et al. 1987). We note that combined shear-and-fold zones are not restricted to the Alps, having been identified in other deeply eroded transpressional fault zones (e.g. Ebert & Hasui 1998) which appear to have similar fabric orientational distributions as those described for the PFS.

Significance for crustal strength and force transmission through the lithosphere Decoupling related to strain partitioning in the viscous continental crust presents an interesting dynamic problem, because intracrustal strain partitioning was previously thought to occur in much shallower levels, at the interface of the viscous crust with the brittle, upper crust (e.g. Richard & Cobbold 1989; Molnar 1992). Molnar (1992) attributed this to the apparent inability of a horizontal viscous layer to transmit shear stresses to an attached brittle layer above. Yet, we have seen that partitioning of strikeslip and dip-slip shearing along the PFS occurred below the brittle-to-viscous transition. The solution to this discrepancy between Nature and experiment lies with the fact that crustal rocks are mechanically anisotropic, primarily due to the strong mechanical anisotropy of sheet silicates, especially micas in schistose rock. The mechanical effect of rotating a weak, pre-existing anisotropy into parallelism with the shearing plane is two-fold. The resolved shear stress in directions parallel to this anisotropy (i.e. the weak direction) will increase, while the volume proportion of deforming rock that is optimally oriented to maximize simple shear in these directions also increases. Folding or shearing an anisotropic rock therefore weakens it, an effect known in rock mechanics as 'rotation-weakening' (Cobbold 1977)or 'foliation-weakening'

270

M.R. HANDY ET AL.

(Jordan 1988). By analogy with experimental and theoretical work on bimineralic aggregates (Handy 1990), the weakening associated with the formation of new faults depends on the volume proportion of mylonitic faults oriented parallel to the shearing plane, and on the relative strength of the fault rocks and the less-deformed country rock. Thus, even small volumes of fault rock suffice to weaken the entire crust significantly (Handy 1994). For example, if the strike-slip faults make up 5-10% of the crustal, and the strength ratio of less-deformed rock to fault rock is taken to be 3:1 (a conservative estimate), then polyphase flow laws predict a strength drop of about 20%. Although greater strength drops are expected for larger strength contrasts, this estimate may be a maximum for energetic reasons (Handy 1994). The notion that plate-scale fault systems rotate into stable end-orientations to accommodate high shear strain is similar to the 'easy glide' concept first proposed by E. Schmid (1928) to explain the rotation of crystallographic slip systems toward stable orientations of high resolved shear stress during crystal-plastic deformation of polycrystalline aggregates. An important difference, however, is that the potential glide direction (i.e. the Burger's vector) in an intracrystalline slip system is limited by the crystallography and orientation of the host grain in a polycrystalline aggregate (e.g. Taylor 1938), whereas the shearing direction in a mylonitic shear zone is determined primarily by the stress field in the vicinity of the fault array. This local stress field is predictably related to the far-field, plate-scale stress field if the configuration, orientation and length-scale of the fault array have become strain-invariant. An important implication of the above for plate tectonics is that weakening and intracrustal decoupling in the viscous crust do not preclude the vertical transmission of normal and shear stresses through the lithosphere. Although Molnar's (1992) idea that shear stresses cannot be transmitted upward across horizontal viscous layers is correct if these layers have very low viscosity and are both homogeneous and isotropic, it evidently does not apply to heterogeneous, anisotropic media like folded, foliated basement rocks (Cobbold 1977). Decoupling along and across networked folds and mylonitic faults in the viscous continental crust ensures that deformation between the stronger subcontinental mantle and overlying brittle crust remains compatible, even at high shear strains where the incompatibility in strain between these layers with disparate strengths is potentially very large. The increased volume

proportion of weak, folded and sheared rock in the viscous crust decreases the bulk, horizontal shear strength of the viscous crust, and therefore also of the lithosphere. Yet, the strain-dependent rotation of weak anisotropies towards the macroscopic shearing plane also increases the volume proportion of weak rock oriented at high angles to the greatest principal stress direction. Thus, the plate forces transmitted across an intracrustal decoupling zone are inferred to remain constant or even to increase at constant bulk strain rate. The same principle of straindependent, strength reduction at constant or even increasing plate-scale force may also apply to relatively weak, decoupling layers in the mantle. To speak of the continental crust beneath the brittle-viscous transition as a horizontal 'attachment zone' (Tikoff et al. 2002) is misleading, because the vertical transmission of plate forces that potentially facilitates strikeslip faulting along plate boundaries in the upper crust (e.g. Li & Rice 1987) may also involve detachment of the upper crust from the upper mantle, that is, decoupling in the sense defined at the outset of this paper. Therefore, a more appropriate term for a weak, heterogeneously sheared, anisotropic layer that transmits stress through the lithosphere is 'accommodation zone'.

A generic model of decoupling and strain partitioning The block diagram in Figure 14 is a generic model of intracrustal decoupling in orogenic crust based on the Periadriatic fault system. Decoupling is concentrated along the leading edge of the orogenic indenter, where faulting initiates along pre-existing lithological boundaries, in this case between the Early Mesozoic, rifted margin of the southern Alps and the previously subducted and exhumed nappes in the retro-wedge of the Tertiary Alpine orogen. The rheological contrast at the indenter surface is accentuated by lateral thermal gradients during mylonitic thrusting of the hot, orogenic retro-wedge onto the cold indenter. Decreasing temperature in the retro-wedge enhances strain localization and increases the strain rate, leading to viscous instabilities and the generation of pseudotachylite near the indenter surface. Within the warm retro-wedge of the orogen, older foliations are folded and steepen to form subvertical mylonitic faults and constrictional folds that accommodate stretching parallel to subparallel to the indenter surface. The steep mylonitic faults are potential conduits for the rapid ascent of mantle-derived melts, some of

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE

271

Fig. 14. Genetic model of decoupling zones related to strain partitioning along the Periadriatic fault system. which (like the Rieserferner and Bergell plutons) are emplaced in the cores of antiforms below the brittle-to-viscous transition. These mylonitic faults are optimally oriented to accommodate simple shearing and are predicted to weaken the orogenic crust by up to 20%. Further away from the indenter a n d / o r in shallower crustal levels, orogen-parallel extension is a c c o m m o dated by low-angle normal faults, which detach the exhuming ductile crust from the brittle, upper crust. W e have shown that complex patterns of folding and faulting during transpressional deformation in orogenic crust can be interpreted in terms of decoupling in different crustal levels. The future characterization of such structures in other deeply eroded, ancient fault systems may help us to understand how stress and strain are transferred across active continental fault systems. We thank the reviewers, D. Marquer and L. Ratschbacher, for helpful comments, and our colleagues (K. Hammerschmidt, R. Oberh~insli, G. Gosso, I. Spalla, M. Zucali, R. Bousquet, S. M. Schmid, H. Sttinitz) for spirited discussions. This work was facilitated by the mapping of several Diploma students in parts of the areas described

above: M. Albertz, J. Giese, R. H/iusler, M. Janitschke, S. Ltidtke, C. Mrbus, I. Schindelwig and B. Sperber. Our work was financed largely by the Deutsche Forschungs Gemeinschaft (DFG) in the form of two 2.5year grants (Ha 2403/3, Ha 2403/5) and by one-year assistantships to J.B., M.K. and R.W. from the Freie Universit~it Berlin.

References ARGAND, E. 1916. Sur l'arc des Alpes Occidentales. Eclogae geologicae Helvetiae, 14, 145-191. AXEN, G. J., SELVERSTONE, J. 8~ WAWRZYNIEC, T. 2001. High-temperature embrittlement of extensional Alpine mylonite zones in the mid-crustal ductilebrittle transition. Journal of Geophysics, 106, 437-448. BEHRMANN, J. H. 1988. Crustal-scale extension in a convergent orogen. The Sterzing-Steinach mylonite zone in the Eastern Alps. Geodynamica Acta, 2, 63-73. BERGER, A., ROSENBERG, C. ~; SCHMID, S. M. 1996. Ascent, emplacement and exhumation of the Bergell pluton within the Southern Steep Belt of the Central Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 76, 357-382. BIANCHI, A. & DAL PIAZ, G. V. 1963. Gli inclusi di 'micaschisti eclogitici' della zona Sesia nella formazione porfiritica permiana della zone del

272

M.R. HANDY ET AL.

Canavese fra Biella ed Oropa; charatteristiche ed eta dei fenomeni metamorfici. Giornale Geologico Serie 2a, 31, 39-76. BOILLOT, G., BESLIER, M. O. & COMAS, M. C. 1995. Nature, structure and evolution of the oceancontinent boundary: the lesson of the Western Galicia Margin (Spain). In: BANDA,E., TORNI~, M. TALWANI, M. (eds) Rifted Ocean-Continent Boundaries. Kluwer, Dordrecht, 219- 229. BORSI, S., DEL MORO, A. & SASSI, F. P. 1973. Metamorphic evolution of the Austridic rocks to the south of the Tauern Window (Eastern Alps): radiometric and geopetrological data. Memoria Societg~ Geologica Italiana, 12, 549-571. BORSI, S., DEL MORO, A., SASSI, F. P., ZANFERRARI, A. & ZIRPOLI, G. 1978a. New geopetrological and radiometric data on the Alpine history of the Austridic continental margin south of the Tauern Window (Eastern Alps). Memorie Scienze Geologica Universit~ Padova, 32, 1-17. BORSI, S., DEL MORO, A., SASSI, F. P. & ZIRPOLI, G. 1978b. On the age of the Vedrette di Ries (Rieserferner) massif and its geodynamic significance. Geologische Rundschau, 68, 41-60. BRACE, W. F. & KOHLSTEDT, D. L. 1980. Limits on lithospheric stress imposed by laboratory experiments. Journal of Geophysical Research, 85 (BI 1), 6248-6252. CARRARO, F. & FERRARA, G. 1968. Alpine 'Tonalite' at Magliano, Biella (Zona Dioritico-Kinzigitica). A preliminary note. Schweizerische Mineralogische und Petrographische Mitteilungen, 48, 75- 80. CESARE, B. 1992. Metamorfismo di contatto di rocce pelitiche nell' aureola di Vedrette di Ries (Alpi Orientali Italia). Atti Ticinensi di Scienze della Terra, 35, 1-7. CESARE, B. 1994. Hercynite as a product of staurolite decomposition in the contact aureole of Vedrette di RiPs, eastern Alps, Italy. Contributions to Mineralogy and Petrology, 116, 239-246. CHESTER, F. M. 1995. A rheologic model for wet crust applied to strike-slip faults. Journal of Geophysical Research, 100 (B7), 13033-13044. CLIFF, R. A., DROOP, G. T. R. & REX, D. C. 1985. Alpine metamorphism in the south-east Tauern Window, Austria: 2. Rates of heating, cooling and uplift. Journal of Metamorphic Geology, 3, 403-415. CLIFF, R. A., NORRIS, R. J., OXBURGH, E. R. & WRIGHT, R. C. 1971. Structural, metamorphic and geochronological studies in the Riesseck and Southern Ankogel Groups, the Eastern Alps. Jahrbuch des Geologischen Bundesanstalts Wien, 114, 121-272. COBBOLD, P. R. 1977. Description and origin of banded deformation structures. II. Rheology and the growth of banded perturbations. Canadian Journal of Earth Sciences, 14, 2510-2523. COBBOLD, P. R. & GAPAIS, D. 1986. Slip-system domains. I. Plane-strain kinematics of arrays of coherent bands with twinned fibre orientations. Tectonophysics, 131, 113-132. DAVIDSON, C., ROSENBERG, C. & SCHMID, S. M. 1996. Synmagmatic folding of the base of the

Bergell pluton, Central Alps. Tectonophysics, 265, 213-238. DAVIS, G. A. & LISTER G. S. 1988. Detachment faulting in continental extension: perspectives from the Southwestern US Cordillera. In: CLARK, S. P. JR., BURCHFIEL, B. C. & SUPPE J. (eds) Processes in Continental Lithospheric Deformation. Geological Society of America, Special Papers, 218, 133-159. DEE MORO, A., PARDINI, G. C., QUERCIOLI, C., VILLA, I. M. & CALLEGARI, E. 1983. Rb/Sr and K/Ar chronology of Adamello granitoids, southern Alps. Memorie della Societh Geologica Italiana, 26, 285-299. DERCOURT, J., ZONENSHAIN, L. P. et al. 1986. Geological evolution of the Tethys belt from the Atlantic to the Pamirs since the Lias. Tectonophysics, 123, 241-315. DEWEY, J. F., HELMAN, M. L., TURCO, E., HUTTON, D. H. W. & KNOTT, S. D. 1989. Kinematics of the Western Mediterranean. In: COWARD, M. P., DIETRICH, D. & PARK, R. G. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 265-283. DUCHI~NE, S., BLICHERT-TOFT, J., LUAIS, B., TI~LOUK, P., LARDEAUX,J.-M. & ALBARI~DE,F. 1997. The Lu-Hf dating of garnets and the ages of the Alpine high-pressure metamorphism. Nature, 387, 586-588. DUNKL, I., FRISCH, W. & GRUNDMANN G. 2003. Zircon fission track thermochronology of the southeastern part of the Tauern Window and the adjacent Austroalpine margin, Eastern Alps. Eclogae geologicae Helvetiae, 96, 209-217. EBERT, H. D. & HASUI, Y. 1998. Transpressional tectonics and strain partitioning during oblique collision between three plates in the Precambian of southeast Brazil. In: HOLDSWORTH, R. E., STRACHAN, R. A. & DEWEY, J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 231-252. ESCHER, A. & BEAUMONT,C. 1997. Formation, burial and exhumation of basement nappes at crustal scale: a geometric model based on the Western Swiss-Italian Alps. Journal of Structural Geology, 19 (7), 955-974. FISCH, H. R. 1989. Zur Kinematik der stidlichen Steilzone der Zentralalpen, E von Bellinzona. Schweizerische Mineralogische und Petrographische Mitteilungen, 69, 377-393. FOSSEN, H. & TIKOFF, B. 1998. Extended models of transpression and transtension, and application to tectonic settings. In: HOLDSWORTH, R. E., STRACHAN, R. A. & DEWEY, J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 15-33. FRISCH, W., KUHLEMANN,J., DUNKL, I. & BROGEL, A. 1998. Palinspastic reconstruction and topographic evolution of the Eastern Alps during late Tertiary tectonic extrusion. Tectonophysics, 297, 1-15.

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE FRISCH, W., DUNKL, I. & KUHLEMANN,J. 2000. Postcollisional orogen-parallel large-scale extension in the Eastern Alps. Tectonophysics, 327, 239-265. FROITZHEIM, N., SCHMID, S. M. & FREY, M. 1996. Mesozoic paleogeography and the timing of eclogite-facies metamorphism in the Alps: a working hypothesis. Eclogae geologicae Helvetiae, 89 (1), 81-110. FOGENSEHUH, B., SEWARD, D. & MANCKTELOW, N. 1997. Exhumation in a convergent orogen: the western Tauern window. Terra Nova, 9, 213-217. GAPAIS, D. & COBBOLD, P. R. 1987. Slip system domains. 2. Kinematic aspects of fabric development in polycrystalline aggregates. Tectonophysics, 138, 289-309. GAPAIS, D., BALE, P., CHOUKROUNE, P., COBBOLD, P. R., MAHJOUB, Y. & MARQUER, D. 1987. Bulk kinematics from shear zone patterns: Some field examples. Journal of Structural Geology, 9 (5/6), 635 -646. GENSER, J. & NEUBAUER,F. 1989. Low-angle normal faults at the eastern margin of the Tauern window (Eastern Alps). Mitteilungen der Osterreichischen Geologischen Gesellschafi, 81, 233-243. GIGER, M. & HURFORD, A. J. 1989. Tertiary intrusives of the Central Alps: Their Tertiary uplift, erosion, redeposition and burial in the south Alpine foreland. Eclogae geologicae Helvetiae, 82, 857-866. Gosso, G., DAL PIAZ, G. V., PIOVANO, V. & POLINO, R. 1979. High pressure emplacement of early-Alpine Nappes, post-nappe deformations and structural levels (internal northwestern Alps). Memorie Istituto di Mineralogia e Petrografia dell' Universit~ di Padova, 32, 5-15. GRASEMANN, B. & MANCKTELOW, N. 1993. Twodimensional thermal modelling of normal faulting: the Simplon Fault Zone, Central Alps, Switzerland. Tectonophysics, 225, 155-165. GUNZENHAUSER, B. 1985. Zur Sedimentologie und Pal/iogeographie der oligo-miocaenen Gonfolite Lombarda zwischen Lago Maggiore und der Brianza (Stidtessin, Lombardei). Beitriige zur geologischen Karte der Schweiz, NF, 159. HANDY, M. R. 1990. The solid-state flow of polymineralic rocks. Journal of Geophysical Research, 96 (B6), 8647-8661. HANDY, M. R. 1994. The energetics of steady state heterogeneous shear in mylonitic rock. Materials Science and Engineering, A175, 261-272. HANDY, M. R. & ZINGG, A. 1991. The tectonic and rheologic evolution of an attenuated cross section of the continental crust: Ivrea crustal section, southern Alps, northwestern Italy and southern Switzerland. Geological Society of America Bulletin, 103, 236-253. HANDY, M. R. & OBERHANSLI,R. 2004. Metamorphic structure of the Alps: tectonometamorphic age map. In: OBERH~NSLI, R. (ed) Explanatory Notes to the Map of Metamorphic Structures of the Alps. Mitteilungen der 13sterreichischen Mineralogischen Gesellschaft, 49, 201-226. HARLAND, W. B., ARMSTRONG, A. V., Cox, L. E., CRAIG, A. G. & SMITH, D. G. 1989. A Geologic

273

Time Scale. Cambridge University Press, New York. HOBBS, B. E., ORD, A. & TEYSSIER, C. 1986. Earthquakes in the ductile regime? Pure and Applied Geophysics, 124 (1/2), 309-336. HOINKES, G., KOLLER, F., RANTITSCH, D., DACHS, E., HOCK, V., NEUBAUER, F. & SCHUSTER, R. 1999. Alpine metamorphism of the Eastern Alps. Schweizerische Mineralogische und Petrographische Mitteilungen, 79 (1), 155-182. HOLT, W. E. 2000. Correlated crust and mantle strain fields in Tibet. Geology, 28 (1), 67-70. HUNZIKER, J. C. 1974. Rb-Sr and K-Ar age determinations and the Alpine tectonic history of the Western Alps. Memorie Geologica della Universit~ di Padova, 31, 1-54. HUNZIKER, J. C., DESMONS,J. & HURFORD,A. J. 1992. Thirty-two years of geochronological work in the Central and Western Alps: a review on seven maps. Mgmoires de Ggologie (Lausanne), 13, 1-59. HURFORD, A. J. 1986. Cooling and uplift patterns in the Lepontine Alps, South Central Switzerland and an age of vertical movement on the Insubric fault line. Contributions to Mineralogy and Petrology, 92, 413-427. HURFORD, A. J., HUNZIKER, J. C. & STOCKERT, B. 1991. Constraints on the late thermotectonic evolution of the western Alps: Evidence for episodic rapid uplift. Tectonics, 10 (4), 758-769. INGER, S. & CLIFF, e. A. 1994. Timing of metamorphism in the Tauern Window, Eastern Alps: Rb-Sr ages and fabric formation. Journal of Metamorphic Geology, 12, 695-707. INGER, S., RAMSBOTHAM, W., CLIFF, R. A. & REX, D. C. 1996. Metamorphic evolution of the Sesia Lanzo Zone, Western Alps: time constraints from multi-system geochronology. Contributions to Mineralogy and Petrology, 126, 152-168. JORDAN, P. 1988. The rheology of polymineralic rocks - an approach. Geologische Rundschau, 77 (1), 285-294. KLEINSCHRODT, R. 1987. Quarzkorngefiigeanalyse im Altkristallin stidlich des westlichen Tauernfensters (Siidtirol/Italien). Erlanger Geologische Abhandlungen, 114, 1-82. KURZ, W. & NEUBAUER, F. 1996. Deformation partitioning during updonfing of the Sonnblick area in the Tauern Window (Eastern Alps, Austria). Journal of Structural Geology, 18 (11), 1327-1343. LAEASSIN, R. 1989. Plate scale kinematics and compatibility of crustal shear zones in the Alps. In: COWARD, M. P., DIETRICH, D., PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 339-352. LAMMERER, B. & WEGER, M. 1996. Footwall uplift in an orogenic wedge: the Tauern Window in the Eastern Alps of Europe. Tectonophysics, 285, 213-230. LAUBSCHER, H. 1991. The arc of the Western Alps today. Eclogae Geologica Helvetiae, 84 (3), 631-659. LAUBSCHER, H. 1996. Shallow and deep rotations in the Miocene Alps. Tectonics, 15, 1022-1035.

274

M.R. HANDY ETAL.

