Vertical Coupling And Decoupling in the Lithosphere (Geological Society Special Publication No. 227)

Vertical Coupling and Decoupling in 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: GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) 2004. Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227. GIORGIS, S., MARKLEY, M. & TIKOFF, B. 2004. Vertical-axis rotation of rigid crustal blocks driven by mantle flow. In: GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227, 83-99.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 227

Vertical Coupling and Decoupling in the Lithosphere EDITED BY

J. GROCOTT Kingston University, UK

K. J. W. MCCAFFREY University of Durham, UK

G. TAYLOR University of Plymouth, UK and

B. TlKOFF University of Wisconsin, USA

2004 Published by The Geological Society London

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Contents GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. Vertical coupling and decoupling in the lithosphere

1

Geophysical constraints on vertical coupling in the lithosphere SAVAGE, M. K., FISCHER, K. M. & HALL, C. E. Strain modelling, seismic anisotropy and coupling at strike-slip boundaries: applications in New Zealand and the San Andreas Fault TIKOFF, B., Russo, R., TEYSSIER, C. & TOMMASI, A. Mantle-driven deformation of orogenic zones and clutch tectonics

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41

Vertical axis block rotations in the upper crust, horizontal and vertical partitioning and implications for vertical coupling-decoupling in the lithosphere LEGG, M. R., KAMERLING, M. J. & FRANCIS, R. D. Termination of strike-slip faults at convergence zones within continental transform boundaries: examples from the California Continental Borderland

65

GIORGIS, S., MARKLEY, M. & TIKOFF, B. Vertical-axis rotation of rigid crustal blocks driven by mantle flow

83

TEYSSIER, C. & CRUZ, L. Strain gradients in transpressional to transtensional attachment zones

101

CERCA, M., FERRARI, L., BONINI, M., CORTI, G. & MANETTI, P. The role of crustal heterogeneity in controlling vertical coupling during Laramide shortening and the development of the Caribbean-North America transform boundary in southern Mexico: insights from analogue models

117

FABBRI, O., IWAMURA, K., MATSUNAGA, S., COROMINA, G. & KANAORI, Y. Distributed strike-slip faulting, block rotation, and possible intracrustal vertical decoupling in the convergent zone of SW Japan

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Lower crustal flow and topography WHITNEY, D. L., PATERSON, S. R., SCHMIDT, K. L., GLAZNER, A. F. & KOPF, C. F. Growth and demise of continental arcs and orogenic plateaux in the North American Cordillera: from Baja to British Columbia

167

Orogenic examples MCCLELLAND, W. C. & OLDOW, J. S. Displacement transfer between thick- and thin-skinned decollement systems in the central North American Cordillera

177

KLEPEIS, K. A. & CLARKE, G. L. The evolution of an exposed mid-lower crustal attachment zone in Fiordland, New Zealand

197

vi

CONTENTS

MCCAFFREY, K. J. W., GROCOTT, J., GARDE, A. A. & HAMILTON, M. A. Attachment formation during partitioning of oblique convergence in the Ketilidian orogen, South Greenland

231

FERNANDEZ-FERNANDEZ, E., JABALOY, A. & GONZALEZ-LODEIRO, F. Lower Miocene deformation in the hanging wall of the Internal-External Zone boundary of the Betic Cordillera: deformation at the edges of vertical-axis rotation domains in oblique convergent margins

249

Subduction examples - island arcs and marginal basins TAKESHITA, T. & YAGI, K. Flow patterns during exhumation of the Sambagawa metamorphic rocks, SW Japan, caused by brittle-ductile, arc-parallel extension

279

FABBRI, O., MONIE, P. & FOURNIER, M. Transtensional deformation at the junction between the Okinawa trough back-arc basin and the SW Japan island arc

297

Melts and crustal rheology WHITNEY, D. L., TEYSSIER, C. & FAYON, A. K. Isothermal decompression, partial melting and exhumation of deep continental crust

313

VIGNERESSE, J. L. & BURG, J. P. Strain-rate-dependent rheology of partially molten rocks

327

Index

337

Acknowledgements This special publication developed out of two symposia: 'Deformation at Convergent Margins', convened at the European Union of Geosciences meeting (BUG XI) at Strasbourg in April 2001; and 'Vertical Coupling and Decoupling at Convergent Margins' convened at the AGU Fall Meeting in San Francisco in December 2001. We are grateful to all the authors for submitting their work and supporting the volume and we are indebted to our expert panel of reviewers who, through their conscientious work, have ensured the high quality of the volume. We are especially grateful to Dr Stella Bignold (Kingston University) who assisted us in the compilation of this volume. The panel of reviewers for this volume was: M. Anderson, University of Plymouth, UK J.R. Andrews, University of Southampton, UK T.S. Brewer, University of Leicester, UK M. Brown, University of Maryland, USA J.-P. Brim, Universite de Rennes 1, France P. England, University of Oxford, UK J. van Gool, GEUS, Denmark K. de Jong, National Taiwan University, Taiwan H.G.A. Lallemant, Rice University, Houston, USA S. Lamb, University of Oxford, UK L. Lonergan, Imperial College, London, UK B. Murphy, St Francis Xavier University, Nova Scotia, Canada L. Parson, Southampton Oceanography Centre, UK C.W. Passchier, Johannes Gutenberg University, Mainz, Germany B. Pelletier, IRD, New Caledonia C. Rosenberg, Freie Universitat, Berlin, Germany E. Rutter, University of Manchester, UK

M. de St-Blanquat, Universite Paul Sabatier, Toulouse, France M.K. Savage, Victoria University of Wellington, New Zealand G. Schreurs, University of Bern, Switzerland F. Spera, University of California, USA R. Strachan, Oxford Brookes University, UK A. Thomassi, University de Montpellier II, France M. Tolson, Ciudad Universitaria, Mexico J. Turner, University of Birmingham, UK O. Vanderhaeghe, Universite Henri Poincare Nancy 1, France A. Vauchez, University de Montpellier II, France J.-L. Vigneresse, CREGU, Vandoeuvre-les-Nancy, France R.L.M. Vissers, Utrecht University, Netherlands C. Waters, University of Wisconsin, USA G.K. Westbrook, University of Birmingham, UK R. Wintsch, Indiana University, USA

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Vertical coupling and decoupling in the lithosphere JOHN GROCOTT1, KENNETH J. W. McCAFFREY2, GRAEME K. TAYLOR3 & BASIL TIKOFF4 1 Centre for Earth and Environmental Science Research, School of Earth Sciences and Geography, Kingston University, Kingston-upon-Thames, Surrey, KT1 2EE, UK (e-mail: [email protected]) 2 Department of Earth Sciences, University of Durham, Durham, DH1 3LE, UK ^School of Earth, Ocean and Environmental Sciences, University of Plymouth, Devon, PL4 8AA, UK ^Department of Geology and Geophysics, University of Wisconsin, Madison, WI53706, USA Continental tectonics, and the formation of mountain belts, do not adhere to the plate tectonic paradigm (Molnar 1988). Mountain belts at plate boundaries are areas of diffuse deformation in which geologists have recognized that not only are the plates not rigid (Gordon 1998), but parts of the lithosphere (e.g. upper crust) are moving laterally with respect to other parts (e.g. lower crust), such as in thrust belts (Bally et al. 1966). An exciting development in tectonics is the detailed investigation of the behaviour of continental crust during orogenesis. In particular, the role of coupling (attachment) or decoupling (detachment) of the lithospheric layers during continental deformation has significant implications for all aspects of modern and ancient tectonics. Historical development The recognition of regional detachments or decollements, an idea developed in foreland fold and thrust belts, was the first major contribution to our understanding of the vertical stratification in orogenic belts. Elucidation of the structure of foreland fold and thrust belts in the external parts of orogens by the development of the techniques of balanced cross-section construction (Price 1981, 1986) and deep seismic reflection profiling (e.g. Mueller et al. 1980) showed that foreland thrust systems are typically thin-skinned and bounded at depth by a basal detachment, decollement or sole thrust. Shortening in the external parts of cordilleran and collisional orogens was taken up by folding and thrusting above a basement which remained essentially undeformed and part of the foreland (e.g. Bally et al. 1966). Balanced cross-section techniques were developed and refined in the

Alberta foothills district of western Canada (Dahlstrom 1969) and later applied worldwide to show that shortening exceeded 100s of km for many foreland fold-thrust belts (e.g. Allmendinger et al. 1986, 1987 for the Cretaceous Sevier orogenic system, North America). The magnitudes of these displacements carry profound implications for the deep structure of the internal zones of orogens that were explored in landmark papers by Oldow et al (1989, 1990). These authors pointed out that simple massbalance considerations allowed the position of the basal decollement to be extrapolated below the internal zones of orogens, with the implication that entire orogens were underlain by decollement systems and were thus 'floating' on the underlying lithosphere. This concept of 'orogenic float' (Oldow et al. 1990) is also applicable to ocean-continent subduction boundaries as there is often a clear link between orogenic displacements and the relative convergence vectors of the lithospheric plates. For example, where the relative velocity vector of the subducting plate is oblique to the trench, plate deformation of the upper plate is partitioned, typically, into margin-parallel strike-slip displacements and margin-normal dip-slip displacements (Dewey 1982), at least where the bulk strain in the upper plate is an oblique convergence (transpression) (McCaffrey 1996; Grocott & Taylor 2002). Coeval strike-slip and dip-slip displacements can be reconciled easily with the orogenic float model, since it is not a requirement that the strike-slip fault systems traverse the entire lithosphere or even the crust. Strike-slip fault systems, like the dip-slip fault systems, simply detach on the decollement at the base of the orogen. What Oldow et al. (1990) found difficult to account for was how

From: GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) 2004. Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227, 1-7. 0305-8719/04/$15 © The Geological Society of London 2004.