LI, V. C. & RICE, J. R. 1987. Crustal deformation in Great California earthquake cycles. Journal of Geophysical Research, 92 (B 11), 11533-11551. MADDOCK, R. 1983. Melt orogin of pseudotachylites demonstrated by textures. Geology, 11, 105-108. MAGLOUGHLIN, J. F. & SPRAY, J. G. 1992. Frictional melting processes and products in geological materials: introduction and discussion. Tectonophysics, 204, 197-204. MANCKTELOW, N. S. 1990. The Simplon Fault Zone. Beitriige zur Geologie der Schweiz, 163 (Neue Folge), 1-74. MANCKTELOW, N. S. 1992. Neogene lateral extension during convergence in the Central Alps: evidence from interrelated faulting and backfolding around the Simplonpass (Switzerland). Tectonophysics, 215, 295-317. MANCKTELOW, N. S., STOCKLI, D. F. et al. 2001. The DAV and Periadriatic fault systems in the Eastern Alps south of the Tauern window. In: FRISCn, W. & MANCKTELOW, N. S. (eds) Developments in Alpine Geodynamics. International Journal of Earth Sciences, Springer, Berlin, 90 (3), 593-622. MCNULTY, B. A. 1995. Peudotachylyte generated in the semi-brittle and brittle regimes, Bench Canyon shear zone, central Sierra Nevada. Journal of Structural Geology, 17 (11), 1507-1521. MEANS, W. D. 1989. Stretching faults. Geology, 17, 893 -896. MEYRE, C., MARQUER,D., SCHMID, S. M. & CIANCALEONI, L. 1998. Syn-orogenic extension along the Forcola fault: Correlation of Alpine deformations in the Tambo and Adula nappes (Eastern Penninic Alps). Eclogae Geologicae Helvetiae, 91,409-420. MILNES, G. 1974. Structure of the pennine zone (central Alps): A new working hypothesis. Geological Society of America Bulletin, 85, 1725-1732. MOONEY, W. D. & GINZBURG, A. 1986. Seismic measurements of the internal properties of fault zones. Pure and Applied Geophysics, 124 (1/2), 141-157. MOLNAR, P. 1992. Brace-Goetze strength profiles, the partitioning of strike-slip and thrust faulting at zones of oblique convergence, and the stressheat flow paradox of the San Andreas Fault. In: EVANS, B. & WONG, T. F. (eds) Fault Mechanics and Transport Properties of Rocks: A Festschrift in Honor of W. F. Brace. International Geophysics Series, 51, 435-460. MOST, P. 2003. Late Alpine cooling histories of tectonic blocks along the central part of the TRANSALP traverse (lnntal-Gardatal) : constraints from geochronology. PhD thesis, Universit~it Tfibingen. MOLLER, W., MANCKTELOW N. S. & MEIER, M. (2000) Rb-Sr microchrons of synkinematic mica in mylonites: an example from the DAV fault of the Eastern Alps. Earth and Planetary Science Letters, 180, 385-397. MULLER, W., PROSSER, G., MANCKTELOW, N. S., VILLA, I. M., KELLEY, S. P., VIOLA, G. & OBERLI, F. 2001. Geochronological constraints on the evolution of the Periadriatic Fault System

(Alps). In: FRISCH, W. & MANCKTELOW, N. S. (eds) Developments in Alpine Geodynamics. International Journal of Earth Sciences, Springer, Berlin, 90 (3), 623-653. NICOLAS, A., POLINO, R., HIRN, A., NICOLICH, R. & the Ecors-Crop working group. 1991. Ecors-Crop traverse and deep structure of the western Alps. A synthesis. In: ROURE, F., HEITZMANN, P. & POLINO, R. (eds) Deep Structure of the Alps. Mrmoire Socirt6 Grologique de France, Paris, 156; Mrmoire Socirt6 Grologique de Suisse, Zarich, 1; Volume speciale de Societh Geologica Italiana, Roma, 1, 15-28. NIEVERGELT, P., LINIGER, M., FROITZHEIM, N. & FERREIRO MAHLMANN, R. 1996. Early to mid Tertiary crustal extension in the Central Alps: The Turba Mylonite Zone (Eastern Switzerland). Tectonics, 15 (2), 329-340. OBERLI, F., MEIER, M., BERGER, A., ROSENBERG, C. & GIERI~, R. 2004. U - T h - P b and 230Th/238U disequilibrium isotope systematics: precise accessory mineral chronology and melt evolution tracing in the Alpine Bergell Intrusion. Geochimica et Cosmochimica Acta, 68, 2543-2560. OLIVER, G. J. H. & COGGON, J. H. 1979. Crustal structure of Fiordland, New Zealand. Tectonophysics, 54, 253-292. P1ERI, M. & GROPPI, G. 1981. Subsurface geological structure of the Po Plain, Italy. Progettto Finalizzato Geodinamica, CNR Publication, 414, 113321, Rome, Italy. POL1NSKI, R. K. & EISBACHER, G. H. 1991. Deformation partitioning during polyphase oblique convergence in the Karawanken Mountains, southeastern Alps. Journal of Structural Geology, 14 (10), 1203-1213. PROSSER, G. 1998. Strike-slip movements and thrusting along a transpressive fault zone: the Northern Giudicarie line (Insubric line, northern Italy). Tectonics, 17, 921-937. RANALLI, G. & MURPHY, D. C. 1987. Rheological stratification of the lithosphere. Tectonophysics, 132 (4), 281-296. RATSCHBACHER, L., MERLE, O., DAVY, P. & COBBOLD, P. 1991a. Lateral extrusion in the eastern Alps, Part 1: boundary conditions and experiments scaled for gravity. Tectonics, 10 (2), 245-256. RATSCHBACHER, L., FRISCI-I, W., LINZER, H. G. & MERLE, O. 1991b. Lateral extrusion in the eastern Alps, Part 2: structural analysis. Tectonics, 10 (2), 257-271. REDDY, S. M., CLIFF, R. A. & EAST, R. 1993. Thermal history of the Sonnblick Dome, south-east Tauern Window, Austria: implications for heterogeneous uplift within the Pennine basement. Geologische Rundschau, 82, 667-675. REDDY, S. M., KELLY, S. P. & WHEELER, J. 1996. A 40Ar/39Ar laser probe study of micas from the Sesia Zone, Italian Alps: implications for metamorphic and deformation histories. Journal of Metamorphic Geology, 14, 493-508. REDDY, S. M., WHEELER, J. • CLIFF, R. A. 1999. The geometry and timing of orogenic extension: an

DECOUPLING AND STRAIN PARTITIONING IN CONTINENTAL LITHOSPHERE example from the Western Italian Alps. Journal of Metamorphic Geology, 17, 573-589. REDDY, S. M., WHEELER, J. et al. 2003. Kinematic reworking and exhumation within the convergent Alpine Orogen. Tectonophysics, 365, 77-102. REINHARDT, B. 1966. Geologie und Petrographie der Monte-Rosa Zone, der Sesia Zone und des Canavese im Gebiet zwischen Valle d'Ossola und Val Loana (Prov. di Novara, Italien). Schweizerische Mineralogische und Petrographische Mitteilungen, 46 (2), 553-678. RICHARD, P. & COBBOLD, P. R. 1989. Structures en fleurs positives et d6crochements crustaux: mod61isation analogique et interpr6tation m6canique. Comptes Rendus Acaddmie des Sciences, Paris S6rie II, 308, 553-560. ROMER, R. L., SCHARER,U. & STECK, A. 1996. Alpine and pre-Alpine magmatism in the root-zone of the western Central Alps. Contributions to Mineralogy and Petrology, 123, 138-158. ROSENBERG, C. L. 2004. Shear zones and magma ascent: A model based on a review of the Tertiary magmatism in the Alps. Tectonics, 23, TC3002, doi: 10.1029/2003TC001526. ROSENBERG, C., BERGER, A. & SCHMID, S. M. 1995. Observations from the floor of a granitoid pluton: Inferences on the driving force of final emplacement. Geology, 23 (5), 443-446. ROYDEN, L. & BURCHFIEL, B. C. 1989. Are systematic variations in thrust belt style related to plate boundary processes? (The western Alps versus the Carpathians). Tectonics, 8 (1), 51-61. RUBATTO, D., GEBAUER,D. & COMPAGNONI,R. 1999. Dating of eclogite-facies zircons: the age of Alpine metamorphism in the Sesia-Lanzo Zone (Western Alps). Earth and Planetary Science Letters, 167, 141-158. RUFFET, G., GRUAU, G., BALLI~VRE, M., FERAUD, G. & PHILLIPOT, P. 1997. Rb-Sr and 40Ar-39Ar laser probe dating of high-pressure phengites from the Sesia Zone (Western Alps): underscoring of excess argon and new age constraints on the high-pressure metamorphism. Chemical Geology, 141, 1-18. SCHARBART, S. 1975. Radiometrische Altersdaten von Intrusivgesteinen im Raum Eisenkappel (Karawanken, K~irnten). Verhandlungen der Geologischen Bundesanstalt, Wien, 4, 301-304. SCH,~,RER, U., COSCA, M., STECK, A. & HUNZIKER, J. C. 1996. Termination of major ductile strike-slip shear and differential cooling along the Insubric line (Central Alps): U-Pb, Rb-Sr and 40Ar/39Ar ages of cross-cutting pegmatites. Earth and Planetary Science Letters, 142, 331 - 351. SCHMID, E. 1928. Zn-normal stress law. In: Proceedings of the International Congress on Applied Mechanics, Delft. SCHMID, S. M. & HANDY, M. R. 1991. Towards a genetic classification of fault rocks: geological usage and tectonophysical implications. In: Hs(2, K. J., MACKENZIE, J. & MULLER, D. (eds) Controversies in Modern Geology. Academic Press, London, 339-361.

275

SCHMID, S. M., ZIN~G, A. & HANDY, M. R. 1987. The kinematics of movements along the Insubric Line and the emplacement of the Ivrea Zone. Tectonophysics, 135, 47-66. SCHMID, S. M., AEBLI, H. R., HELLER, F. & ZINGG, A. 1989. The role of the Periadriatic Line in the tectonic evolution of the Alps. In: COWARD, M. P., DIETRICH, D. & PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 153-171. SCHMID, S. M., PFIFFNER, O. A., FROITZHEIM, N., SCHONBORN, G. & KISSLING, E. 1996. Geophysical-geological transect and tectonic evolution of the Swiss-Italian Alps. Tectonics, 15 (5), 1036-1064. SCHMID, S. M. & KISSLING, E. 2000. The arc of the western Alps in the light of geophysical data on deep crustal structure. Tectonics, 19 (1), 62-85. SCHM1D, S. M., FUGENSCHUH, B. & LIPPITSCH, R. 2003: The Western Alps-Eastern Alps transition: tectonics and deep structure. Extended Abstracts TRANSALP conference, Trieste. Memorie di Scienze Geologische, 54, 257-260. SCHMID, S. M., FOGENSCHUH, B., KISSLING, E. & SCHUSTER, R. 2004. TRANSMED Transects IV, V and VI: Three lithospheric transects across the Alps and their forelands. In: CAVAZZA, W., ROURE, F., SPAKMAN, W., STAMPFLI, G. M. & ZIEGLER, P. A. (eds) The TRANSMED Atlas: the Mediterranean Region from Crust to Mantle. Springer Verlag, Berlin. SCHMID, S. M., FUGENSCHUH, B., KISSLING, E. & SCHUSTER, R. 2004. Tectonic map and overall structure of the Alpine orogen. Eclogae Geologicae Helvetiae, 97, 93-117. SCHOLZ, C. n. 1990. The Mechanics of Earthquakes and Faulting. Cambridge University Press, Cambridge. SCHONBORN, G. 1999. Balancing cross sections with kinematic constraints: the Dolomites (northern Italy). Tectonics, 18, 527-545. SELVERSTONE, J. 1988. Evidence for East-West crustal extension in the eastern Alps: implications for the unroofing history of the Tauern Window. Tectonics, 7, 87-105. SENARCLENS-GRANCY, W. 1965. Zur Grundgebirgsund Quart~irgeologie der Deferegger Alpen und ihrer Umgebung. Verhandlungen der Geologischen Bundesanstalt, Sonderheft G, 246-255. SIBSON, R. H. 1975. Generation of pseudotachylite by ancient seismic faulting. Geophysical Journal of the Royal Astronomical Society, 43, 775-784. SIBSON, R. H. 1977. Fault rocks and fault mechanisms. Journal of the Geological Society, London, 133, 191-213. STECK, A. & HUNZIKER, J. 1994. The tertiary structural and thermal evolution of the Central Alps - compressional and extensional structures in an orogenic belt. Tectonophysics, 238, 229-254. STEENKEN, A., SIEGESMUND, S. & HEINRICHS, T. 2002. The emplacement of the Rieserferner Pluton (Eastern Alps, Tyrol): constraints from field observations, magnetic fabrics and microstructures. Journal of Structural Geology, 22, 1855-1873.

276

M.R. HANDY ET AL.

STIPP, M., F~)GENSCHUH,B., GROMET,L. P., ST(INITZ, H. & SCHMID, S. M. 2004. Contemporaneous plutonism and strike-slip faulting along the easternmost segment of the Insubric line: the Tonale fault zone north of the Adamello pluton (Italian Alps). Tectonics, 23, doi: 10.1029/ 2003TC001515. STOCKHERT,B. 1984. K - A r determinations on muscovites and phengites from deformed pegmatites, and the minimum age of Old Alpine deformation in the Austridic basement to the South of the western Tauern Window (Ahrn valley, Southern Tyrol, Eastern Alps). Neues Jahrbuch der Mineralogie, Abhandlungen, 150 (2), 103-120. STOCKHERT, B., BRIX, M. R., KLEINSCHRODT, R., HURFORD, A. J. & WIRTH, R. 1999. Thermochronometry and microstructures of quartz - a comparison with experimental flow laws and predictions on the temperature of the brittle-plastic transition. Journal of Structural Geology, 21, 351-369. SYLVESTER,A. G. 1988. Strike-slip faults. Geological Society of America Bulletin, 100, 1666-1703. TAYLOR, G. I. 1938. Plastic strain in metals. Journal of the Institute of Metals, 62, 307-324. TEYSSIER, C. & TIKOFF, B. 1998. Strike-slip partitioned transpression of the San Andreas fault system: a lithospheric-scale approach. In: HOLDSWORTH, R. E., STRACHAN, R. A. & DEWEY, J. F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 143-158. TEYSSIER, CH., TIKOFF, B. & WEBER, J. 2002. Attachment Between Brittle and Ductile Crust at Wrenching Plate Boundaries. EGS Stephan Mueller Special Publication Series, 1, 119-144. TIKOFF, B., TEYSSIER, CH. & WATERS, CH. 2002. Clutch Tectonics and the Partial Attachment of Lithospheric Layers. EGU Stephan Mueller Special Publication Series, 1, 57-73. TRANSALP Working Group. 2002. First deep seismic reflection images of the Eastern Alps reveal giant crustal wedges and transcrustal ramps. Geophysical Research Letters, 29 (10), 1029/2002GL014911. TSE, S. T. & RICE, J. R. 1986. Crustal earthquake instability in relation to the depth variation of frictional slip properties. Journal of Geophysical Research, 91 (B9), 9452-9472. VALASEK, P. (~ MI3LLER, ST. 1997. A 3D crustal model of the Swiss Alps based on an integrated interpretation of seismic refraction and NRP 20 seismic reflection data. In: PFIFFNER, O. A., LEHNER, P., HEITZMANN, P., MUELLER, ST. & STECK, A. (eds) Results of NRP 20: Deep Structure of the Swiss Alias. Birkh~iuser Verlag, Basel, Boston, Berlin, 305-325. VAUCHEZ, A. (~ TOMMASI, A. 2003. Wrench faults down to the asthenosphere: geological and geophysical evidence and thermo-mechanical effects. In: STORTI, F., HOLDSWORTH, R. E. & SALVANI, F. (eds) Intraplate Strike-Slip Deformation Belts. Geological Society, London, Special Publications, 210, 15-34.

VILLA, I. & VON BLANCKENBURG,F. 1992. A hornblende Ar-Ar age traverse of the Bregaglia tonalite (SE Central Alps). Schweizerische Mineralogische und Petrographische Mitteilungen, 71, 73-87. VIOLA, G., MANCKTELOW, N. & SEWARD, D. 2001. Late Oligocene evolution of Europe-Adria collision: New structural and geochronological evidence from the Giudicarie fault system (Italian Eastern Alps). Tectonics, 20 (6), 999-1020. VON BLANCKENBURG,F., VILLA, I. M., BAUER, H., MORTEANI, G. & STEIGER, R. H. 1989. Time calibration of a PT-path from the Western Tauem Window, Eastern Alps: the problem of closure temperatures. Contributions to Mineralogy and Petrology, 101, 1- 11. VON MISES, R. 1928. Mechanik der plastischen Form~inderung von Kristallen. Zeitschriftfiir angewandte Mathematik und Mechanik, 8, 161 - 184. WAGNER, G. A., MILLER, D. S. & J,~GER, E. 1977. Fission track ages on apatite of Bergell rocks from central Alps and Bergell boulders in Oligocene sediments. Earth and Planetary Science Letters, 45, 355-360. WAGNER, R. 2005. Tertiary tectonics and magmatism along the TRANSALP-Profile between the Tauern Window and the Periadriatic Line (Eastern Alias). PhD thesis, Freie Universit~it Berlin, Germany. WAWRZYNIEC, T. F., SELVERSTONE,J. 8z AXEN, G. J. 2001. Styles of footwall uplift along the Simplon and Brenner normal fault systems, Central and Eastern Alps. Tectonics, 20, 748-770. WEMMER, K. 1991. K/Ar-Altersdatierungsmrglichkeiten ftir retrograde Deformationsprozesse im sprrden und duktilen Bereich - Beispiele aus der KTB Vorbohrung (Oberpfalz) und dem Bereich der Insubrischen Linie (N-Italien). G~ttinger Arbeiten der Geologie und Paliiontologie, 51, 1-61.

WENK, H.-R. 1978. Are pseudotachylites the product of fracture or fusion? Geology, 6, 507-511. WIEDENBECK, M. 1986. Structural and isotopic age profile across the Insubric Line, Mello, Valtellina, N. Italy. Schweizerische Mineralogische und Petrographische Mitteilungen, 66, 211 - 217. WHEELER, J. & BUTLER, W. H. 1993. Evidence for extension in the western Alpine orogen: the contact between the oceanic Piemonte and overlying continental Sesia units. Earth and Planetary Science Letters, 117, 457-474. ZIMMERMANN,R., HAMMERSCHMIDT,K. • FRANZ, G. 1994. Eocene high-pressure metamorphism in the Pennine units of the Tauern window (Eastern Alps): Evidence from 40Ar-39Ar dating and petrological investigations. Contributions to Mineralogy and Petrology, 117, 175-186. ZINGG, A. & HUNZIKER,J. C. 1990. The age of movements along the Insubric Line West of Locarno (northern Italy and southern Switzerland). Eclogae Geologicae Helvetiae, 83 (3), 629-644. ZINGG, A., HANDY, M. R., HUNZIKER, J. C. & SCHMID, S. M. 1990. Tectonometamorphic history of the Ivrea Zone and its relation to the crustal evolution of the Southern Alps. Tectonophysics, 182, 169-192.

Lithospheric-scale analogue modelling of collision zones with a pre-existing weak zone E R N S T W I L L I N G S H O F E R 1, D I M I T R I O S S O K O U T I S t & J E A N - P I E R R E B U R G 2

1Faculty of Earth and Life Sciences, Vrije Universiteit Amsterdam, De Boelelaan 1085, 1081 HV Amsterdam, The Netherlands (e-mail: ernst, willingshofer@ falw. vu.nl) 2Geologisches Institut, ETH-Ziirich and Universitdt Ziirich, Sonneggstrasse 5, 8092 Ziirich, Switzerland Abstract: Lithospheric-scale analogue experiments have been conducted to investigate the

influence of strength heterogeneities on the distribution and mode of crustal-scale deformation, on the resulting geometry of the deformed area, and on its topographic expression. Strength heterogeneities were incorporated by varying the strength of the crust and upper mantle analogue layers and by implementing a weak plate or part-of-a-plate between two stronger ones. Three (brittle crust/viscous crust/strong viscous upper mantle) and four (brittle crust/viscous crust/brittle upper mantle/strong viscous upper mantle) layer models were confined by a weak silicone layer on one side in order to contain but not oppose lateral extrusion. Experimental results show that relative strength contrasts between converging plates and intervening weak plates control the location and the shape of deformation sites taken as 'collision orogens'. If the contrast is small, internal deformation of the strong plates through fore- and backthrusting occurs early in the deformation history. However, the bulk system is dominated by buckling that nucleates on the weak plate whose antiformal topography is highest; model Moho of the bordering stronger plates is deepest under these conditions. If the contrast is large, deformation remains localized within the weak plate for a larger amount of shortening and develops a root zone below a narrow deformation belt, which coincides with the locus of maximum topography. Implementing a buoyant, low-viscosity layer above the model Moho of the weak plate favours the development of asymmetric model orogens notwithstanding the initial symmetric setup. Once the asymmetry is established strain remains localized in thrust faults and ductile shear zones documenting foreland directed displacement of the model orogen. Such laterally and vertically irregular configurations have applications in continentcontinent collision settings such as the Eastern Alps. First-order mechanical boundary conditions recognized from modelling to be favourable to the post early Oligocene tectonics of the Eastern Alps include: (1) subtle rather than high-strength contrasts between the Adriatic indentor and the strongly deformed region comprising Penninic and Austroalpine units to the north of it; (2) decoupling of Penninic continental upper crust from its substratum to allow for crustal-scale buckling of the Tauern Window; (3) weak mechanical behaviour of the European lower crust during collision to account for its constant thickness along the TRANSALP deep seismic transect; and (4) the direct continuation of the basal detachment underlying the fold and thrust belt in the hangingwall of the European plate with a wide ductile shear zone in the core of the orogen, which separates the European from the Adriatic plate.

The mechanical properties of the continental lithosphere are non-uniform in space and time (Ranalli 1997). This heterogeneity is primarily due to changes in composition and thermal conditions expressed in the theological stratification of the lithosphere with the M o h o being the most important discontinuity (e.g. Ranally & Murphy 1987). Laterally, the rheology of the continental lithosphere may be modified because of

tectonics, leading for example to the separation or collision of continents. Such processes may result in changes of composition (e.g. continental next to oceanic rheology) and lithospheric thicknesses, both in compression as well as extension, and are usually associated with a pronounced thermal perturbation, which influences the strength of the lithosphere transiently. In that way, the thermo-mechanical age of the

From: GAPAIS,D., BRUN, J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and

Tectonics:from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 277-294. 0305-8719/05/$15.00

(~ The Geological Society of London 2005.

278

E. WILLINGSHOFER ET AL.

lithosphere is reset, which emphasizes the strong time-dependence of lithospheric strength (Cloetingh & Burov 1996). Additionally, the lithosphere is affected by faulting and shearing producing a number of metastable rheological discontinuities that are prone to reactivation (Ranalli 2000). Subsequent deformation of the lithosphere will be steered by pre-existing lateral strength variations (e.g. Ziegler et al. 1998) with relative strength differences among deforming minerals (Handy 1990), rock layers (Hudleston & Lan 1993), crustal-scale layers (Gerbault & Willingshofer 2004), or lithospheric plates (e.g. Molnar & Tapponnier 1975) as controlling factor in terms of strain distribution, structure and style of deformation. Our study focuses on this strength contrast across plate boundaries. In particular, we investigate differences in the structural evolution of collision zones, their deep structure, the relationship to higher-level deformations and the resultant topographic expression for conditions of continental convergence as a function of the relative strength contrast of the colliding plates. For this purpose we use a fully mechanical approach, namely lithospheric-scale analogue modelling, which is not restricted by the amount of imposed strain and allows incorporating lateral material transfer. We subsequently discuss implications of our modelling results on aspects of the tectonic evolution of the Eastern Alps in Europe, from where a wealth of surface and subsurface data allow constraining the large-scale geometry of the mountain range as well as its evolution through time.

Experimental concept: why a weak zone? When continents collide, the continental crust thickens through complex deformation processes involving thrusting and folding. Simplified views invoke one plate thrust over the other. As a consequence, the lower plate heats up in response to thermal re-equilibration and experiences Barrovian-type metamorphism. The return of the lower plate rocks to the Earth's surface occurs by a combination of erosion and tectonic processes. In terms of rheology, strengthening of the thickened lithosphere may occur because of the underthrusting of a cold plate. Conversely, weakening is expected due to an increased radiogenic heat production in the thickened continental crust and upward advective heat transport during exhumation (e.g. Mancktelow & Grasemann 1997), which causes an increase of heat flow and hence a shallowing of the brittle-ductile layer transition zone (see for example Fig. 5 in Willingshofer &

Cloetingh 2003). Orogenic wedges such as the European Alps or the Pyrenees mainly consist of a stack of upper crustal slices detached from the subducting mantle and lower crust (Mufioz 1992; Schmid et al. 1996). Associated thickening juxtaposes upper crust material next to that of the lower crust or upper mantle, which are both stronger for the prevailing temperature and pressure conditions at depth. Furthermore, increased fluid activity arising from dehydration of watersaturated sediments during underthrusting probably also weakens the orogenic wedge (e.g Mainprice & Paterson 1984). We argue that the aforementioned processes result in significant lateral strength variations during the progressive evolution of collisional mountain belts in a way that the bulk strength of the orogenic wedge is less than that of the bordering undeformed continental lithospheres. Such conditions are the starting point for our modelling study, which elaborates on the role of a weak zone during subsequent collision and indentation tectonics.

Modelling setup M a t e r i a l s a n d initial setups

In our experiments the continental lithosphere consists of three or four layers representing the brittle crust, ductile crust, brittle upper mantle and ductile mantle in nature. Details of the model setup and the properties of the materials are given in Figures l a to c and in Table 1. Lateral changes of lithospheric strength are incorporated by laterally varying the thickness of the brittle and ductile crustal layers and the rheology of the brittle or ductile upper mantle (Fig. 1). In nature, such strength variations are expressed in the changing depth to the brittleductile boundary in the crust and in differences in the shear strength of the upper mantle as suggested for the present-day lithospheric strength across the European Central and Eastern Alps (Okaya et al. 1996; Willingshofer & Cloetingh 2003). Dry quartz and feldspar sands, both MohrColomb-type materials (Table 1), were used as rock analogues for the brittle crust and the brittle upper mantle (Fig. 2a-c). Non-Newtonian viscous layers representing the viscous crust (silicone mix I in Fig. l a - c ) and the viscous upper mantle (silicone mix I for the weak zone and silicone mix II for the foreland and indentor plates) are mixtures of Rhodorsile Gommetype or PDMS-type (material properties in Weijermars 1986) silicone with barite powder (Table 1). Additionally, we used a mixture of Rhodorsile Gomme-type silicone with Acid Oil

A N A L O G U E M O D E L L I N G OF C O L L I S I O N Z O N E S

279

Fig. 1. Initial setup for experiments A, B, and C. Location of the corresponding strength profiles are marked with A and B in the sketches of the model configuration. Strength profiles have been calculated using a failure criterion for brittle deformation in the form (o.1 - o.3) = 2[c +/xpgz(1 - A)]/(/~2 + 1) 1/2 - / x where (o.1 - o.3) is the differential stress, c the cohesion,/x the coefficient of friction, pgz the overburden pressure and A the ratio of water to overburden pressure (0 in our modelling); and for ductile deformation in the form (o.1 - o'3) = r/~/where (o'1 - o'3) is the differential stress, 77 the viscosity, and ~/the experimental shear strain rate. Material parameters are listed in Table 1. The black arrow indicates the direction of push. LVL, low-viscosity layer in experiment C.

280

E. WILLINGSHOFER ET AL.

Table 1. Mechanical properties of analogue materials used in this study Layer Brittle crust Ductile crust Brittle upper mantle Viscous upper mantle Asthenosphere

Analogue material

Density, p (kg m-3)

Coefficient of friction*,/x

Dry feldspar sand Dry quartz sand Silicone mix I LVL, silicone mix Dry quartz sand Silicone mix I Silicone mix II Sodium polytungstate solution mixed with glycerol

1300 1510 1520 1100 1510 1520 1540 1800

0.75 0.9

Viscosity*, 77(Pa s)

Power n

1.8 • 105 1.8 • 104

1.8 2.2

1.8 x 105 7.2 • 105 1.2

1.8 2.0

0.9

Models B A, A, C B A, A, A,

C B, C B, C B, C B, C

*The coefficient of friction has been inferred with a Hubbert-type shear apparatus (e.g. Krantz 1991 ). Note that sand densities and/x values have been deduced for sifted sand layers. The cohesion of the brittle materials, which is actually the resistance of the sand grains to sliding, is 4 0 - 7 0 Pa. *The effective viscosity of the silicone mixtures for laboratory strain rates (,-~1 x 10 -6 s - i ) has been determined with a coni-cylindrical viscometer under room temperature (21 _+ 1 ~

as a low-viscosity layer (LVL in Fig. lc) above the model Moho. The above named layers are floating on a dense, low-viscosity fluid simulating the asthenosphere. Mantle and crustal layers were laid down one by one atop the model asthenosphere. The uppermost layer, the brittle upper crust, has been added by sprinkling layers of coloured sand, including a 4 • 4 cm grid on the surface to facilitate strain analysis. The models were allowed to re-equilibrate isostatically for a time period of c. 16 hours. The experiments discussed in this paper have been conducted within a 50 • 40 • 15 cm plexiglass tank. Glass walls attached to the long sides of the box reduced friction. Solid walls confined three sides of the model while the fourth, a layer of silicone putty labelled 'Lateral Confinement' (Fig. 1), allowed limited amounts of lateral escape. At variance with Ratschbacher et al. (1991a), we deliberately chose a deformable, yet confined lateral boundary to avoid a priori facilitated lateral extrusion. Experiment A is the reference model. During modelling, the surfaces of all experiments have been scanned with a three-dimensional video-laser every 30 min. The gathered data have then been converted to digital elevation models (DEM, Figs 2 to 4), which are crucial for the understanding of the relationships between evolving structures and topography. The strength of the model indentor and foreland plates is the main difference between experiments A and B (Fig. 1). Experiment B contains a brittle upper mantle, which, compared to experiment A, increases the strength difference between the central weak zone and the bordering strong plates. The rheological stratification of the weak zone is the main difference between

experiments A and C (Fig. 1). The buoyant low-viscosity layer placed above the model C Moho simulates a decoupling horizon between the crust and the upper mantle and aspires to explore the influence of a deep-seated rheological weakness such as partial melts or zones of high fluid activity on the structural response of the shallower crust. Scaling procedure Geometric and dynamic similarity is a prerequisite for a valid comparison between analogue experiments and natural prototype (e.g. Weijermars & Schmeling 1986). The experiments are geometrically scaled by applying a 7 length ratio of 5 x 1 0 - , equating 1 cm in the models to 20 km in nature. For both the foreland and indentor plates we assumed that the brittle crust has twice the thickness of the ductile crust, a configuration that embodies a strong lithosphere of pre-Mesozoic thermotectonic age (Cloetingh & Burov 1996). This thickness ratio has been reversed for the preexisting weak zone simulating young, thermally reset Alpine-type crust (e.g. Genser et al. 1996). The dimensionless stress-scale factor for the brittle Mohr-Coulomb-type materials of grain size 100-300 txm used in the models is determined by the relationship or* = p* g*~*

(1)

where * stands for model-prototype ratios of stress, density, gravitational acceleration and thickness, respectively (e.g. ~ =O'model/O'prototype). All experiments have been performed under normal gravity conditions, hence g* equals 1.

ANALOGUE MODELLING OF COLLISION ZONES p* = 2.8 x 10 -7 and 2.3 • 10 -7 for densities of the upper crust and upper mantle of 2700 and 3300 kg m -3 in nature, respectively. The Ramberg number (Rm) (Weijermars & Schmeling 1986; Sokoutis et al. 2000) has been adopted as a measure for dynamic similarity of the ductile layers of the model and the natural example, where

taken into account as large error bars are attached to the quantity of such anomalies (crustal and lithosphere thickness) for the geologic past. The models also ignored natural recovery processes such as erosion and re-sedimentation and possible influences stemming from earlier subduction of oceanic lithosphere. In our experiments, the weak zone/strong plate boundaries are vertical, which is not likely to occur in nature at the scale of the lithosphere. However, we are confident that this simplification does not bias the main results derived from lateral strength gradients. For a concise discussion on the influence of different initial geometries of the weak zone on collision and indentation tectonics the reader is referred to Willingshofer et al. (2004).

pgh 2 Rm -- - r/V

(2)

and p and h are the density and thickness of the ductile layer, respectively, g is the acceleration of gravity, r / t h e viscosity of the ductile layer and V the compression rate. To fulfil the criterion of dynamic similarity, Rm of the experiment and the prototype must be approximately the same. A summary of the scaling parameters is presented in Table 2.

Simplifications Analogue models simplify the complexity of nature such that only the first-order rheological stratification can be accounted for. Although ductile materials with power law rheology were incorporated to simulate creep processes, they maintained their properties throughout the thickness of the layer and hence temperaturedependent variations of the creep strength with depth were not integrated. We attempted to translate temperature and thickness variation effects into implicit mechanics. For this purpose, we (1) assigned the same rheology to the ductile crust and ductile upper mantle within the weak zone, and (2) only considered that part of the lithosphere that has significant strength. As such, the mantle layers in the experiments are thinner than in nature and lateral thickness variations possibly due to previous thickening events, which might reduce the bulk strength of the lithosphere as argued in the 'experimental concept' section, are neglected. By assuming the same initial thickness for all lithosphere layers in the models, gravitational effects in relation to pre-existing mass anomalies are not

281

Modelling

results

The following sections summarize and display the modelling results in chronological order. The terms 'thrust' and 'backthrust' refer to thrusting in or against the direction of push from the moving wall, respectively. The plate in contact with the moving wall is the 'indentor' and the distal plate is the 'foreland'.

Experiment A Surface deformation and evolution of topography. The first structures were thrusts and backthrusts that formed simultaneously (Fig. 2a). These faults cut the model surface at or close to the boundaries between the strong and weak plates, and are clearly visible as linear features in the digital elevation model (DEM). Surface uplift of the weak plate is small and affects all the area between the thrust faults. As convergence continued, thrusting propagated outwards with respect to the central weak plate (Fig. 2b). A sequence of backthrusts developed within the brittle crust of the indentor. After 10% bulk-shortening, thrusting flipped to the foreland side of the weak plate activating thrust 5 in Figure 2b. The weak plate folded while thrusting propagated away into the indentor and

Table 2. Scaling parameters of analogue experiment and prototype Layer Ductile crust model Lower crust nature Ductile mantle model Ductile mantle nature

Density, p (kg m -3)

Viscosity, r/ (Pa s)

Layer thickness, h (m)

1520 2900 1540 3300

1.8 • 105 ,-~1022 7.2 • 105

5 X 10 -4

1021- l 0 22

2 x 104

1 x 10 4 1 x 10 -3

Velocity, v (m s- )

Rm

1.9 1.6 1.9 1.6

1.1 1.7 1.2 8.1

x x x x

10 -6 10 -a~ 10 -6 10 -1~

282

E. WILLINGSHOFER ET AL.

Fig. 2. Modelling results of experiment A for different stages of deformation are displayed as: (a-d) topview images (left panel), digital elevation models (DEMs) (right upper pannel), line drawings showing the structural interpretation (right lower pannel), and (e, f) cross-sections. The green rectangle in (a) indicates the area extent of the DEM. The numbers in the figures indicate the sequence of the evolving structures. Black arrows in (a-f) indicate the direction of push. Note that the model surface corresponds to the uppermost coloured sand layer.

ANALOGUE MODELLING OF COLLISION ZONES foreland plates. Consequently, the uplifted weak plate had an antiformal surface flanked on both sides by depressions (DEM, Fig. 2c, d). The foreland depression was weakly developed and originated from thrust loading. The indentor depression spanned approximately the width of the backthrust system and formed a syncline followed by an anticline close to the moving wall. Noteworthy at this stage of convergence is the development of small piggy-back basins in the hangingwall of thrusts 1 (Fig. 2b). Such thrust-parallel basins would capture detritus derived from the raising antiformal weak plate. After 20% bulk-shortening, outward thrust propagation stopped and convergence was taken up by out-of-sequence thrusting along thrusts 1 and smaller-scale backthrusts (9) within the weak plate (Fig. 2c). Coeval folding continued and led to surface uplift on the weak plate and subsidence in the backthrust region (DEM, Fig. 2c). Little distortion of the surface grid indicates that lateral escape was then minor. Until the end of the experiment, shortening was mainly taken up by out-of-sequence thrusting (Fig. 2d). The topographic low at the weak plate/indentor boundary coincided with a trench in front of the main backthrust of the model orogen (DEM, Fig. 2d). This backthrust carried material over the earlier backthrusts. As a consequence, no noticeable change of surface altitude occurred within the weak plate during the last few centimetres of shortening. Minor strike-slip faulting along the weak plateindentor boundary is a boundary effect impelled by friction along the sidewall and hence contains no implication for geological interpretation. Dextral strike-slip faulting on the lateral confinement side was related to the low resistant lateral material. The final widths of the weak and strong plates, deduced from measurements at the model surface, indicate that both strong plates have been shortened by about 45% on average. The weak zone was shortened by 17% and underwent stretching orthogonal to the shortening direction by 10%. The deformed surface grid reveals that stretching was unevenly distributed along strike with minimum stretch close to the glass wall (7%) and maximum stretch close to the lateral deformable confinement (19%).

Cross-sections. Cross-sections of experiment A document the finite state of shortening, which has arisen from interplay of thrusting and folding (Fig. 2e, f). The upper weak plate is an anticline with a thick core of ductile crust and upper mantle. Thickening of the weak plate is due to homogeneous thickening of its ductile

283

layers (about 90%). The anticline is slightly asymmetric, exhibiting an overturned limb facing the indentor. Thrusting led to considerable thickening of the brittle crust, especially in the backthrust area where initial thickness has doubled (Fig. 2e). These backthrusts rotated towards steeper dips to accommodate folding of the indentor. In contrast, the brittle layer of the weak plate was not faulted. Shortening was taken up within the lower crusts of the foreland and indentor plates by homogeneous thickening, which is largest (c. 25%) close to the back wall and the moving wall and least (c. 13%) at the boundary with the weak plate. Along strike differences in the model orogen reflect differences in the strength of the confining material (glass wall versus deformable silicone putty). Shortening of the indentor close to the rigid glass wall can only be taken up by folding and backthrusting. These structures are less developed close to the lateral confinement that allowed the model to escape laterally.

Experiment B Surface deformation and evolution of topography. A backthrust appeared at the surface of experiment B at the weak plate/indentor boundary after 1.6% bulk-shortening (Fig. 3a). Thereafter, thrusting mainly took place within the weak plate (Fig. 3b). Imbrication in its brittle layer is associated with relief growth characterized by model orogen-parallel ridges and troughs. This structurally controlled landscape is distinctly different from the smooth dome-shaped topography of experiment A (compare DEM of Fig. 2b, c and Fig. 3b, c). Deformation essentially remained within the weak plate up to 15% bulk shortening and was taken up by thrusting and backthrusting. Dextral strike-slip faults close to the lateral confinement are accommodation structures formed in response to thrust tectonics (Fig. 3c). Subsequently, convergence was accommodated by conjugate thrusts defining pop-up structures away from the weak plate and by backthrusting along fault 1 (Fig. 3c). Thrusting occurred dominantly close to both the back- and advancing walls. A pronounced increase of topography over the weak plate suggests that thickening of the ductile layers controlled the topographic evolution at this stage (DEMs of Fig. 3b, c). At the end of the experiment, shortening was distributed throughout the model, as shown by the distribution of active thrusts and backthrusts (Fig. 3d). Folding of the indentor had a lower amplitude than in experiment A, which reflects higher strength (strength profiles of Fig. la

284

E. WILLINGSHOFER ET AL.

Fig. 3. Modelling results of experiment B for different stages of deformation. ( a - d ) topview images (left panel), digital elevation models (DEMs) (right upper pannel), line drawings showing the structural interpretation (right lower pannel), and (e, f) cross-sections. The green rectangle in (a) indicates the area extent of the DEM. The numbers in the figures indicate the sequence of the evolving structures. Black arrows in (a-f) indicate the direction of push. Note that the model surface corresponds to the uppermost coloured sand layer. Symbols as in Fig. 2.

ANALOGUE MODELLING OF COLLISION ZONES and b). For the same reason, loading of the indentor by the model orogen along backthrust 1 yielded a shallower trench. At the end of the deformation history, the weak zone was shortened by 44% and stretched by 9% parallel to the strike of the model orogen. Compared to experiment A, more shortening was taken up by deformation of the weak zone and less (28%) by the adjacent strong plates.

Cross-sections. Cross-sections of experiment B show a doubly vergent orogen and gentle folding of the adjacent stronger lithospheres (Fig. 3e, f). The brittle crust of the weak plate is truncated by several thrusts and backthrusts. Repeated activation yielded a complex pattern of partly interfering reverse faults. In both the foreland and indentor plates, thrusting within the brittle layers took place away from the weak plate, in response to stress concentration close to the distal wall and the moving wall. The initially parallelepipedic ductile layers of the weak plate were converted into elongate bodies with up to six times the original thickness due to intense horizontal shortening (Fig. 3e). The ductile crust of both the foreland and indentor plates has been thickened by 8% on average. Besides being folded with the rest of the model (Fig. 3e, f), the brittle upper mantle shows little signs of deformation but strongest thickening in the hinge zones. In section A (Fig. 3e), one single thrust developed in the foreland plate close to the contact with the weak plate. The fault plane, which lines up with the top surface of the ductile mantle (white layer) is oriented at a high angle to the main compression direction, suggesting that thrusting pre-dated folding. Homogeneous thickening of the ductile mantle material absorbed shortening of the weak plate. The ductile mantles of the foreland and indentor plates show long-wavelength folds. Thickening was highest in the fold hinges near the weak plate. Experiment C Surface deformation and evolution of topography. The formation of a pop-up uplift on the weak plate started soon after shortening was imposed (Fig. 4a). Surface uplift was due partly to active faulting and partly to the buoyancy of the low-viscosity layer. As shortening continued, a backthrust at the weak plate-indentor boundary became active and the weak plate got folded with the former pop-up structure lying as a tectonic klippe in a syncline between two anticlines (Fig. 4b). At the same time, material was expelled from the weak plate

285

towards the lateral confinement, guided by a set of strike-slip faults. The evolving topography was distinctly asymmetric with the anticline close to the indentor being highest. Dominant foreland-directed thrusting initiated at 11% bulk shortening along thrust 5, which remained active until the end of the experiment (Fig. 4c, d). Thrusting of the model orogen over the foreland plate induced flexure and hence a foreland-type basin parallel to the frontal thrust 5 (DEM, Fig. 4c). In contrast, no similar feature was observed on the indentor side. The raising topography of the weak plate cannot solely be related to thrusting because shortening also tightened folds, with normal faulting occurring along anticline crests (Fig. 4c). Such structures demonstrate internal deformation of the model orogen during thrusting over the foreland. The foreland-facing anticline became overturned and transformed into a recumbent fold-nappe during the final phases of shortening. A mature foreland basin developed while thrusting continued to load the foreland plate (DEM, Fig. 4d). The finite model configuration documents 58% shortening and 20% orogen-parallel lengthening of the weak zone. Unlike the above-described experiments where shortening was distributed rather evenly over the foreland and indentor plates, distinctly more of the convergence was accommodated in experiment C on the foreland plate site (55%) through long-lasting activation of thrust 5. Shortening of the indentor plate was about 16%.

Cross-sections. Cross-sections of experiment C (Fig. 4e, f) show upright and overturned folds in the weak plate. Thrusts and backthrnsts cut the brittle layer at inflection points of the folds and normal faults dissect the crest of upright anticlines. A major thrust (no. 5) separates the weak plate from the underthrust foreland plate bent by the load of the model orogen. This thrust links up with a broad ductile shear zone within the weak zone, which separates the foreland from the indentor plates. The brittle layers of the foreland and indentor plates are free of faulting, close to the moving wall excepted. The buoyant low-velocity layer remained trapped at depth. Part of it accumulated beneath the strong ductile mantle of the indentor and the other part was expelled towards the weak lateral confinement. The fact that the indentor only shows a small inclination towards the weak plate reflects the combined effects of the buoyancy of the accumulated low-viscosity layer underneath the indentor and the small loading by the model orogen. This former effect was probably decisive for the switch to dominant foreland-directed

286

E. WILLINGSHOFER ETAL.

Fig. 4. Modelling results of experiment C for different stages of deformation. ( a - d ) topview images (left panel), digital elevation models (DEMs) (right upper pannel), line drawings showing the structural interpretation (right lower pannel), and (e, f) cross-sections. The green rectangle in (a) indicates the area extent of the DEM. The numbers in the figures indicate the sequence of the evolving structures. Black arrows in (a-f) indicate the direction of push. Note that the model surface corresponds to the uppermost coloured sand layer. Symbols as in Fig. 2.

ANALOGUE MODELLING OF COLLISION ZONES thrusting at 11% bulk-shortening. Short-wavelength folding accommodated shortening of the ductile mantle below the low-viscosity layer. The ductile crust and upper mantle of both the foreland and indentor plates thickened increasingly from the boundaries with the weak plate (0-5%) towards the confining walls (< 15%). The laterally escaped part of the weak plate spread into the lateral confinement and, therefore, had thinner layers and slightly lower topography.

Summary and general discussion of modelling results In this section we highlight similarities as well as differences of the models and discuss them in the frame of previously published modelling studies. (1)

(2)

Our modelling results are consistent with previous numerical and analogue models (e.g. Cloetingh et al. 1999 and references therein; Davy & Cobbold 1991; Martinod & Davy 1994; Burg et al. 1994) and show that folding is a primary response to shortening of the continental lithosphere. The above quoted authors have shown that homogeneous lithosphere (without strength variations) starts to fold soon after compression commences. Modelling with a weak zone within the lithosphere deviates on the following points. (a) Folding is suppressed in case of high-strength contrasts between weak and strong plates (experiment B). In such a situation, shortening is accommodated by thickening of the weak zone for a long period of time and folding of the indentor and to a lesser extent the foreland plate are late-stage features. (b) The presence of a mechanical instability within the weak zone (our buoyant lowviscosity layer in experiment C) controls the locus of folding through strain localization such that after about 10% bulk-shortening most of the remaining convergence was taken up along thrust number 5, which is interpreted to continue towards the centre of the model orogen as a ductile shear zone (Fig. 4e, f). Consequently, folding was confined to the weak zone and underthrusting of the foreland plate was the more energetic response to compression. Indentation tectonics requires strength differences between colliding plates. Strength differences are manifested by stronger deformation in the weaker plate than in the stronger plate. Depending

(3)

287

on the regional mechanical boundary conditions, indentation causes thickening and/or activation of large strike-slip faults guiding lateral escape of crustal blocks. The stronger plate becomes deformed when thickening can no longer absorb shortening. Such an evolution has been described from various natural examples like the Anatolian region (McKenzie 1972), Asia (Molnar & Tapponnier 1975) and the European Alps (Ratschbacher et al. 1991a, b). Indentation tectonics have been approached experimentally by using rigid indentors (Tapponnier et al. 1982; Davy & Cobbold 1988; Ratschbacher et al. 1991a; Martinod et al. 2000; Rosenberg et al. 2004), hence forcing deformation to take place within the deformable plate. Previous work focused on lateral extrusion and deformation in front of the indentor, which are both strongly controlled by the degree of lateral confinement. Although our experiments have not been designed to specifically study this process, comparison of experiments A and B (Figs 2 and 3) with C (Fig. 4) confirms the findings of earlier works that the degree of confinement is important for extrusion tectonics. In our work the deformable indentor and foreland plates had the same properties and were subject to lithospheric-scale folding (experiments A and B), bending due to loading (experiment C), and thrusting in the brittle layers. We infer that a strong lateral confinement favoured folding of the indenter (Figs 2 and 3), whereas a weak confinement enhanced lateral material transport towards the deformable confinement (Fig. 4). The amount of orogen-parallel extension, deduced from the deformed surface grid, was in all experiments least close to the glass wall and highest in proximity to the deformable confinement. It is worth noting that shortening of the model with the higher strength contrast (experiment B) did not increase the amount of orogenparallel extension (9%), compared to experiment A (10%). Instead, shortening resulted in a narrower and thicker weak zone. The relative strength contrast between the weak plate and the adjacent indentor and foreland plates appears to determine the sequence of deformation. High strength contrasts (experiment B) concentrated deformation within the weak plate, while low strength contrasts favoured deformation within both the indentor and

288

(4)

E. WILLINGSHOFER ETAL. foreland plates (experiment A). These results are consistent with models that employed indentor and indented plates of equal strength (Davy & Cobbold 1988) and models using a strong, cratonic-type, lithosphere as indentor (Keep 2000). In the absence of any inherited asymmetry from preceding subduction, we ascribe the model deep structures to intracontinental processes. We infer that a low strength contrast (experiment A) between the weak and strong plates leads to buckling dominated deformation in the weak plate and, to a lesser amount in the adjacent strong plates. As a result, the topographic maximum located within the weak zone does not correspond to the deepest Moho, which is the trough syncline of the indentor. High strength contrasts (experiments B) favour the formation of crustal and lithospheric roots. In this model the topographic maximum correlates well with the locus of the crustal root. A distinctly asymmetric structure of the model orogen is only observed where a mechanical instability above the Moho within the weak zone was initially present (experiment C). Once the shape of this low density instability has become asymmetric due to the imposed kinematic boundary conditions (push from one side), it facilitated strain localization over long periods of time. As a consequence, most of the shortening was taken up by a thrust at the base and a thick ductile shear zone within the central zone of the model orogen (Fig. 4e, f). Analogue plates accommodated shortening in the ductile crust by homogeneous thickening (Models A, B, and C) thereby largely maintaining its initial geometry. These results are consistent with those of a numerical modelling study focusing on the role of lower crust rheology during collision tectonics (Gerbault & Willingshofer 2004). The same authors suggest that a strength difference of about 20 MPa between the lower crust of the weak zone and the indentor is sufficient to favour lower crust indentation above lateral (opposite to the direction of shortening) ductile flow. In our experiments, indentation occurs on the scale of the lithosphere and interference patterns with the weak zone are simple. Conditions favourable for lower crust indentation as envisaged in the European central Alps demand decoupling at upper crust/lower crust and at Moho levels (e.g. Pfiffner et al. 2000).

Discussion of modelling results in the light of post-early Oligocene tectonics of the Eastern Alps In the following sections we discuss our modelling results in the frame of the late Oligocene to present geodynamic evolution of the Eastern Alps. At this point we like to emphasize that our experiments have not been designed to reproduce the deformation history of the Eastern Alps in great detail. Instead, we have chosen a conceptual approach, which restricts us to the discussion of first-order aspects relevant to tectonics of the Eastern Alps and other orogens in which a weak zone has been compressed. As such, we do not present a 'best fit model', but aim at deducing favourable mechanical conditions for Late Oligocene to Neogene Eastern Alps tectonics from experiments described in the previous sections.

The Eastern Alps as geological example

The European Alps are among the youngest features of the European lithosphere and owe their origin to continent-continent collision related to the Africa-Europe convergence (e.g. Frisch 1979; Dewey et al. 1989; Schmid et al. 1996). We regard the orogenic wedge of the Eastern Alps, which existed during the early Oligocene (c. 30 Ma ago) between the Adriatic plate to the south of the Periadriatic Line and the European foreland in the north (Fig. 5) (e.g. Frisch et al. 1998), as a pre-existing weak plate or partof-plate based on the following arguments. (1) The strong concentration of post-Eocene deformation, which affected Austroalpine and Penninic units (Fig. 5), in particular in the Tauern Window region, may be envisaged as expressing significant lateral strength variations, which controlled the locus of deformation (see also Ratschbacher et al. 1991a, b). (2) Cloetingh & Burov (1996) have shown that the strength of the European continental lithosphere strongly depends on the time elapsed since it was subject to loading and tectonic activity. Adopting this concept, it follows that the tectonic units comprising the orogenic wedge, which underwent thermal rejuvenation and loading during the Late Cretaceous (Austroalpine unit) and Eocene-Miocene (Penninic units), are weaker than the European and Adriatic plates, which have been affected by Variscan tectonics (e.g. Frey et al. 1999). (3) The radiogenic heat production, hence thermally induced softening within the orogenic wedge, which dominantly consists of continental slices (Austroalpine and

ANALOGUE MODELLING OF COLLISION ZONES

289

Fig. 5. Simplified tectonic map of the Eastern Alps showing the location of the TRANSALP deep seismic reflection and the ALP75 refraction lines. Palaeogeographic units after Froitzheim et al. (1996). Inset shows position of reflection seismic profiles across the European Alps. 1, TRANSALP; 2, NRP20-EAST (EGT); 3, NRP20-WEST; 4, ECORS-CROP.

Zentralgneiss units) with minor oceanic units, was probably higher than in the adjacent plates of normal crustal thickness. (4) Alpine ductile fabrics are related to Eocene-Miocene tectonics and are confined to the Tauern Window and a small strip of Austroalpine units flanking the Tauern Window to the south. The understanding of certain aspects of Eastern Alps tectonics, therefore, may justifiably refer to first-order features identified in the analogue models. We perceive relevance on several points, namely lateral extrusion and coeval orogen-parallel extension, backthrusting on Adria, buckling in the Tauern region, and interpretation of the deep seismic structures. Indentation and lateral escape

Late Oligocene to middle Miocene lateral extrusion in the Eastern Alps is interpreted to reflect the combined effects of horizontal forces related to indentation by the Adriatic plate and gravitational forces stemming from preceding crustal thickening in the Tauern Window region (Ratschbacher et al. 1991a and references

therein). Eastward lateral extrusion, towards the Pannonian Carpathian region, was conditioned by contemporaneous subduction and slab retreat along the Carpathian arc providing space for the extruding blocks (Royden 1988). Lateral extrusion ceased during the early late Miocene when subduction along the Carpathians terminated and the incoming buoyant European plate jammed the subduction process (Sandulescu 1988). Until that time, deformation of the Adriatic indenter was restricted to its northern margin, which became involved in extrusion tectonics (Frisch et al. 1998) and at the transition zone to the Dinarides (Castellarin & Cantelli 2000). Internal deformation of the Adriatic indentor commenced during the Late Middle Miocene through activation of the south vergent Valsugana thrust system, which documents southerly directed backthrusting (Castellarin & Cantelli 2000). Subsequent deformation events portray south-directed thrust propagation with active present-day thrusting in the Po Plain (Benedetti et al. 2000). In summary, significant deformation of the Adriatic indentor interior, amounting to about 50 km

290

E. WILLINGSHOFER ET AL.

north-south shortening (Sch6nborn 1999), postdates the climax of lateral extrusion in the Eastern Alps north of the Periadriatic Line and subduction along the Carpathian arc. In our study, only experiment B with the largest strength contrasts shows significant deformation of the indentor proper after about 15% bulk-shortening. This jump in deformation site takes place once the indented plate has been thickened to the point that vertical stress components hamper further thickening of this plate. As to whether the increased resistance at the lateral confinement contributes to activating deformation within inner indentor regions cannot be deduced with certainty from our experiments. The experiment with a low strength contrast (experiment A) records simultaneous deformation of the weak zone and the indentor more or less throughout the experiment, and hence exhibits little resemblance with the Eastern Alps. Strong strain localization over a long period of time (experiment C) also retards indenter deformation. From our experiments, we deduce that strength contrasts between the weak zone and the indentor are important to keep deformation localized in a particular region over a long period of time. At the same time one needs to bear in mind that our simple geometric and kinematic model setup neglects the influence of other processes, such as subduction along the Carpathians, which is intrinsically related to the opening of the Pannonian Basin (e.g. Horwith 1993), and probably facilitated lateral extrusion in the Alps through slab pull forces, thereby reducing the demand for high strength contrasts between indentor and indented region. Model B therefore most likely exaggerates natural conditions, for which more subtle strength variations have been suggested at least for the present-day situation across the Eastern Alps (Willingshofer & Cloetingh 2003). Geometric boundary conditions adopted in this study can on13/ account for about 20% bulk orogen parallel extension within the weak zone. Extension is not evenly distributed, but increases towards the deformable weak zone. Both the total amount of Oligocene to Miocene extension, which was in the order of 50% (Ratschbacher et al. 1991b; Frisch et al. 1998), and the locus of maximum extension, which was in front of the indenter dip (Frisch et al. 2000) differ from our modelling results. While the total amount of extension can be accounted for by a weaker or no lateral confinement (Ratschbacher et al. 1991a), the locus of extension might be controlled by the indenter geometry as recently suggested by analogue modelling by Rosenberg et al. (2004).

The structure o f the Tauern W i n d o w

The Tauern Window is an elongate dome formed during cooling after the thermal peak at 30 Ma (Lammerer & Weger 1998 and references therein). Its internal structure is characterized by short-wavelength (several kin), tight, upright folds (Fig. 6a) with N E - S W trending fold axis in the west and N W - S E trending fold axes in the east. Fold limbs are associated with sinistral shear zones in the western and dextral shear zones in the eastern Tauern Window (e.g. Behrmann & Frisch 1990; Kurz & Neubauer 1996). The western and eastern terminations of the Tauern Window against the overlying Austroalpine units are low-angle shear zones documenting normal displacement of hangingwall units (Selverstone 1988; Genser & Neubauer 1989). Structures within and close to the Tauern Window are consistent with models employing an Oligocene-Miocene transpressional stress regime (e.g. Lammerer 1988). It has been argued, against earlier work, that the window shows similarities with extensional core complexes (e.g. Selverstone 1988) genetically linked with lateral extrusion tectonics (Frisch et al. 2000). In models A and C highamplitude and short-wavelength buckle folds formed within the weak zone and controlled surface uplift. While buckling in model A occurred on the scale of the lithosphere, buckling in model C was confined to the crust above the buoyant low-viscosity layer, which allowed for full crust-mantle decoupling within the weak zone. Our results support the findings of Lammerer & Weger (1998), who emphasized the importance of crustal-scale buckling for the structural evolution of the western Tauern Window. Unlike the amount and spatial distribution of orogen-parallel extension accompanying folding (see previous section), buckling in front of the indentor appears to be a process independent of the indentor geometry and direction of shortening. Evidence stems from comparing our results with a straight indenter with those of Ratschbacher et al. (1991a) or Rosenberg et al. (2004), who used triangular indentors. Those works showed that contemporaneous shearing in strike-slip mode critically depends on indenter geometry and convergence direction. The strong lateral confinement in our models also suggests that at least part of the orogen-parallel component of stretch is related to the folding process itself. Our models show that a crust weaker than its surroundings was essential in the development of high-amplitude buckle folds. Inversely, if such folds are diagnostic of weak crusts

ANALOGUE MODELLING OF COLLISION ZONES

291

Fig. 6. (a) Structure of the Eastern Alps along the TRANSALPprofile based on surface geology. Redrawn after TRANSALP Working Group (2002). (b) Slightly modifiedinterpretation of the TRANSALPdeep seismic line from TRANSALP Working Group (2002). IF, Inntal Fault; PA, PeriadriaticFault; Val, ValsuganaThrust. compressed between stronger plates, the deep Tauern Window has to be treated separately in mechanical terms from either European and Adria plates to account for initial plate heterogeneity. In this respect, numerical modelling suggests that the Tauern Window region was distinctly weaker than the surrounding areas while folding and lateral escape took place (Genser et al. 1996). In accordance with the structure of the Tauern Window, we consider crustal-scale buckling as documented in model C as important processes in collisional settings described above. The deep structure of the Eastern Alps

Similarities in structures and evolution between the eastern Alps and the models are mechanical references to interpret deep seismic profiles. Key features of the deep structure of the TRANSALP transect (Fig. 6; TRANSALP Working Group 2002) include (1) a southdipping European lower crust of rather constant thickness, (2) a distinctly thicker Adriatic lower crust, which dips to the north, and (3) crustalscale ramps such as the Sub-Tauern Ramp carrying Adriatic lower and upper crusts on to the underthrusting European crust and the SubDolomite Ramp, which acted as a backthrust during imbrication of the Adriatic crust. The Moho of both the European and Adriatic plates

dips towards the mountain root. All of our experiments resulted in lower crust geometries of the foreland plate similar to that of the European lower crust (TRANSALP Working Group, 2002), experiment C being most resembling (compare cross-sections in Figs 2, 3, and 4 with Fig. 6). In the Central Alps, geophysical evidence suggests that the European lower crust is a strong layer of approximately constant thickness (Pfiffner et al. 2000 and references therein). Analogue plates accommodated shortening by homogeneous thickening (Models A, B, and C) thereby largely maintaining initial overall geometry. Accordingly, constant thickness does not imply that a layer is not deformed and that it has a high strength. Models A and B, where folding affected the indentor, result in Moho dipping towards the orogen as published by the TRANSALP Working Group (2002), whereas Model C shows a flat indentor Moho consistent with the Moho shape identified from receiver functions utilizing teleseismic events (Kummerow et al. 2004). The flat Moho in experiment C is a consequence of strain localization along the basal thrust, which places the model orogen over the foreland plate. This thrust (5 in Fig. 5), which can be interpreted as equivalent to the basal detachment of the northern fold and thrust belt in the Eastem Alps, joins a wide shear zone separating the foreland from the

292

E. WILLINGSHOFER ET AL.

indentor plates. Both the thrust and the shear zone remained active during the second half of the experiment and took up most of the shortening, hence hampering buckling and outward propagation of the backthrusts. The shear zone in experiment C dips towards the indentor plate and has the same function as the Sub-Tauern Ramp (Fig. 6), placing the indentor plate above the foreland plate. Inspired by experiment C, we view the Sub-Tauern Ramp as structurally continuous with the detachment at the base of the northern fold and thrust belt but, contrasting with the interpretation of the TRANSALP Working Group (2002), discontinuous with respect to the Inntal Fault.

Conclusions Conclusions from analogue experiments for collisional settings involving a weak plate are generic and applicable to nature. The relative strength between converging plates and intervening weak plates controls both the deformation location and style. Small contrasts favour buckling of both weak and indentor plates, whereas strong contrasts focus deformation onto the weak plate for a larger amount of shortening and result in the formation of important crustal roots. Buckling is a fundamental shortening mode of the lithosphere, even in cases with strength heterogeneities. The presence of a buoyant mechanical instability above the model Moho within the weak plate favours the development of asymmetric orogens notwithstanding the initially symmetrical setup. Furthermore, the mechanical weakness promotes strain localization and hence influences the sequence of deformations and their topographic expression. The absence of lower crustal wedges along the TRANSALP line does not necessarily reflect a lack of strain in the lower crust and does not allow drawing unambiguous conclusions on its strength. If the lower crust is weak, as in our experiments, strain is diffusely distributed over the foreland and indentor and results in homogeneous thickening of the lower crust layer. This type of deformation is consistent with the aseismic behaviour of the European lower crust along the TRANSALP transect. Experiment C suggests that the basal detachment, which floors the fold and thrust belt, is the northern structural continuation of the SubTauern Ramp. Experiment C further suggests that short-wavelength buckling within the Tauern Window requires decoupling of the upper crust from its upper mantle and probably also from its lower crust. The simple geometric

and kinematic boundary conditions used in this study show that folding is associated with orogen-parallel extension, which upon strong lateral confinement, is insufficient on its own to explain large-scale extension in the Eastern Alps. We are indebted to O. Merle and C. Rosenberg for their thorough reviews. E. W. acknowledges funding of this study by NWO, the Netherlands Organization for Scientific Research, project 810.31.003. D. S. thanks ISES (the Netherlands Centre for Integrated Solid Earth Science) and NWO for the financial support.

References BEHRMANN, J. & FRISCH, W. 1990. Sinistral ductile shearing associated with metamorphic decompression in the Tauern Window, eastern Alps. Jahrbuch der Geologischen Bundesanstalt, 133, 135-146. BENEDETTI, L., TAPPONNIER, P., KING, G. C. P., MEYER, B. & MANIGHETTI, I. 2000. Growth folding and active thrusting in the Montello region, Veneto, northern Italy. Journal of Geophysical Research, 105, 739-966. BURG,J.-P., DAVY,P. & MARTINOD,J. 1994. Shortening of analogue models of the continental lithosphere: new hypothesis for the formation of the Tibetian plateau. Tectonics, 13, 475-483. CASTELLARIN,A. & CANTELLI,L. 2000. Neo-Alpine evolution of the southern Eastern Alps. Journal of Geodynamics, 30, 251-274. CLOETINGH, S. • BUROV, E. 1996. Thermomechanical

structure of European continental lithosphere; constraints from rheological profiles and EET estimates. Geophysical Journal International, 124, 695 -723. CLOETINGH, S., BUROV, E. & POLIAKOV, A. 1999.

Lithospheric folding: primary response to compression? (from central Asia to Paris basin). Tectonics, 18, 1064-1083. DAVY, P. 8,: COBBOLD, P. 1988. Indentation tectonics in nature and experiment: 1. Experiments scaled for gravity. Bulletin of the Geological Institutions

of the University of Uppsala, 14, 129-142. DAVY, P. & COBBOLD, P. R. 1991. Experiments on shortening of a 4-layer model of the continental

lithosphere. Tectonophysics, 188, 1-25. DEWEY, J.

F.,

HELMAN, M.

L.,

TURCO, E.,

HUTTON, D. H. W. & S. D. K. 1989. Kinematics of the western Mediterranean. In: COWARD,M. P., DIETRICH, D. & PARK, R. G. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 265-284. FREY, M., DESMONS, J. & NEUBAUER, F. 1999. Metamorphic map of the Alps. Schweizerische Minera-

logisch Petrographische Mitteilungen, 79. FRISCH, W. 1979. Tectonic progratation and plate

tectonic evolution of the Alps. Tectonophysics, 60, 121-139. FRISCH, W., KUHLEMANN,J., DUNKL, I. & BRf3GEL, A.

1998. Palinspastic reconstruction and topographic evolution of the Eastern Alps during late Tertiary tectonic extrusion. Tectonophysics, 297, 1-15.

ANALOGUE MODELLING OF COLLISION ZONES FRISCH, W., DUNKL, I. & KUHLEMANN,J. 2000. Postcollisional orogen-parallel large-scale extension in the Eastern Alps. Tectonophysics, 327, 239-265. FROITZHEIM, N., SCHMID, S. M. & FREY, M. 1996. Mesozoic paleogeography and the timing of eclogite-facies metamorphism in the Alps: A working hypothesis. Eclogae Geologicae Helvetiae, 89, 81-110. GENSER, J. & NEUBAUER,F. 1989. Low angle normal faults at the eastern margin of the Tauern Window (Eastern Alps). Mitteilungen der Osterreichischen Geologischen Gesellschaft, 81, 233-243. GENSER, J., VAN WEES, J. D., CLOETINGH, S. ~z NEUBAUER, F. 1996. Eastern Alpine tectonometamorphic evolution; constraints from twodimensional P-T-t modeling. Tectonics, 15, 584-604. GERBAULT, M. • WILLINGSHOFER, E. 2004. Lower crust indentation or horizontal ductile flow during continental collision? Tectonophysics, 387, 169-187. HANDY, M. 1990. The solid-state flow of polymineralic rocks. Journal of Geophysical Research, 95, 8647-8661. HORV,~TH, F. 1993. Towards a mechanical model for the formation of the Pannonian basin. Tectonophysics, 226, 333-357. HUDLESTON, P. J. & LAN, L. 1993. Information from fold shapes. Journal of Structural Geology, 15, 253-264. KEEP, M. 2000. Models of lithospheric-scale deformation during plate collision: effects of indentor shape and lithospheric thickness. Tectonophysics, 326, 203-216. KRANTZ, R. W. 1991. Measurements of friction coefficients and cohesion for faulting and fault reactivation in laboratory models using sand and sand mixtures. Tectonophysics, 188, 203-207. KUMMEROW, J., KIND, R. ETAL. 2004. A natural and controlled source seismic profile through the Eastern Alps: TRANSALP. Earth and Planetary Science Letters, 225, 115-129. KURZ, W. &: NEUBAUER,F. 1996. Deformation partitioning and shear localization during the updoming of the Sonnblick area in the Tauern Window (Eastern Alps, Austria). Journal of Structural Geology, 18, 1327-1343. LAMMERER, B. 1988. Thrust-regime and transpression-regime tectonics in the Tauern Window (Eastern Alps). Geologische Rundschau, 77, 143-156. LAMMERER, B. & WEGER, M. 1998. Footwall uplift in an orogenic wedge: the Tauern Window in the Eastern Alps of Europe. Tectonophysics, 285, 213-230. MAINPRICE, D. U. & PATERSON, M. S. 1984. Experimental studies of the role of water in the plasticity of quartzites. Journal of Geophysical Research, 89, 4257-4269. MANCKTELOW, N. S. & GRASEMANN,B. 1997. Timedependent effects of heat advection and topography on cooling histories during erosion. Tectonophysics, 270, 167-195.

293

MARTINOD, J. & DAVY, P. 1994. Periodic instabilities during compression of the lithosphere 2. Analogue experiments. Journal of Geophysical Research, 99, 12057-12069. MARTINOD, J., HATZFELD, D., BRUN, J. P., DAVY,P. & GAUTIER, P. 2000. Continental collision, gravity spreading, and kinematics of Aegea and Anatolia. Tectonics, 19, 290-299. MCKENZIE, D. 1972. Active tectonics of the Mediterranean region. Geophysical Journal of the Royal Astronomical Society, 55, 109-185. MOLNAR, P. & TAPPONNIER, P. 1975. Cenocoic tectonics of Asia: effects of a continental collision. Science, 189, 419-426. Mur~oz, J. A. 1992. Evolution of a continental collision belt: ECORS-Pyrenees crustal balanced cross-section. In: MCCLAY, K. R. (ed) Thrust Tectonics. Chapman & Hall, London, Chapter 8, 235 -246. OKAYA, N., CLOETINGH, S. 8z Mt)LLER, S. 1996. A lithospheric cross section through the Swiss Alps (part II): constraints on the mechanical structure of a continent-continent collision zone. Geophysical Journal International, 127, 399-414. PFIFFNER, O. A., ELLIS, S. & BEAUMONT, C. 2000. Collision tectonics in the Swiss Alps: insight from geodynamic modeling. Tectonics, 19, 10651094. RANALLI, G. 1997. Rheology of the lithosphere in space and time. In: BURG, J. P. & FORD, M. (eds) Orogeny Through Time. Geological Society, London, Special Publications, 121, 19-37. RANALLI, G. 2000. Rheology of the crust and its role in tectonic reactivation. Journal of Geodynamics, 30, 3-15. RANALLI, G. & MURPHY, D. C. 1987. Rheological stratification of the lithosphere: Tectonophysics, 132, 281-295. RATSCHBACHER, L., MERLE, O., DAVY, P. & COBBOLD, P. 1991a. Lateral extrusion in the eastern Alps, part l: boundary conditions and experiments scaled for gravity. Tectonics, 10, 245 -256. RATSCHBACHER, L., FRISCH, W., LINZER, H.-G. & MERLE, O. 1991b. Lateral extrusion in the eastern Alps, part 2: structural analysis. Tectonics, 10, 257-271. ROSENBERG, C. L., BRUN, J. P., CAGNARD, F. & GAPAIS, D. 2004. Coeval shortening, thickening and orogen-parallel extension: laboratory models applied to the Eastern Alps. In: BERTOTTI, G., BUITER, S., RUEEO, P. & SCHREURS, G. (eds) From Mountains to Sedimentary Basins: Modelling and Testing Geological Processes. Geomod 2004 International Conference, Emmetten, Switzerland, Bolletino di Geofisica, 45, 136-140. ROYDEN, L. H. 1988. Late Cenocoic tectonics of the Pannonian Basin system. In: ROYDEN, L. H. & HORVATH, F. (eds) The Pannonian Basin, A Study in Basin Evolution. AAPG Memoir, Tulsa, 27 -48. SANDULESCU, M. 1988. Cenozoic tectonic history of the Carpathians. In: ROYDEN,L. H. & HORV~TH, F.

294

E. WILLINGSHOFER ETAL.

(eds) The Pannonian Basin, A Study in Basin Evolution. AAPG Memoir, Tulsa, 17-25. SCHMID, S. M., PFIFFNER, O. A., FROITZHEIM, N., SCHONBORN, G. & KISSLING, E. 1996. Geophysical-geological transect and tectonic evolution of the Swiss-Italian Alps. Tectonics, 15, 1036-1064. SCH(SNBORN, G. 1999. Balancing cross sections with kinematic constraints: the Dolomites (northern Italy). Tectonics, 18, 527-545. SELVERSTONE,J. 1988. Evidence for east-west crustal extension in the eastern Alps: implications for the unroofing history of the Tauern Window. Tectonics, 7, 87-105. SOKOUTIS, D., BONINI, M., MEDVEDEV, S., BOCCALETT1,M., TALBOT, C. J. & KOYI, H. 2000. Indentation of a continent with a built-in thickness change: experiment and nature. Tectonophysics, 320, 243-270. TAPPONNIER, P., PELTZER, G., LE DAIN, A. Y., ARMIJO, R. & COBBOLD, P. 1982. Propagating extrusion tectonics in Asia: new insights from simple experiments with plasticine. Geology, 10, 611-616. TRANSALP Working Group. 2002. First deep seismic reflection images of the Eastern Alps reveal giant crustal wedges and transcrustal ramps. Geophysical Research Letters, 29, 10.1029/2002GL014911.

WEIJERMARS, R. 1986. Flow behaviour and physical chemistry of bouncing putties and related polymers in view of tectonic laboratory applications. Tectonophysics, 124, 325-358. WEIJERMARS, R. & SCHMELING,H. 1986. Scaling of Newtonian and non-Newtonian fluid dynamics without inertia for quantitative modelling of rock flow due to gravity (including the concept of rheological similarity). Physics of the Earth and Planetary Interiors, 43, 316-330. WILLINGSHOFER, E. & CLOETINGH, S. 2003. The present-day lithospheric strength of the Eastern Alps and its relationship to neotectonics. Tectonics, 22, 1075, doi: 10.1029/2002TC001463. WlLLINGSHOFER, E., SOKOUTIS, D. & BURG, J.-P. 2004. Influence of the initial geometry of lateral strength variations during continental convergence: an analogue modelling study. In: BERTOTTI, G., BUITER, S., RUFFO, P. & SCHREURS, G. (eds) From Mountains to Sedimentary Basins: Modelling and Testing Geological Processes. Geomod 2004 International Conference, Emmetten, Switzerland, Bolletino di Geofisica, 45, 164-168. ZIEGLER, P. A., VAN WEES, J. D. & CLOETINGH, S. 1998. Mechanical controls on collision-related compressional intraplate deformation. Tectonophysics, 300, 103-129.

Origin of the current stress field in the western/central Alps: role of gravitational re-equilibration constrained by numerical modelling BASTIEN DELACOU, CHRISTIAN SUE, J E A N - D A N I E L C H A M P A G N A C & MARTIN BURKHARD I n s t i t u t d e G d o l o g i e , U n i v e r s i t ~ d e Neuchgttel, S w i t z e r l a n d

Abstract: We interpret the strain and stress fields of the western/central Alpine arc on the

basis of 2.5D finite element modelling and a recent seismotectonic synthesis. Models have fixed boundary forces and different crustal geometries, so that they respond to buoyancy forces (variations in gravitational potential energies). The seismotectonic regime, characterized by orogen-perpendicular extension in the high topographic core of the belt and local orogenperpendicular compressional/transpressional deformation in the external zones, appears to be very close to the modelled gravitational regime. Rotation of Apulia has a minor effect on the current strain or stress fields of the Alpine realm. Nevertheless, it could help to explain the orogen-parallel dextral faulting that is observed all along external zones, from the northern Valais to the Argentera external crystalline massif. Our results highlight the consequences for the Alpine realm of ongoing convergence between the African and European plates. Our interpretation is that collision is no longer ongoing and that buoyancy-driven stresses dominate the present-day geodynamics of the western/central Alps.

The Alpine belt has resulted from Tertiary collision between the Apulian micro-plate (considered as an African promontory) and the European plate, following Late Cretaceous to Eocene subduction of the Alpine Tethys (Coward & Dietrich 1989; Dewey et al. 1989; Laubscher 1991; Stampfli et al. 1998; Schmid & Kissling 2000). Whereas compressional structures, such as nappes, metamorphic zones and phases of folding have been well documented (e.g. Choukroune et al. 1986; Fry 1989; Burkhard 1990; Pognante 1991; Butler 1992; Spalla et al. 1996; Duch~ne et al. 1997; Burkhard & Sommaruga 1998; Becker 2000) the current tectonic context remains debatable. Is collision still active or has the Alpine belt come to the end of its compressive history? The relatively recent discovery of extensional tectonics, through seismotectonic and structural analyses (Mancktelow 1992; Maurer et al. 1997; Eva et al. 1998; Ftigenschuh et al. 1999; Sue et al. 1999; Bistacchi et al. 2000; Kastrup 2002; Sue & Tricart 2002, 2003; Champagnac et al. 2003; Champagnac et al. 2004) goes a long way toward answering the question. Extensional earthquakes have been known for a long time (Pavoni 1961; Ahorner et al. 1972; Frtchet 1978). The large-scale seismotectonic synthesis

of Delacou et al. (2004) demonstrates that an extensional regime operates throughout all the internal zones of the belt. In addition, structural analyses of fault slip data indicate that extensional tectonics have been prevalent in these zones since at least Miocene times (Mancktelow 1992; Bistacchi et al. 2000; Tricart et al. 2001; Sue & Tricart 2003; Champagnac et al. in press). Extension is therefore a major feature of the recent to present-day geodynamics of the Alpine arc. Various contradictory models have been put forward to explain such intra-orogenic extensional tectonics: (1) large-scale buckling under compressive conditions combined with outer-arc extension (Burg et al. 2002), (2) lateral extrusion in an active convergent belt (Ratschbacher et al. 1991; Frisch et al. 2000; Sachsenhofer et al. 2000), (3) slab break-off re-equilibration (Davies & yon Blanckenburg 1995; Sue 1998), (4) rotational tectonics (Calais et al. 2002; Collombet et al. 2002), and (5) gravitational re-equilibration of an overthickened crust (Bada et al. 2001). While overall convergence between the African and European plates is still ongoing at a rate of 3 to 8 m m a -~ (Argus et al. 1989; Demets et al. 1994; Albarello et al. 1995; Crttaux et al. 1998; Nocquet 2002), the boundary conditions

From: GAPAIS,D., BRUN,J. P. & COBBOLD,P. R. (eds) 2005. Deformation Mechanisms, Rheology and Tectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 295-310.

0305-8719/05/$15.00

9 The Geological Society of London 2005.

296

B. DELACOU ET AL.

around the Alpine belt, as estimated by recent GPS results (Calais et al. 2002; Nocquet 2002; Vigny et al. 2002; Nocquet & Calais 2003, 2004), reveal no clear relative movements between the Apulian and European microplates. Velocities across the belt are between 1 and 2 mm a - l and they provide no clear indication of convergence or divergence. At best, the GPS data indicate anticlockwise rotation of Apulia with respect to Europe at an angular velocity of 0.52~ about a pole near Milan (Calais et al. 2002). In this study we use numerical modelling and the large-scale seismotectonic analysis of Delacou et al. (2004), to test the effects of gravitational body forces, coupled with rotation, on the current stress and strain fields of the western/central Alps. Numerical modelling has proved to be a powerful tool for analysing the geodynamics of different areas, such as the Himalayas (Cattin & Avouac 2000; Cattin et al. 2001), New Zealand (Liu & Bird 2002), southern Spain and northern Africa (Negredo et al. 2002), the United States and Mexico (Bird 2002), the Baikal rift zone (Lesne et al. 1998), the Basin and Range province (Hassani & Chrry 1996) and Central Europe (Griinthal & Stromeyer 1992; Golke & Coblentz 1996). Here, we use a three-dimensional model of the Alps to study the origins of the current stress and strain fields of the western/central Alpine arc.

Seismotectonic data We test our numerical models by comparing calculated strain and stress fields with those obtained from earthquake analysis (Delacou et al. 2004). Our database is a compilation of 389 reliable focal mechanisms (Mrnard 1988; Thouvenot 1996; Eva & Solarino 1998; Sue et al. 1999; Baroux et al. 2001; Kastrup, 2002), covering the entire arc of the western/central Alps, from eastern Switzerland to the Ligurian margin (Fig. 1). Local magnitudes (M1) range from 0.7 to 6.0, for earthquakes recorded between 1969 and 2000. Foci are mainly in the upper crust (first 20 km), especially in the core of the belt where no deeper earthquakes have occurred. There are a few exceptions in external areas (30 km under the Swiss Molasse basin, 25 km under the western Po plain and 20 km under the Ligurian margin).

transcurrent), principal directions of deformation (P- and T- axes) and principal stress axes (o'l, 0"2, 0"3) obtained from inversion of focal plane solutions. We used the dips of P- and T-axes to calculate an r-parameter (P-axis dip - T-axis dip) that summarizes the type of deformation, that is, compressional, extensional or transcurrent (Delacou et al. 2004). In Figure 1 the r-parameter is shown by coloured dots at epicentres, whereas interpolation provides the background colour. This large-scale regionalization reveals large zones of homogeneous deformation. In the internal zones, a continuous zone of extension follows the crest line from the southern Valais to the Argentera massif. Extension is also found in eastern Switzerland, over topographic highs, but continuity with the main zone is not proven, because the Lepontine dome is almost seismically active. Other notable features are local zones of compressional/transpressional deformation along the edges of the Alpine belt, in the eastern Helvetic domain, the front of the Belledonne massif, the front of the Digne nappe and the western Po plain. We made a map of P- and T-axis trajectories, by projecting the axes onto a horizontal plane and interpolating vectorially (Fig. 1). In internal zones, orogen-perpendicular extension prevails, T-axes striking N - S in the Valais, E - W behind the Pelvoux massif, and S W - N E behind the Argentera massif. In external zones, P-axis trajectories define a large-scale fan, convergent toward the Po plain. Orogen-perpendicular compressive axes swing through 120 ~ from a NNW trend in eastern Switzerland, to NW in front of the Belledonne massif, and SW in front of the Digne nappe. This orogen-perpendicular configuration confirms earlier results, which were based on far fewer data (Frrchet 1978; Pavoni 1986). Stress inversion methods have been applied to subsets of the focal mechanism data, to constrain the present-day stress field of the Alpine arc. For details of the analysis and calculations, see Delacou et al. (2004). The results (Fig. 1) reveal a generalized extensional stress field in the core of the belt. Orogen-perpendicular 0.3, contrasts with localized zones of transpression in external zones, where fan-shaped orogenperpendicular 0.1 converges toward the Po plain. Strike-slip faulting occurs everywhere in the belt, but is especially abundant in external zones.

S e i s m o t e c t o n i c strain a n d stress fields

The strain and stress states of the Alpine realm are defined via three parameters (Fig. 1): type of deformation (compressional, extensional or

Correlations with crustal thickness

We have used a Digital Elevation Model (DEM), GTOPO30, to calculate average Alpine

M O D E L L I N G G R A V I T A T I O N A L RE-EQUILIBRATION IN THE W E S T E R N / C E N T R A L ALPS

297

Fig. 1. Seismotectonic overview of the study area (Delacou et al. 2004). Top. Left: Digital elevation model (DEM) and geological contours. Note correspondence between topographically high areas and extensional zones of deformation (bottom map). Right: Regionalization of deformation draped on smooth DEM (radius 25 km). Note extension in inner areas that follows crest of belt and localized compressive/transpressive areas at feet of topographic gradients. Bottom. Strain and stress fields of the Alpine realm. Background colour represents type of deformation, small coloured lines represent earthquake P-axes (red) and T-axes (blue), black arrows are o-1 axes and white arrows o,3 axes. Note the orogen-perpendicular pattern of both tensile axes (in the core of the belt) and compressive axes (in external areas).

298

B. DELACOU ET AL.

topography, where each point of the grid represents average altitude within a radius of 25 km (Fig. 1). This average topography provides a proxy for topographic loading at the scale of the lithosphere, high average altitude being associated with over-thickened crust. The resulting map closely matches gravimetric maps (e.g. Masson et al. 1999), high average topography (higher than 2500 m) corresponding to strong negative Bouguer anomalies ( - 160 to - 220 mgal). On draping the map of regionalized deformation over the average DEM (Fig. 1), internal areas of high topography appear to match closely with areas where the state of strain/stress is extensional. In contrast, transpressive external zones coincide with zones of concave-upward curvature, between high mountains and low foreland. To explain the close correlation between areas of large crustal thickness (directly correlated with high average topography) and generalized Alpine extensional tectonics, we favour a geodynamic model, where the current Alpine regime is controlled, at least partly, by internal gravitational body forces. In this model, gravitational potential anomalies (GPA), driven by crustal thickness heterogeneities between internal and external zones, will induce extension in high internal zones. In response to this extensional regime, external areas will undergo compression/transpression. This kind of model will induce orogen-perpendicular extensional stress axes in high internal zones and orogenperpendicular compressional stress axes in low external zones. In what follows, we use numerical techniques to test this model of gravitational re-equilibration, alone or combined with rotation.

2.5D finite element modelling

A numerical code (2.5D thin-shell finite element code, SHELLS) has been used to model the stress and strain field of the western/central Alpine arc. Basically, this code solves for stress equilibrium and conservation of mass, given the rheology and density at each point (Bird 1989, 1999; Kong & Bird 1995). Models include three-dimensional variations in topography and thickness of crust and lithosphere. Because the code solves a momentum equation in a vertically integrated form (two-dimensional approximation), it is referred to as a 2.5D finite element method. The thin-shell approximation yields only horizontal components of the momentum equation (the vertical component being replaced by an isostatic approximation) and no vertical shear traction is considered on vertical planes (flexural strength is ignored). Material behaviour is assumed to be anelastic: thermally activated non-linear dislocation creep in the lower crust and mantle, and Mohr-Coulomb frictional plasticity in the shallow parts of crust and upper mantle (Table 1). Given values of initial surface heat flow and steady thermal conduction are used to compute a three-dimensional temperature distribution with constant but distinct heat productivity and conductivity for crust and mantle (Table 1). In summary, the assumptions and approximations of this method enable modelling of large-scale geodynamic systems over long time-scales (given the anelastic assumption, time-scales smaller than a few thousand years are not adequately modelled). For orogenic systems like the one in this study, the thin-shell

Table 1. Thermal parameters, densities and rheological parameters of models. For detailed description of rheological parameters, see Bird (1989) Parameter Heat conductivity Heat productivity Mean densities (P = 0 and T = 0) Mohr-Coulomb frictionnal parameters Fault friction coefficient Continuum friction coefficient Biot coefficient (efficacy of pore pressure) Dislocation-creep parameters ACREEP (shear stress coefficient) BCREEP (temperature coefficient) CCREEP/G/p (pressure coefficient) DCREEP (max. shear stress) ECREEP (exponent) = 1/n

Values (crust/mantle)

Units

2.7/3.2 7.27E-7/3.2E-8 2816/3332

J m -1 s -1 K -1 J m -3 s -1 kg m -3

0,03 0,85 1

2.3E9/9.5E4 4000/18 314 0/0.0171 5.00E + 08 0.333333

P a s 1/3

K K Pa -1 Pa

MODELLING GRAVITATIONAL RE-EQUILIBRATION IN THE WESTERN/CENTRAL ALPS 299 code can efficiently model the response to gravitational potential anomalies (GPA), but will not account for flexural strength (or isostatic rebound). However, even in processes such as post-glacial rebound or erosional denudation, where flexure is a significant component, the models would probably yield a stress pattern that is close to the one indicated by earthquakes (that is, extensional tectonics in internal uplifted areas). Another limitation of the 2.5D approximation is that decoupling of the stress field cannot occur at depth. Thus, it is not possible to model compression in the deep lithosphere and simultaneous extension at shallower depths. However, this limitation may not be serious, because no vertically decoupled tectonics of this kind have yet been identified at a large scale in the Alpine arc. Given the assumptions, the models in this study are limited to the analysis of the stress or strain field generated by re-equilibration of gravitational potential anomalies (GPA) and its possible combination with rotational tectonics. The boundaries of the models have been chosen to reflect the limits of the western/central Alps, as well as the limits of our seismotectonic study (Fig. 2). In the north, the boundary follows the outer edge of the Molasse Basin; in the northwest and west, the outer edges of the Jura and Subalpine chains; in the southwest, the lower Rhone valley; in the south, the Ligurian margin; and in the southeast, the Po plain. The eastern boundary of the model is an arbitrary north-south line, which is assumed to be frictionless and that limits our study area to the western/central Alps. The models in this study all have 295 cells, used in the finite element technique (Fig. 2). Another feature of the code SHELLS is that it can take into account faults (Fig. 2, Table 1). In the western/central Alps, the problem has been to identify large faults that are potentially active. Indeed, recognized seismically active faults are scarce and of limited extents. Moreover, an exhaustive list of active faults is difficult to establish, as every new local seismic swarm defines a new active fault system. In our models, we have decided to take into account large-scale inherited structures that are supposed to play an important role in the current dynamics of the studied area, these being the Pennine front and the Insubric line (Fig. 2).

Models with fixed boundaries In order to test the effects of buoyancy forces alone, models are assumed to have fixed

boundaries. The strain/stress field is generated only by contrasting gravitational potential anomalies (GPA) between the inner areas of thickened crust and external normal ones.

Isostatic model (model A) As a first step, a simple three-dimensional model has been constructed under the assumption of isostatic equilibrium (Figs 3 and 4). From the surface topography (taken from the GTOPO30 DEM data, smoothed at the mesh spacing size) and the surface heat flow (compiled from the European Geotraverse experiments; Blundell et al. 1992), SHELLS calculates routinely the three-dimensional structure of the crust and the lithosphere that satisfies isostatic equilibrium (Fig. 3) and steady-state thermal conduction, by taking into account the densities and thermal properties of crust and mantle (Table 1). We assume that all boundary nodes are stationary. For model A, the calculated stress field is characterized by orogen-perpendicular extension in regions of high topography in the core of the belt and by orogen-perpendicular compression in external zones (Fig. 4). This pattern results from equilibration between regions of positive GPA in the inner areas, where high topography correlates with large crustal thickness (according to the assumption of isostatic equilibrium) and regions of normal GPA (near zero) in external zones, where altitudes are small and the Moho is close to its normal depth (around 30 km). This configuration results in an extensional stress state in the core of the belt, tending to reduce the over-thickened crust, and a compressional stress state in external regions. In terms of strain rate (Fig. 4), a belt of horizontal stretching appears to follow the high topography, especially on its external side, from the Aar massif to the Argentera massif. Extensional strain rates are between 1 and 7 x 10 - 1 6 s - 1 . Compressional strain rates are about 2 x 10 - ~ 6 s - J in external zones. They seem to be guided by the fixed boundaries of the model, reaching a maximum in front of the Jura, the Po plain and the Rhone valley. This could be explained by GPA equilibration, whereby crustal thickness decreases over the whole system as far as the boundaries of the model, where it creates compression. In reality, external boundaries (that can be considered as fixed, far away from the Alps) are not as sharp as they are in our models, so that shortening should be more distributed.

300

B. DELACOU ET ALo 4~

5~

6OE

7OE

8~

9~

____._,,,=IO~ ....

~48~

47~

46~

~ilan

44~

Nice

43~ " 4~

'

.......... 5OE

6OE

IIII 7OE

'

I 8~

,, 9~

,. . . . . . . 10~

Fig. 2. Grid and configuration in our finite element models. Models have 295 elements, regularly spaced in the area of the western/central Alps, Bold lines inside models represent faults: Pennine Front (PF), Simplon fault (Si) and Periadriatic Line (PL). Aa, Aar external crystalline massif; Ar, Argentera external crystalline massif; B, BeUedonne external crystalline massif; Br, Brian~onnais area; Di, Digne nappe; G, Grisons; H, Helvetic zones; J, Jura fold and thrust belt; Le, Lepontine dome; Li, Ligurian margin; M, Molasse basin; Mb, Mont-Blanc external crystalline massif; NV, Northern Valais; Pi, Piemontais area; Pe, Pelvoux external crystalline massif; Po, Po plain; SV, Southern Valais.

The surface velocity field also follows the shape of the model, velocities reaching about 0.15 to 0.25 mm a -~ in directions (NW to SW) that are perpendicular to the belt. In internal zones, stretching leads to southeast-directed surface velocities, which reach 0.3 mm a -a in the northern part (Valais).

Motion on faults is mainly manifest as extensional reactivation of the Pennine front. Slip perpendicular to this fault zone reaches 0 . 7 m m a -z on its northern segment and decreases progressively toward the south. Near the Mediterranean, the fault appears to be locked. The Periadriatic line does not slip at all.

M O D E L L I N G G R A V I T A T I O N A L RE-EQUILIBRATION IN THE W E S T E R N / C E N T R A L ALPS

301

MODEL CONFIGURATIONS

Fig. 3. Model configuration. Elevation and surface heat flow are common to all models. Note differences in crustal thicknesses between isostatic models (models A and C), where Moho depth is directly related to topography, and realistic models (models B and D), characterized by a Moho dipping toward the E/SE on the European side of the belt, and a complex geometry at the eastern Po plain boundary. Moho geometry is taken from Waldhauser et al. (1998).

302

B. DELACOU ET AL. MODEL A ISOSTATIC MODEL FIXED BOUNDARIES

Fig. 4. Model A: Model with isostatic three-dimensional crustal geometry (see Fig. 3) and fixed boundaries. This starting model represents tectonic response of Gravitational Potential Anomalies (GPA) in a simple model of the western/central Alps (see text for explanations).

M O D E L L I N G G R A V I T A T I O N A L RE-EQUILIBRATION IN THE W E S T E R N / C E N T R A L ALPS

MODEL B REALISTIC CRUSTAL MODEL FIXED BOUNDARIES

Fig. 5. Model B: Model with realistic three-dimensional crustal geometry (see Fig. 3) and fixed boundaries. This model exhibits a more complex tectonic response than model A, as a result of complex crustal geometry (see text for explanation).

303

304

B. DELACOU ETAL.

Realistic crustal m o d e l ( m o d e l B)

For a more realistic three-dimensional crustal structure, we have constructed a model (model B, Fig. 5), where the Moho geometry (Fig. 3) has been interpreted from wide-angle seismic experiments (Waldhauser et al. 1998). Given the topography, surface heat flow and realistic crustal geometry, lithosphere thickness is calculated, so as to respect the thermal properties of crust and mantle (assuming steady thermal conduction). This results in a complex threedimensional geometry (Fig. 3), where the highest altitudes do not directly overlie the crustal root. The latter reaches a depth of about 50-55 km at a point that appears to be shifted toward the SSE with respect to large-scale topography. A consequence of such a setting is that GPAs do not correlate in a simple manner with topography (as they do in model A), but depend on the whole three-dimensional crustal structure. In regions of high topography and relatively shallow Moho depth, the GPA is positive and extension is expected in the anomalous lithosphere, whereas in regions of deep Moho and relatively moderate topography, the GPA is negative and compression is expected. Thus, the resulting stress field (Fig. 5), appears to be more complex than in model A. A general trend, from inner extension to outer compression, is still present (as in model A), but with regional variations. This is so in the Jura, where extension now occurs in the northern inner part, in the southwestern external Alps, where there is a mix of compression and extension, and in the northwestern Po plain, where extension is observed. These can be considered as local effects of crustal thickness variations, which were not present in model A. In the southwestern Alps, where focal mechanisms are of mixed type (compressional, extensional and transcurent), the model seems to fit the observations. In the northern part of the internal zone, the general orogen-perpendicular extension is cross-cut by an E - W band of N - S extension. This correlates with the northern edge of the Apulian crustal wedge (Fig. 3), which may be correlated to the Val d'Aosta extensional fault zone (Bistacchi et al. 2001). Three bands of high strain rate (up to 3 x 10 -15 s -1) can be recognized: a band of W N W - E S E shortening in the external zones beyond the Belledonne and Mont-Blanc massifs, a band of fan-shaped stretching that follows the topographic high (with two peaks in the Aar and Pelvoux regions), and E - W shortening in the western Po plain. These bands correlate fairly well with the seismotectonic setting and the concentrations

of epicentres (see Fig. 1). There are three zones of high surface velocities: NW-directed velocities of up to 0.75 mm a -~ in external zones, E-directed velocities of up to 0.7 mm a-~ in the southern inner area, and a complex zone over the Aar massif. The pattern of slip along faults is more complex than in model A. Extensional fault slip (up to 1.3 m m a -~) occurs along the middle segment of the Pennine front, whereas the northern branch seems to accommodate complicated local movements, due to dextral transtension in the Simplon area and compression in the Valais. The southern branch of the Pennine front is now accommodating local thrusting (less than 0.3 mm a-~), decreasing toward the south to reach a locked state near the Argentera massif. The Periadriatic line is almost inactive, except along its western segment, where dextral slip rates reach 0.15 mm a -

Models with rotational boundaries Rotation may have played an important role in the dynamics of the western/central Alps, since at least Oligo-Miocene times (Gidon 1974; Anderson & Jackson 1987; M6nard 1988; Vialon et al. 1989; Thomas et al. 1999; Collombet et al. 2002). On the strength of GPS monitoring, involving French, Swiss and Italian stations (Calais et al. 2002), it is claimed that the Apulian promontory is rotating anticlockwise with respect to stable Europe at a rate of 0.52~ around a pole located at 45.36~ 9.10~ (near Milan). In order to test the effects of such a rotation on the strain and stress field of the western/central Alpine arc, boundary nodes for the Po plain have been given appropriate velocities in model C (same three-dimensional crustal structure as model A) and model D (same three-dimensional crustal structure as model B). In terms of stress (Figs 6 and 7), results for the rotation models appear to be quite similar to those for fixed models at a large scale. Orogenperpendicular extensional stress is present in the internal zones, and orogen-perpendicular fan-like compression in the external zones (at least in the northwestern part). Only the regional/local pattern of stress axes is different from that of the fixed models. Thus rotation induces frontal compression at the eastern edge of the SW Alps and near the Po plain. This is especially true for model C, where compressional axes follow the rotational motion of boundary nodes. For model D, stress axes deviate less than for model C. This may be because GPAs are more variable in

MODELLING GRAVITATIONAL RE-EQUILIBRATION IN THE W E S T E R N / C E N T R A L ALPS

305

MODEL C ISOSTATIC MODEL ROTATIONAL BOUNDARIES

Fig. 6. Model C: Model with isostatic three-dimensional crustal geometry and rotational Po plain boundary nodes. See Figure 4 for comparison. Differences between models A and C are only due to rotational Po plain boundary. Curved arrow on surface velocity map indicates rotation pole (see text for explanation).

306

B. D E L A C O U ETAL.

MODEL D REALISTIC CRUSTAL MODEL ROTATIONAL BOUNDARIES

Fig. 7. Model D: Model with realistic three-dimensional crustal geometry and rotational Po plain boundary nodes. See Figure 5 for comparisons. Differences between models B and D are only due to rotational Po plain boundary (see text for explanation).

MODELLING GRAVITATIONAL RE-EQUILIBRATION IN THE WESTERN/CENTRAL ALPS 307 the non-isostatic model (model D), so that body forces are more dominant. In terms of strain rate, as well as stress, axes are almost the same as for fixed models and only small regional reorientations are observed. The main differences are at boundaries. For example, at the south Ligurian boundary, anticlockwise rotation induces large shortening/ Surface velocities of up to 1.4-1.5 m m a concentrate in internal zones and appear to be strongly linked to the rotational boundary. Velocity vectors follow this rotation, pointing more to the south than for fixed models. Rotation seems to have little effect on velocities of external zones. Slip directions along faults are more southdirected than those of fixed models. For the Pennine front, this implies S-verging stretching of up to 1.7 mm a -1 in the northern segment (model C), dextral transtension in the middle seg~ ment (1.1 + 2 mm a-1 for model C, 2.1 mm a for model D), and dextral transpression in the southern segment (1.4mm a -~ for model C, 1.2 mm a-1 for model D). Another major difference with fixed models is the small dextral motion on the Periadriatic line (up to 0.22 mm a-1 for model C and 0.4 mm a - for model D).

Geodynamic implications Numerical modelling and comparison with largescale seismotectonic analysis have shown that body forces play a major role in determining the current stress field of the western Alpine arc. Balance of GPAs explains the orogenperpendicular contrasted stress field in the western/central Alps (extensional in the core of the belt, locally compressional at the periphery). The role of rotational boundary forces is less obvious, as only local stress reorientations appear in our models. Nevertheless, rotation models seem suitable to explain dextral strikeslip faulting along the external zones (from the northern Valais to the Argentera massif). In addition, our results are consistent with GPS studies (e.g. Calais et al. 2002) that give velocities of about 1 mm a-1 within the Alpine realm, compatible with velocities of 1-1.5 m m a -~ obtained in the models of this study. More precisely, GPS studies reveal extension in the core of the belt, including lengthening of the line Lyon-Turin (0.5 _+ 0.9 mm a-1 to the SE at La Feclaz in the Subalpine chains and 1.4 +_ 0.4 mm a-~ to the SE at Modane (in the internal Vanoise area). This geodetic stretching correlates well with the values obtained for the core of the belt by seismotectonic analysis and numerical modelling. Moreover, GPS results

also indicate shortening in the western Po plain (1.0-0.5 mm a -1 of E - W shortening between Modane and Turin) and in Provence (1.40.5 mm a-z of N - S shortening between Grasse and Turin). Despite this qualitative agreement with the results of our modelling and seismotectonic analysis, more detailed comparisons cannot be made, because the GPS data still have insufficient resolution. This study addresses the consequences of ongoing convergence between Europe and Africa. The convergence velocity is estimated at 3 + 8 mm a-~ in a N to NW direction at the longitude of the Alps (Argus et al. 1989; Demets et al. 1990; Demets et al. 1994; Albarello et al. 1995; Cr6taux et al. 1998; Kreemer & Holt 2001; Nocquet 2002). It could be taken up in different areas between the European and African stable continents, such as Northern Africa, the Apennines, the Dinarides or the Calabrian subduction zone. In the vicinity of the western/central Alpine arc, the interaction between boundary forces and body forces is still a matter for debate. Studies, such as the one by Thatcher et al. (1999) in the Basin and Range province, show that gravitational extensional tectonics can interact with boundary conditions, leading to reorientation of extensional axes parallel to plate tectonic directions. However, in our study, direct effects of plate tectonics are less useful to explain the stress field of the western/central Alps, which appears to be controlled mostly by internal body forces. A more detailed analysis of the possible interactions between boundary forces and body forces in the Alpine belt would require a detailed three-dimensional geometry of the models (accounting for lithospheric complexities), a fully three-dimensional finite element code, as well as more constraints on boundary conditions between Apulian and European microplates. Recent tomographic studies (Lippitsch 2002) have yielded a complex three-dimensional geometry at great depth, which has been interpreted in terms of lithospheric slabs, possibly detached in the western Alps and subvertical under the central Alps. These lithospheric structures cannot be modelled by the techniques used in this study. Their consequences for the current stress field and recent tectonics of the Alpine arc remain to be analysed.

Conclusions A seismotectonic investigation along the entire arc of the western/central Alps has revealed contrasting stress regimes. Within a zone of extension that follows the arcuate crest line

308

B. DELACOU ETAL.

from the southern Valais to the Argentera massif, extensional axes are perpendicular to the orogen. Compression is limited to the external zones, where compressional axes are also perpendicular to the orogen. Strike-slip faulting occurs in both external and internal zones, but is particularly abundant in the latter, where it is right-lateral, all the way from the northern Valais to the Durance fault (northwest of the Argentera massif). This well-defined seismotectonic stress state, comparable to the ones computed with 2.5D numerical modelling, highlights the essential role of gravitational body forces, which are able to produce orogen-perpendicular extension in the topographic highs and resulting orogenperpendicular compression at the periphery. The role of rotation, which has been tested in our models, is more ambiguous, but could explain the arcuate right-lateral faulting prevailing in the external zones. In a context of ongoing far-field convergence between the European and African plates, no direct evidence of collision has been found in the Alpine realm, either by seismotectonic analyses, numerical modelling, or GPS studies. This suggests that the current stress field in the western/central Alps is post-collisional. Neuch~tel University and the Swiss National Science Foundation (grant # 21-61684.00) supported this study. We wish to thank Peter Bird for free online access to his finite element codes (http://element.ess.ucla.edu/); Jean Chrry and Rob Butler for their reviews; Peter Cobbold for his great help to the manuscript; and Nicole B&houx, Jean-Mathieu Nocquet and Riad Hassani for fruitful discussions.

References AHORNER, Z., MURAWSKI,H. & SCHEINDER,G. 1972. Sismotektonische Traverse von der Nordsee bis zum Apennin. Geologische Rundschau, 61, 915-942. ALBARELLO, D., MANTOVANI, E., BABBUCCI, D. & TAMBURELLI,C. 1995. Africa-Eurasia kinematics main constraints and uncertainties. Tectonophysics, 243 (1-2), 25-36. ANDERSON, H. & JACKSON, J. 1987. Active tectonics in the Adriatic region. Geophysical Journal of the Royal Astronomical Society, 91, 937-983. ARGUS, D. F., GORDON, R. G., DEMETS, C. & STEIN, S. 1989. Closure of the Africa-Eurasia-north America plate motion circuit and tectonics of the Gloria fault. Journal of Geophysical Research, 94, 5585-5602. BADA, G., HORVATH, F., CLOETINGH, S., COBLENTZ, D. & TOTH, T. 2001. Role of topography-induced gravitational stresses in basin inversion: the case study of the Pannonian basin. Tectonics, 20 (3), 343 -363.

BAROUX, E., BI~THOUX, N. & BELLIER, O. 2001. Analyses of the stress field in southeastern France from earthquake focal mechanisms. Geophysical Journal International, 145, 336-348. BECKER, A. 2000. The Jura Mountains - an active foreland fold-and-thrust belt? Tectonophysics, 321 (4), 381-406. B1RD, P. 1989. New finite element techniques for modeling deformation histories of continents with stratified temperature-dependent rheology. Journal of Geophysical Research, 94 (B4), 3967-3990. BIRD, P. 1999. Thin-plate and thin-shell finite-element programs for forward dynamic modeling of plate deformation and faulting. Computers & Geosciences, 25, 383-394. BIRD, P. 2002. Stress-direction history of the western United States and Mexico since 85 Ma. Tectonics, 21 (3), 5-12. BISTACCH1, A., DAL PIAZ, G. V., MASSIRONI, M., ZATTIN, M. • BALESTRIERI, M. L. 2001. The Aosta-Ranzola extensional fault system and Oligocene - Present evolution of the AustroalpinePenninic wedge in the northwestern Alps. International Journal of Earth Sciences, 90 (3), 654-667. BISTACCHI, A., EVA, E., MASSIRONI, M. & SOLARINO, S. 2000. Miocene to Present kinematics of the NW-Alps: evidences from remote sensing, structural analysis, seismotectonics and thermochronology. Journal of Geodynamics, 30, 205 -228. BLUNDELL, D., FREEMAN, R. & MUELLER, S. 1992. A

Continent Revealed: The European Geotraverse. Cambridge University Press. Oxford, London, Edinburgh, Boston, Melbourne. BURG, J. P., SOKOUTIS, D. & BONINI, M. 2002. Modelinspired interpretation of seismic structures in the Central Alps: Crustal wedging and buckling at mature stage of collision. Geology, 30, 643-646. BURKHARD, M. 1990. Aspects of the large-scale Miocene deformation in the most external part of the Swiss Alps (Subalpine Molasse to Jura fold belt). Eclogae Geologicae Helvetiae, 83 (3), 559-583. BURKHARD, M. & SOMMARUGA, A. 1998. Evolution of the western Swiss Molasse basin: structural relations with the Alps and the Jura belt. In: MASCLES, A., PUIGDEF.~BREGAS, C., LUTERBACHER, H. P. & FERNANDEZ, M. (eds)

Cenozoic Foreland Basins of Western Europe. Geological Society of London, Special Publications, 134, 279-298. BUTLER, R. W. H. 1992. Thrusting patterns in the NW French Subalpine chains. Annales Tectonicae, 6, 150-172. CALAIS, E., NOCQUET, J. M., JOUANNE, F. & TARDY, M. 2002. Current strain regime in the Western Alps from continuous Global Positioning System measurements, 1996-2001. Geology, 30, 651-654. CATTIN, R. & AVOUAC,J. P. 2000. Modeling mountain building and the seismic cycle in the Himalaya

MODELLING GRAVITATIONAL RE-EQUILIBRATION IN THE WESTERN/CENTRAL ALPS 309 of Nepal. Journal of Geophysical Research, 105 (B6), 13 389-13 407. CATTIN, R., MARTELET, G., HENRY, P., AVOUAC,J. P., DIAMENT, M. & SHAKYA,T. R. 2001. Gravity anomalies, crustal structure and thermo-mechanical support of the Himalayas of Central Nepal. Geophysical Journal International, 147, 381-392. CHAMPAGNAC, J. D., SUE, C., DELACOU, B. & BURKHARD, M. 2003. Brittle orogen-parallel extension in the internal zones of the Swiss Alps (south Valais). Eclogae Geologicae Helvetiae, 96, 325 -338. CHAMPAGNAC, J. D., SUE, C., DELACOU, B. & BURKHARD, M. 2004. Brittle deformation in the inner northwestern Alps: from early orogenparallel extrusion to late orogen-perpendicular collapse. Terra Nova, 16(4), 232-242, doi: 10.1111/j. 1365-31 21.2004.00555.x. CHOUKROUNE, P., BALLI~VRE, M., COBBOLD, P., GAUTIER, Y., MERLE, O. & VUICHARD, J. P. 1986. Deformation and motion in the western alpine arc. Tectonics, 5 (2), 215-226. COLLOMBET, M., THOMAS, J. C., CHAUVIN, A., TRICART, P., BOUILLIN, J. P. & GRATIER, J. P. 2002. Counterclockwise rotation of the western Alps since the Oligocene: new insights from paleomagnetic data. Tectonics, 21, 352-366. COWARD, M. & DIETRICH, D. 1989. Alpine tectonics: an overview. In: COWARD, M., DmTRICH, D. & PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 1-29. CRETAUX, J.-F., SOUDARIN, L., CAZENAVE, A. & BOUILLI~ F. 1998. Present-day tectonic plate motions and crustal deformations from the DORIS space system. Journal of Geophysical Research, 103, 30167-30181. DAVIES, J. H. & VON BLANCKENBURG,F. 1995. Slab breakoff: A model of lithosphere detachment and its test in the magmatism and deformation of collisional orogens. Earth and Planetary Sciences Letters, 129 (1-4), 85-102. DELACOU, B., SUE, C., CHAMPAGNAC, J. D. & BURKHARD, M. 2004. Present-day geodynamics in the bend of the western and central Alps as constrained by earthquake analysis. Geophysical Journal International, doi: 10.1111/j. 1365-1246X. 2004.02320.x. DEMETS, C., GORDON,R. G., ARGUS,D. F. & STEIN, S. 1990. Current plate motions. Geophysical Journal International, 101, 425-478. DEMETS, C., GORDON, R. G., ARGUS,D. F. & STEIN, S. 1994. Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions. Geophysical Research Letters, 21, 2191-2194. DEWEY, J. F., HELMAN, M. L., TURCO, E., HUTTON, D. W. H. & KNOTT, S. D. 1989. Kinematics of the western Mediterranean. In: COWARD, M., DIETRICH, D. & PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 265-283. DUCHI~NE, S., BLICHERT-TOFT, J., LUAIS, B., TI~LOUK, P., LARDEAUX, J. M. & ALBARI~DE, F. 1997. The L u - H f dating of garnets and the ages

of the Alpine high-pressure metamorphism. Nature, 387, 586-589. EVA, E., PASTORE, S. & DEICHMANN, N. 1998. Evidence for ongoing extensional deformation in the Western Swiss Alps and thrust-faulting in the southwestern Alpine foreland. Journal of Geodynamics, 26 (1), 27-43. EVA, E. & SOLARINO, S. 1998. Variations of stress directions in the western Alpine arc. Geophysical Journal International, 135, 438-448. FRECHET, J. 1978. Sismicitg du sud-est de la France et une nouvelle mdthode de zonage sismique. Thbse de doctorat d'Etat, Universit6 des Sciences Technologiques et Medicales, Grenoble. FRISCH, W., DUNKL, I. & KUHLEMANN, J. 2000. Post-collisional orogen-parallel large-scale extension in the Eastern Alps. Tectonophysics, 327 (3-4), 239-265. FRY, M. 1989. Southwestward thrusting and tectonics of the Western Alps. In: COWARD,M., DIETRICH, D. & PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 83-109. FI3GENSCHUH, B., LOPRIENO, A., CERIAN1, S. & SCHMID, S. 1999. Structural analysis of the Subbrianqonnais and Valais units in the area of Mofitiers (Savoy, Western Alps): paleogeographic and tectonic consequences. International Journal of Earth Sciences, 88, 201-218. GIDON, M. 1974. L'arc alpin a-t-il une origine tourbillonnaire? Comptes Rendus de l'Acaddmie des Sciences de Paris, 278, 21-24. GOLKE, M. & COBLENTZ, D. 1996. Origins of the European regional stress field. Tectonophysics, 266 (1-4), 11-24. GRt3NTHAL, G. & STROMEYER, D. 1992. The recent crustal stress field in Central Europe: trajectories and finite element modeling. Journal of Geophysical Research, 97 (B8), 11 805-11 820. HASSANI, R. & CHI~RY, J. 1996. Anelasticity explains topography associated with Basin and Range normal faulting. Geology, 24 (12), 1095-1098. KASTRUP, U. 2002. Seismotectonics and stress-field variations in Switzerland. PhD thesis, ETHZ, Ziirich. KONG, X. & BIRD, P. 1995. SHELLS: A thin-shell program for modeling neotectonics of regional or global lithosphere with faults. Journal of Geophysical Research, 100 (Bll), 22 129-22 131. KREEMER, C. & HOLT, W. E. 2001. A no-net-rotation model of present day surface motion. Geophysical Research Letters, 28, 4407-4410. LAUBSCHER, H. 1991. The arc of the Western Alps today. Eclogae Geologicae Helvetiae, 84 (3), 631-659. LESNE, O., CALAIS, E. & DEVERCHERE,J. 1998. Finite element modeling of crustal deformation in the Baikal rift zone: new insights into the activepassive rifting debate. Tectonophysics, 289, 327-340. LIPPITSCH, R. 2002. Lithosphere and upper mantle P-wave velocity structure beneath the Alps by high-resolution teleseismic tomography. PhD thesis, ETHZ, Zurich.

310

B. DELACOU ETAL.

LIu, Z. & BIRD, P. 2002. Finite element modeling of neotectonics in New Zealand. Journal of Geophysical Research, 107 (B 12), 2328-2346. MANCKTELOW, N. S. 1992. Neogene lateral extension during convergence in the Central Alps: evidence interrelated faulting and backfolding around the Simplonpass (Switzerland). Tectonophysics, 215, 295-317. MASSON, F., VERDUN,J., BAYER, R. & DEBEGLIA,N. 1999. Une nouvelle carte gravim&rique des Alpes occidentales et ses cons6quences strncturales et tectoniques. Comptes Rendus de l'Acaddmie des Sciences de Paris, 329, 865-871. MAURER, H., BURKHARD, M., DEICHMANN, N. & GREEN, G. 1997. Active tectonism in the central Alps: Contrasting stress regimes north and south of the Rhone Valley. Terra Nova, 9, 91-94. M~NARD, G. 1988. Structure et cindmatique d'une cha~ne de collision: Les Alpes occidentales et centrales. Th~se de Doctorat d'6tat, Universit6 Joseph Fourier, Grenoble. NEGREDO, A. M., BIRD, P., SANZDE GALDEANO,C. & BUFORN, E. 2002. Neotectonic modeling of the Ibero-Maghrebian region. Journal of Geophysical Research, 107 (Bll), 2292. NOCQUET, J.-M. 2002. Mesure de la d~formation crustale en Europe occidentale par Gdoddsie spatiale. Th~se de doctorat, Universit6 de Nice. NOCQUET, J. M. & CALAIS, E. 2003. Crustal velocity field of western Europe from permanent GPS array solutions, 1996-2001. Geophysical Journal International, 154 (1), 72-88. NOCQUET, J. M. & CALAIS, E. 2004. Geodetic measurements of crustal deformation in the Western Mediterranean and Europe. Pure & Applied Geophysics, 161 (3), 661-681. PAVONI, N. 1961. Faltung durch horizontal verschiebung. Eclogae Geologicae Helvetiae, 54, 515-534. PAVONI, N. 1986. Regularities in the pattern of major fault zones of the earth and the origin of arcs. In: WEZEL, F. C. (ed) Origin of Arcs. Elsevier, Amsterdam, 63-78. POGNANTE, U. 1991. Petrological constraints on the eclogite- and blueschist-facies metamorphism and P-T-t paths in the Western Alps. Journal of Metamorphic Geology, 9, 5-17. RATSCHBACHER, L., FRISCH, W., LINZER, H.-G. & MERLE, O. 1991. Lateral extrusion in the Eastern Alps; part 2: structural analysis. Tectonics, 10 (2), 257-271. SACHSENHOFER, R. F., KOGLER, A., POLESNY, H., STRAUSS, P. & WA6REICH, M. 2000. The Neogene Fohnsdorf Basin: basin formation and basin inversion during lateral extrusion in the Eastern Alps (Austria). International Journal of Earth Sciences, 89 (2), 415-430. SCHMID, S. M. & KISSLING, E. 2000. The arc of the western Alps in the light of geophysical data on deep crustal structure. Tectonics, 19 (1), 62-85. SPALLA, M. I., LARDEAUX, J. M., DAL PIAZ, G. V., GOSSO, G. & MESSIGA, B. 1996. Tectonic

significance of Alpine Eclogites. Journal of Geodynamics, 21, 257-285. STAMPFLI, G. M., MOSAR, J., MARQUER, D., MARCHANT, R., BAUDIN, Y. & BOREL, G. 1998. Subduction and obduction processes in the Swiss Alps. Tectonophysics, 296 (1-2), 159-204. SUE, C. 1998. Dynamique actuelle et rdcente des Alpes occidentales internes - Approche structurale et sismologique. Th~se de doctorat, Universit~ Joseph Fouriere, Grenoble. SUE, C., THOUVENOT, F., FRI~CHET, J. & TRICART, P. 1999. Widespread extension in the core of the western Alps revealed by earthquake analysis. Journal of Geophysical Research, 104 (Bll), 25 611-25 622. SUE, C. & TRICART, P. 2002. Widespread Post-Nappe Normal Faulting in the Internal Western Alps: A New Constraint on Arc Dynamic. Journal of the Geological Society, 159, 61-70. SUE, C. & TRICART, P. 2003. Neogene to current normal faulting in the inner western Alps: a major evolution of the late alpine tectonics. Tectonics, 22 (5). THATCHER, W., FOULGER, G. R., JULIAN, B. R., SVARC, J., QUILTY, E. & BAWDEN, G. W. 1999. Present-day deformation across the Basin and Range Province, Western United States. Science, 283, 1714-1718. THOMAS, J. C., CLAUDEL, M. E., COLLOMBET, M., TRICART, P., CHAUV1N, A. & DUMONT, Y. 1999. First paleomagnetic data from the sedimentary cover of the French Penninic Alps: evidence for Tertiary counterclockwise rotations in the Western Alps. Earth and Planetary Sciences Letters, 171 (4), 561-574. THOUVENOT, F. 1996. Aspects gdophysiques et structuraux des Alpes occidentales et de trois autres orogdn~ses (Atlas, Pyrdndes, Oural). Th~se de doctorat d'Etat, Universit6 Joseph Fourier, Grenoble. TRICART, P., SCHWARTZ, S., SUE, C., POUPEAU, G. & LARDEAUX, J.-M. 2001. La d~nudation tectonique de la zone ultradauphinoise et l'inversion du front brian~onnais au sud-est du Pelvoux (Alpes Occidentales): une dynamique mioc6ne ~ actuelle. Bulletin de la Socidtd Gdologique de France, 172 (1), 49-58. VIALON, P., ROCHETTE, P. & Mf~NARD, G. 1989. Indentation and rotation in the Alpine arc. In: COWARD, M., DIETRICH, D. & PARK, R. (eds) Alpine Tectonics. Geological Society, London, Special Publications, 45, 329-338. VIGNY, C., CHERY, J. ETAL. 2002. GPS network monitors the Western Alps' deformation over a five-year period: 1993-1998. Journal of Geodesy, 76 (2), 63 -76. WALDHAUSER, F., KISSLING, E., ANSORGE, J. & MUELLER, S. 1998. Three-dimensional interface modelling with two-dimensional seimic data: the Alpine crust-mantle boundary. Geophysical Journal International, 135, 264-278.

Index Note: Page numbers in italics refer to figures; those in bold refer to tables. Abitibi, Canada, quartz veins 68, 76-77 accommodation zone 269 Afar area 76 aftershocks 25 Aheim dunite 148, 150, 151 allochthon 71 Alpine Belt see Western/Central Alps Alpine Chain structural domain 233-235 tectonic sketch map 231 Amindav samples 47 ammonium chloride solution 92 amphibole brittle deformation 121 BSE images 108 crystallographic preferred orientation (CPO) 100, 115-116 deformation mechanisms 97, 120-121 grain, SPO 114 grain boundary preferred orientation (GBPO) 114 grain size 113 modal proportions 101 point and contour pole figures 119 representative microprobe analyses 111 amphibole porphyroclasts compositions 108, 109 in protomylonites 108 modal proportion 105 amphibole-plagioclase boundaries, degree of preferred orientation 114 amphibole- plagioclase system, rheology 121 - 122 amphibolite-facies metagabbro, microstructure 12 metamorphism 178-179 retrogression 190 amphibolitic deformation 186 kinematics of 185 anatectic rocks, Domaine Sud-Armoricain (DSA) 209-217 Anita Bay dunite 148, 150, 151 Aosta Fault 2 Apennines, Italy 71-72, 71 Archaean granulites, steep extrusion of 193-204 area increase 20-22 simulation 18-19 attachment zone 269 augen mylonite 101, 104-105, 105 augen mylonitic hornblendes 121

Badshot dolomite 40, 43-44, 46, 55, 55, 57, 58-59, 59, 61, 62 differential stress-axial strain curves 41 experimental results 38, 39 Badshot Formation 38, 53 banded mylonite 101 Belle-Ile Group 207-208 Bergen Arcs, Norway geological map 177 geological setting 176 syneclogite deformation in 175-192 Bohemian Massif 98 bond breaking 13 bookshelf-type deformation 185,185 boundary conditions 16, 18 boundary effects 23 boundary particles 13 box boundaries 23 breaking springs 13, 13 breaking strength 17-18 breaking thresholds 16 British Columbia 38 brittle deformation 11, 46, 76, 135 inferred from quartz recrystallized grain size 127-142 of dolomite, textural controls on 51-66 brittle-ductile transition zone 219 brittle faulting 22 to cataclastic flow 51-66 bulk deformation rate 83 bulk rock chemistry 105-106 bulk rock composition, representative analyses 107 buoyancy forces and shear stresses 189 calcite 64-65 automated EBSD map 4 crystallization 83 deformation 5-8, 83 slip systems 5 veins 71-72, 71 calcite monocrystal indentation 91 rate limiting process 90-92 pressure solution creep (PSC) rates on 81-95 typical hole in 87 Caledonian contractional structures 187 Caledonian orogenesis 176, 186, 190 tectonic model 187 Cambro-Ordivician metamorphic event 101

312

INDEX

Cambro-Ordovician protolith 98 Canavese fault 260-262, 266 Carboniferous deformation 217 Carboniferous leucogranites 219-220 Carboniferous metamorphism 99 Caribbean-South American plate boundary 266 cataclasite 22, 252 cataclastic deformation 71, 83 cataclastic flow, brittle faulting to 51-66 catastrophic failure 22 Cesi-S. Martino segment 27, 29, 29-32, 30-31, 34 Channel 5 software 3 chemical potential 92 chemically induced grain boundary migration (CIGM) 98, 120 chlorite-Phengite local equilibria 129 chrysotile 71 Coble creep 144, 154 Colfiorito Fault System (CoFS) 30 fault interaction 31-33 geological setting 26-27, 26 main fault segments 28 schematic structural map 28 segmentation and interaction of normal faults 25-36 stereographic projections 28 throw profiles 27-29 collision zones lithospheric-scale analogue modelling of 277-294 with pre-existing weak zone 277-294 collisional eclogitic deformation 186 collisional exhumation models 189-190 compaction event 19 compressive stresses 16, 18, 19, 20, 22 conjugate faults 12 conjugate hybrid shear fractures 12 conjugate shear fractures 16, 23 continental fault zones at strike-slip or oblique-slip plate boundaries 249 continental lithosphere, subduction of 164-166 continental subduction and exhumation 159-174 Coulomb's failure criterion 38, 44 Coulomb's relationship 46 crack length, maximum 77 crack-seal fibrous veins, computer models 69 crack-seal inclusion bands, geometry of 70 crack-seal increments 70 regularity of 69 crack-seal mechanism and fluid valving 69 kinematic numerical model 68 crack-seal patterns 67-79 data analysis 72, 73 Fast Fourier Transform (FFF) analysis 75 statistical analysis 72-76 statistical properties 69

crack-seal processes, indicators of 67 crack-seal samples 69-72 statistical properties 69 crack-seal sequence 70 autocorrelation function 76 cycles of breakage events and periods of quiescence 76 interruptions 70 crack-seal structures 77 fossil record of non-correlated crack sequences 77 crack-seal textures 70 crack-seal thickness distribution 76 crack-seal veins displacement-tracking conditions 68 origin and mineralogy 69 statistical properties 76 crack-seals stress release variations recorded by 69 uncorrelated distribution 76 crack surface areas 77 crack thickness distribution 74, 76 exponential distribution 72-74 variations 76-77 crack thickness histograms, characteristic length scale 72, 72 Cretaceous limestone 27 critically resolved shear stress (CRSS) 8 crustal evolution from midcrustal strain localization to detachment formation 140 crustal rocks, stress release fluctuations in 67-79, 68 crystal indentation diffusion-controlled 82 kinetic-controlled 82 crystal plastic deformation 83 crystal plastic processes 65 crystallographic dispersion 9 crystallographic pole dispersion 5, 8 crystallographic preferred orientation (CPO) 3, 100, 115-116, 115, 144, 146, 149, 155, 180 Cyclades, Greece 127-128 Cycladic Blueschists 129 decoupling and its relation to strain partitioning in continental lithosphere 249-276 deformation and melting, synergistic effects 205-225 deformation heterogeneity 228 deformation history 9, 19 fracture patterns as indicators of 11-24 vein patterns as indicators of 11-24 deformation mechanism map 138 deformation-metamorphism interactions 230-237 Dervio-Olgiasca TMU 239-240

INDEX detachment fault 134 detachment formation, crustal evolution from midcrustal strain localization to 140 dextral en-rchelon array 27 dextral shearing under amphibolite- to greenschist-facies conditions 261 dextral transtensive deformation 212 diamond grain structures 144 differential stress-axial strain curves 59 differential stress curve 15 Digital Elevation Model (DEM) 281-283, 296 dimethylaniline (DMA) 83 dip-slip kinematics 30 discrete particle model 13 dislocation creep 119, 155, 215 and piezometer parameters 131 flow laws 131 dispersion axes 6, 7 - 8 displacement transducer (DCDT) 52 dolomite peak differential stress 44 peak strength, correlation coefficients 42, 42 peak strength controls 41-45 textural controls on brittle deformation 37-49, 51-66 textural controls on peak strength 37-49 textural properties 37 see also Badshot dolomite; Rock Creek dolomite Domaine Sud-Armoricain (DSA) 216 anatectic rocks 209-217 core complex formation 220-221 development of centimetric peritectic cordierite 218 geological map 207 geology 206-209 lower crustal unit 219 migmatites 216, 220-221 P-T path 206 tectono-metamorphic evolution 209 tectono-stratigraphic units 206-208 Domaso-Cortaf6 TMU 239- 240 downdip mineral lineation 197 ductile deformation 135 ductile strain localization, stress history of 131 - 134 dykes growth 219 structural analysis 217-218 thicknesses 218 dynamic recrystallisation 8 olivine aggregates 143-158 earthquake behaviour 22 Eastern Alps as geological example 288-289 deep structure 291-292, 291

indentation and lateral escape 289 post-early Oligocene tectonics 288-292 tectonic map 289 easy glide concept 269 eclogite-facies metamorphism 177-178, 179 reactions 176 eclogitic deformation in large shear zones 181-183, 181 in little transformed zones 184-185 interpretation 188, 188 kinematics of 180-185 eclogitic shear zones/fractures 184 elastic constant 14-18, 17 elastic modulus 77 electron backscatter diffraction (EBSD) 1, 9, 100, 115 electron backscatter diffraction patterns 3, 100 'Elle' environment 12-13, 14 E1 Pilar-Coche Fault Zone 266 empirical failure criterion 42-45 empirical reaction order for dissolution and precipitation processes 82 European Variscides 99 exhumation fault-related 251 mechanisms 175-192 of continental crust 159-174 exhumation rate and strain rate 137 exponential probability density function 76 extension components 18-19, 18, 20 extension directions 18 extension fracture 16-18, 23 extensional midcrustal shear zones 127 fabric evolution degree of 229 stages 1 to 6 230 Fast Fourier Transform (FFT) analysis of crack-seal patterns 75 fault dimensions 25 fault discontinuities 25 fault growth 25 fault interaction 25 fault length and throw values 34 fault planes 27 fault-related exhumation 252, 254 fault scaling 33 fault segmentation 25, 31-33 fault throw distribution 25 flaw density 47, 48 flaw length 47 flaws and grain size 45-48 fluid pressure effect 22-23 Fourier analysis 74-76

313

314

INDEX

fracture development modelling 23 simulations 14-15 fracture modes 11-12, 15, 17-19, 21, 22-23 fracture orientation 60 fracture patterns as indicators of deformation history 11-24 modelling 12-14 pure shear deformation 15-16 fractures filled with vein material 12 Franciscan subduction complex 70 garnet porphyroclasts 200 garnetiferous mylonites 201 geological throw 26, 29-31, 3 0 - 3 2 gold quartz-bearing veins 68, 68, 69-70, 70, 70 grain and grain-boundary texture 48, 63-65 grain boundaries 14 alignments 144 diffusion and sliding (GBS) 144, 148-149, 153 preferred orientation (GBPO) 114, 114 quantitative microstructural analysis 100 springs 14 texture 48, 63-65 grain crushing 62-63, 63 grain size 14, 43, 48, 63, 65 and flaws 45-48 distribution 133, 136, 136 evolution across microscopic shear band 131-134 evolution within Tinos Island shear zone 134 histograms 57 quartz shear stress as function of 138 reduction 118 wall-rock 76 see also recrystallized grain size grain size insensitive (GSI) creep 147, 154 grain size insensitive (GSI) flow laws 154 grain size insensitive (GSI) power law creep 144, 146, 148, 148-149, 153-154, 156 grain size sensitive (GSS) creep 154 grain size sensitive (GSS) flow 118 grain size sensitive (GSS) linear creep 144, 147-148, 148, 149, 153, 153, 154-156 recrystallization-induced transition to 154-155 grain size sensitive (GSS) power law creep 146-147, 149, 152, 153, 156 flow law 153 granite, petrographic continuity 214 granite dykes 213-216 strike orientation 218 granite-filled structures 213 granulite-facies protolith 176-177

gravitational collapse concept of 193 fabrics resulting from 193 gravitational loading 19-23 gravitational potential anomalies (GPA) 298-299, 303 Graymont Dolime quarry 38, 53 Great Dyke 201 greenschist, retrogression 129 greenschist-facies 129, 179-180, 187 extensional shear zones 127 Gressoney extensional fault 266 Gressoney Shear Zone 1-2, 2, 9 Gubbio normal fault 29-30 Hall-Petch relationship 43, 45-46 high-pressure rocks 160 high-strain zones (HSZ) 9 coordinate reference frame 3 kinematic rotation axis 1-10 temporal evolution 9 Holsncy Island 175, 178 eclogites 175 deformation within Caledonides and Bergen Arcs kinematic framework 186-187 field relations 176-180 geological map 178 kinematics-geometrical model, geodynamic interpretation 186 P-T-t paths 190 post-eclogitic rotation 187-188 top-to-the-E sense of shear 190 horizontal extension 18 horizontal fractures 18 hornblende crystallographic preferred orientation (CPO) 115 crystals in ultramylonites 121 deformation mechanisms 98 grain boundaries 105 grains 104 microstructural study 97-98 microstructures 120 mineral chemistry 100 porphyroclasts 102 syndeformational chemical reactions 98 hornblende-plagioclase thermometer 109 Hundskjeften breccia zone 184 hybrid extension 20 hybrid extension fracture 12, 22 hybrid shear fracture 12, 13 hydrostatic compaction 21 Imasco Minerals quarry 53 inferred melt-bearing structures 215

INDEX initial statistical noise 16-18 interconnected weak layers (IWL) 97 interfacial reaction 82 mtracrystalline slip system 269 intrafolial isoclinal folds 197 mtragranular cleavage 65 mtragranular deformation 60, 62-65 mtragranular fractures 14 kinematic evolution 9 kinematic history 9 kinematic partitioning 2, 9 kinematic rotation axis 1-2, 9 in high-strain zones (HSZ) 1-10 microstructural characterization 8-9 kyanite-zoisite-rich matrix 183 Labview Software 52 Lake Como pre-Alpine basement, tectonometamorphic outline 238 Lake Como Southalpine deep crust deformation-metamorphism interactions 237- 239 geological setting 237 Languard-Campo Nappe deformation- metamorphism interactions 240-242 geological setting 240 polycyclic basement 240-243 Languard-Tonale TMU structural-petrographic map 241 tectono-metamorphic outline 243 Large Sample Rig (LSR) triaxial rock press 38 apparatus description 52, 52 calibration 52-53 chemistry and mineralogy 53 experimental results 56 experiments 58 geochemical data 54 macro and micro structures 59-61 mechanical results 58 results 58-61 sample assembly 52 starting material 53 stiffness 53 textural analysis grain size 53-55 hand sample and grain scale textures 55-58 porosity 55 lattice anisotropies 23 lattice directions 23 lattice effects 14 leptyno-amphibolite complex 98 Les Sables d'Olonne 220 leucosome, petrographic continuity 214, 216 leucosome-filled structures 213

315

Limpopo Belt 193, 194, 197, 202 late Archaean tectonics 203 Lindas Nappe, P-T-t evolution 180 lithosphere dry 152-154, 153-154 extrapolation from laboratory 152 hydrated 154 rheological regime maps 152-154, 153-154 lithosphere- asthenosphere transition 144 lithospheric mantle, deformation of 143 lithospheric-scale analogue modelling of collision zones 277-294 experimental concept 278 materials and initial setups 278-284, 279 mechanical properties of analogue materials 280 modelling results 284- 292 for different stages of deformation 281-283 modelling setup 278-284 scaling parameters of analogue experiment and prototype 284 scaling procedure 280-284 surface deformation and evolution of topography 284- 287 load-bearing framework (LBF) 97 localized ductile deformation 135 localized shear zones 100 Lower Eldsfjellet shear zone 180 Lugian domain 98 M. Cavallo-M. Fema segment 27 M. LeScalette segment 27, 29, 29, 30-31, 30-32, 34 M. Pennino segment 27 Marne a Fucoidi marker 29, 29 matrix bands 72 matrix grains 108 mechanical decoupling in rocks 251-252 melt-bearing structures, structural analysis 216-217 melt-flow network 213 melting and deformation, synergistic effects 205 -225 metagabbro microstructures 100-105 metagabbro mylonites, contrasting microstructures and deformation mechanisms 97-125 metagabbro protomylonite 102 metagabbro sheet deformation of eastern (lower) 102 deformation of western (upper) 104-105 metagabbro structures 101 metamorphic core complexes 127 metamorphic crust, tectonic analysis in 229 metamorphic reaction rate 242 metamorphic transformations 227-247

316 Metapelite samples 129 MgC12.6H20 83 mica fish structures 2 microcracks in diffusion controlled PSC 93-94 microfracture fillings 200 orientations 201 microscopic shear band 132 grain size evolution across 131 - 134 microstructural memory 9 microstructural preservation 9 midcrustal shear zone 134 stress-strain rate history 127-142 midcrustal strain localization to detachment formation, crustal evolution from 140 mid-oceanic ridges 76 Mighty White Dolomite quarry 38, 53 migmatite- granite relations 217 - 219 migmatites anisotropy 218 melt extraction 210 structural evolution 210-213 mineral chemistry and zoning 106-109 mineral solubility effect 92 mineral stretching lineations 9 misorientation analysis 3 misorientation angles 3, 5, 5 misorientation axis 3, 5, 7, 7, 8 geometry 3 misorientation data 5 MnNCKFMASH P-T pseudosection 208 Mohr circle 12, 15-16, 19, 20, 22 Mohr circle diagram 11 - 12, 12, 15, 22 Monte Mars region 237 Monte Mucrone-Monte Mars area 232 tectono-metamorphic outline 236 Monte Muggio TMU 239-240 morphological throw 26, 29-31, 30-32 Mtilikwe shear zone 196, 197-198, 201-202 gneissic fabrics 199 kinematic indicators 199- 201 mylonitic fabrics 198 optical and scanning electron microscope photomicrographs of fabrics 200 mylonite 100, 101,252 eastern 103-104, 103 petrology of 198-199 western banded 106 NIO fault 31, 32 nappe systems 244 Nelder-Mead routine 43 NH4C1 solution 93 Niagara 43, 55, 55, 57, 58, 59, 60-61, 64 Niagara dolomite 40 differential stress-axial strain curves 41

INDEX Niagara Formation 38, 53 Niagara samples 46 experimental results 38, 39 noise effects 16 North Limpopo Thrust Zone 202 Northern Marginal Zone (NMZ) 193-204 collapse features 202 contoured stereoplots of poles to foliations and lineations 196 crustal rheology 202 crustal thickening 201 geochronological constraints 201 geothermal gradient 203 magmatic accretion 201 mylonites 199 orogenic evolution 202 regional geological setting 195-197 study methods and samples 197 Norwegian Caledonides 175 Oligo-Miocene PFS 253 olivine aggregates creep curves 145 dry 146-148, 147, 149-152, 154 dynamic recrystallization 143-158 dynamic recrystallized grain size 149-152 high-strain experimental deformation 145 high-strain rheology 144-146 low-strain rheology 143-144 rheological regime maps 146-149, 147 rheology 143-149 strain softening 143-158 stress- strain curves 145 temperature dependence of recrystallized grain size 150-152 wet 148, 148, 149-152, 149, 154 orientation data 6 Ormos Isternia-Panormos Bay cross-section 129 orogens, infrastructure 205 Ortler Nappe 161 over-relaxation factor 13 Oygarden Gneisses 187 Pacific-North American plate boundary 265 peak differential stress 43 and mean stress 62 dolomite 44 peak strength 47 dolomite controls on 41-45 correlation coefficients 42, 42 variations in 37-49 Periadriatic fault system (PFS) 249-276 age and depth of exhumation across 256 age interval contours 255

INDEX age ranges 264 central segment 252, 255, 257 decoupling at indenter surface 264-266 decoupling within orogenic crust 265-267 eastern segment 252- 255 exhumation accommodated by Insubric mylonite belt ranges 256 former decoupling 252 generic model of decoupling and strain partitioning 270- 271,271 Insubric deformation 253 Insubric segments 252 intracrustal decoupling 252, 265-270 mylonitic segments 266 overview 252-253 segments of 250 significance for crustal strength and force transmission through lithosphere 269-270 southern steep belt 260 strike-slip activity 254 structural and age relations 263 structural map 253 symmetry and kinematics of mylonitic shearing 268 synthesis of structural, kinematic and age relations of Insubric faulting 261-264 tectonic map 251 tectonometamorphic age map of western end 259, 261 timing of strike-slip faulting and exhumation 254-259 Tonale and Pustertal segments 262 upper mantle structures 266 western end 261 western segment 259-261 age of faulting and exhumation 258-261 Insubric exhumation 260 tectono-metamorphic age map 258 peridotites 71 Petit Mont, Morbihan 209, 210-211, 214-216 phengitic substitutions 137-138 plagioclase BSE images 108 composition 107 crystallographic preferred orientation (CPO) 100, 115, 115 crystallographic reference frame 117 deformation 97, 116-120 fine-grained 118 grain, shape preferred orientation 113-114, 114 grain size distribution 112-113 inverse pole figures 117 matrix grains 120 microstructuraI study 97-98 microstructures 105 mineral chemistry 100

317

modal proportions 102 point and contour pole figures 116, 118 quantitative textural analyis 112 recrystallization 116-120 representative microprobe analyses 110 rheology 98 syndeformational chemical reactions 98 planar and linear fabric elements 258 plastic behaviour 20 plate-scale fault systems 269 point and contour pole figures 119 Poisson effects 19 porosity, influence of 64-65 Port Navalo Plage 211, 216 powder compaction 83 power-law correlations 67 Presqu'~le de Rhuys, Morbihan 211, 217 pressure solution compaction experiments 83 pressure solution creep (PSC) diffusion-controlled 81-82, 91-93 effect of fluid chemistry 83-84 effect of microcracks 84, 93-94 high-resolution experiments 85, 88 experimental results 8 9 - 9 0 profiles of indented regions 89-90 rates of deformation at different time intervals 89 with continuous deformation recording 86, 86, 90 effect of stress 92-93 with continuous displacement recording 87-89 increase in recording sensitivity 84 indenter experiments 85 experimental conditions and results 86 with ex situ deformation measurement 84-87 comparison with halite 92-93 indenter experiments with ex situ deformation measurement dominant deformation mechanism 90 effect of stress 92 kinetic-controlled 82, 91-92 rates on calcite monocrystals 81-95 reaction-kinetics-controlled 81 single contact experiments 82 pressure solution models 82 principal stresses 15, 19-20, 19, 38, 41 protomylonite 100, 101, 102, 198 grain size distribution 113 pseudotachylite 252, 266 P-T conditions 176 and timing for shear zone formation 129 during vein formation 76 P-T eclogitic field 176 P-T loop 228 P-T path 228 and temperature variation 137

318 P-T-d-t loop 244 P-T-d-t path 228, 239 P-T-t evolution 180 P-T-t path 160-163, 163 pyroxene, representative microprobe analyses 111 Quantitative Textures Analyses (QTA) 242 quartz dislocation creep flow law 139 grain size distribution 136, 136 grain size piezometer 127 recrystallized grain size 135, 139 brittle deformation inferred from 127-142 evolution 128 shear stress 135 and strain rate within midcrustal shear zone 127-142 as function of grain size 138 quartz veins, Abitibi, Canada 68, 76-77 quartzo-feldspathic rocks 118 quasi steady-state behaviours 22 recrystallization-induced transition to grain size sensitive (GSS)-linear creep 154-155 recrystallized grain size 132, 132 decrease with increasing strain 131-133 evolution 134 memory 133-134 piezometer 128 scaling law 155 relaxation algorithm 13 relaxation threshold 23 theological behaviour 17-18 Rock Creek dolomite 38, 40, 43, 46-47, 53, 55, 55, 57, 58-59, 59, 60, 61, 62 differential stress-axial strain curves 41 experimental results 38, 39 Rossland amphibolitic shear zone 185 Rupati Pools 197, 200-201 San Andreas system, California 70-71, 71 scanning electron microscope (SEM) 2 Schmid Factor analysis 5 - 7 segment interaction 31 - 33 seismic rupture 25 nucleation 32 sense of shear deduced from asymmetrical shapes of granulite boudins 182 deduced from S-C structures 182 serpentine veins 70-71, 71

INDEX Sesia-Lanzo Zone (SLZ) correlation between degree of fabric evolution and progress of metamorphic transformation 237 polycyclic basement 230-237 structural-petrographic map 232 tectonic sketch 232 Sestri Levante 13 shear bands 2 shear deformation 9, 14-15, 15, 18-20, 19, 23 shear displacement 23 shear failure 18, 48 shear fracture 11 - 12, 16-18, 20, 22- 23 shear stresses 132, 133-137, 269 and buoyancy forces 189 and temperature 133 shear zone 134 shear zone formation, P - T conditions and timing for 129 Sigma deposit 70 Silesian domain 99 slip systems 6, 7-8, 16 role of 5 - 8 South Tibetan Detachment Fault 202 spatial correlation, absence of 74-76 spring network model 76 stacking fault energy (SFE) 152 Star~ Mesto (SM) belt 98, 99 Stipp and Tullis piezometer 131 strain localization mechanism 128 strain partitioning in continental lithosphere 249-276 strain rate 132, 133-137, 135 and exhumation rate 137 and grain size relationship 129 and temperature 132 quantification 128 strain softening 46, 118 olivine aggregates 143-158 strength-depth profiles 250 stress field in Western/Central Alps 295- 310 stress history of ductile strain localization 131 - 134 stress release fluctuations in crustal rocks 67-79, 68 variations recorded by crack-seals 69 stress shielding mechanism 17 stress-strain curves 19-20, 21 stress-strain rate history of midcrustal shear zone 127-142 stress- strain relationship 17, 18, 19 during deformation 15, 15 stromatic leucosome, anisotropic nature 213 stromatic migmatite 209 petrographic features 211

INDEX subduction continental crust 159-174 continental lithosphere 164-166 geometry 186 syndeformational chemical reactions 106 syneclogite deformation in Bergen Arcs, Norway 175-192 Tasmanian-Pacific plate boundary 266 Tauern window 252, 253-254, 264, 268 Taylor models 8 tectonic analysis in metamorphic crust 229 tectonic deformation 14, 20, 21, 23 tectonic extension 21 tectonic transport direction 1 tectono-metamorphic evolutions of Austroalpine and Southalpine lithostratigraphic units, complexes and TMUs 233-235 tectono-metamorphic unit (TMU) 227-247 boundaries 228 identification 228 size definition 227 synoptic outline 233-235 tectono-thermal memory of rocks 227-247 temperature and strain rate 132 derived from quartz dislocation creep flow law 128 evolution 134-137, 135 variation and P - T path 137 tensile fractures 18, 20 tensile stresses 12, 15-16, 17, 18, 20, 22 textural controls on brittle deformation of dolomite 51 - 66 throw distribution 30 Tinos Island cross section and sample location 130 extensional shear zone 127-131 validation of proposed method for quantification of stress, strain rate and temperature 137 grain size evolution within shear zone 134 metamorphic core complex 127 metapelites 138 microstructures 131 phengitic substitution 130 shear zone, quartz piezometer and flow law 129-131 structural map 130 Tonale fault segment 265 Tonale Nappe 161 'top-to-the-SE rotation' model 187-188 tourmaline 68, 70, 70 transgranular deformation 62-65 Triangle shear zone 202 triangular lattice 13, 13, 23 triaxial rock deformation experiments 51-52

319

triaxial stress state 63 turbidites 71 Twiss piezometer 131 Ulten Unit, Tonale Nappe, Eastern Austroalpine 159-174 computation of temperature distribution and negative buoyancy 165 downdip length of continental crust 169 fast exhumation: buoyancy-driven tectonic extrusion 166-170 geodynamic model 164-171 geodynamic scenario 163-164, 164 geological framework 160-164 P - T - t - p a t h 160-163, 163 petrology 160 rocks of Ulten m~lange 162 slow exhumation: slab break-off and lithosphere extension 170-171 thermoelastic parameters 166 time evolution of negative buoyancy 167-168 upward buoyancy of continental material 169 ultra-high-pressure (UHP) metamorphic rocks 160 ultramylonite 100, 101 eastern 104, 104 hornblende crystals in 121 Umbria Fault System (UFS) 26 Umbria-Marche region 25, 26 stratigraphy 29 Upper Carboniferous brittle-ductile transition zone 220 Val d'Or, Abitibi, Canada 69-70 Variscan Belt, Western France age relations 208-209 evolution of European Variscides 206 synergistic effects of melting and deformation 205 -225 Variscan tectonics 98 vein patterns 22 as indicators of deformation history 11-24 vein sets 13, 22-23 Vilaine Group 207 wall-particles 13 Western/Central Alps 2 correlations with crustal thickness 296-298 finite element modelling 298-299, 300 geodynamic implications 307 isostatic model 299 model configuration 301-302, 304-306 models with fixed boundaries 299-303 models with rotational boundaries 303- 307 realistic crustal model 303

320 Western/Central Alps (Continued) seismotectonic data 296-298 seismotectonic overview 297 seismotectonic strain and stress fields 296 stress field in 295-310 thermal parameters, densities and rheological parameters of models 298 Western Gneiss Region 186 white light interferometer (WLI) microscopy 89, 91

INDEX wing crack coalescence 48 model 45 propagation 46 yield stress 38 Young's modulus 17-18, 38, 41-42, 45, 48, 58 Zimbabwe Craton 194-195, 197

Deformation Mechanisms, Rheology and Tectonics from Minerals to the Lithosphere Edited by D. Gapais, J.-P. Brun and P. R. Cobbold

This book consists of 18 papers on deformation mechanisms, theology and tectonics. The main approaches include experimental rock deformation, microstructural analysis, field structural studies, analogue and numerical modelling. New results on various topics are presented, ranging from brittle to ductile deformation and grain-scale to lithosphere-scale mechanisms. The volume will be of interest to academic and industrial researchers in the fields of structural geology, interactions between metamorphism, fluids and deformation, and large-scale tectonic processes.

Visit our online bookshop: http:llbookshop.geolsoc.org.uk Geological Society w e b site: http:/Ivvww.geolsoc.org.uk

Cover illustration: C-S structures in syn-kinematic granite from the South Armorican ShearZone (Brittany, France)(photograph D. Gapais), and thrust structures within a brittle-ductile analogue model (photograph by L. Barrier)