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forces were transmitted laterally through the orogen from the plate boundary, particularly through the thermally weakened magmatic arc, to the back arc domains where considerable horizontal shortening was accommodated in thinskinned thrust belts. An alternative explanation is that the lithospheric layers are attached, not detached, along sub-horizontal shear zones, which was first explored in a transpressional setting (Teyssier & Tikoff 1998). For example, total strike-slip partitioned transpression in the brittle-ductile upper crust may overlie a domain of homogeneous or heterogeneous transpression in the ductile lower crust. A horizontal fault or shear zone is required between the two layers, to accommodate the different combinations of translation, rotation and strain taking place at different crustal levels. Assuming that the shear zone has a finite thickness, it can be viewed as maintaining kinematic compatibility (i.e. it is an attachment zone) between two vertically coupled domains in which the deformation is partitioned differently. These attachment or coupling zones link rheological domains vertically in the lithosphere. They are fundamentally different from the basal detachment or decollement zones, because they provide kinematic compatibility and attachment (coupling) between domains that have accommodated strains in different ways (Molnar 1992). A similar approach to crustal attachments was taken by Vauchez & Nicolas (1991) and Vauchez etal (1998). Transpressional zones were, in some respects, the obvious tectonic setting to develop the concept of attachment zones, because of the peculiar boundary conditions. Side-loading of transpressional systems, with two blocks moving toward each other and causing an intermediate region to deform, was the original conceptual model for this type of deformation formulated by Sanderson & Marchini (1984). This was a seminal paper in the analysis of deformation patterns, although multiple authors commented on the unrealistic nature of these boundary conditions (e.g. Jones et al. 1997). Transpression is equally well described by bottom boundary conditions: the 'bottom-up' transpression of Teyssier & Tikoff (1998). In this model, transpression is driven by a flow field from below, in the lower crust or mantle lithosphere. This approach does not require any vertical strain, and the compatibility problems of side-driven models of transpression are avoided. The model was applied particularly to central California, in which dextral displacement during oblique convergence between the North American and Pacific plates is accommodated

by a pervasive, steeply-dipping shear fabric in the mantle lithosphere. This mantle deformation is linked to upper crustal strike-slip partitioned transpression by heterogeneous deformation in the mid- to lower crust, where a flat coupling (or attachment) zone allows the translation of upper crustal blocks over the shearing mantle. The model was expanded by Tikoff et al. (2002) and Teyssier et al (2002). The concept of attachments was also applied to obliquely convergent subduction boundaries by Saint Blanquat et al. (1998). These authors envisaged that the subducting plate and the upper plate are partially coupled and that this drives deformation in the overriding plate from below via a series of partial attachment zones (Tikoff et al. 2002). In this way, oblique convergence can drive margin-parallel translation of upper crustal, lower crustal and upper mantle fore-arc slivers in the over-riding plate; each sliver separated by a low-angle zone of partial attachment controlled by the major lithological characteristics of the plate (quartz-, feldsparand olivine-dominated lithologies). This idea was further developed in the Ketilidian orogen of southern Greenland (Garde et al. 2002). While the concept of horizontal attachment zones was being explored in the geological literature, a similar development was occurring in the geophysical literature. For example, Bourne et al. (1998) used geodetic data to infer the presence of a sub-horizontal attachment zone below strikeslip faults in New Zealand and California, again from transpressional settings. Savage et al. (1999) responded directly to this article that the same signal could be created with a detachment zone in the same location. While the debate remains unresolved, it merely indicates that this is an issue of broader consequence, despite the all-too-familiar non-communication between the different sub-disciplines.

This volume The question of attachment and/or detachment horizons in the lithosphere is an exciting field that is actively being researched. This topic has been and is the focus of several dedicated sessions at international meetings and is, in part, the subject of the recent European Geosciences Union volume focussing on the foreland region (Bertotti et al. 2002). In this volume, we have attempted to bring together a variety of different viewpoints on this subject, to discuss the spatial and temporal patterns of crustal attachment and detachment zones. Thus, this introduction does not mean to constitute a thorough review of this field, but rather we will let the individual

VERTICAL COUPLING AND DECOUPLING IN THE LITHOSPHERE

papers speak for themselves. Below, we outline the major topics covered in this book, in order, and try to summarize briefly the context of the contributions.

Geophysical constraints on vertical coupling in the lithosphere Shear wave splitting Lateral and vertical deformation partitioning and clear links between relative plate convergence vectors and upper plate deformation at subduction boundaries, provide strong geological grounds for vertical coupling or partial coupling of deformation in the lithosphere. Of course, samples of mantle lithosphere are rarely obtained at Earth's surface, and direct evidence for the compatibility of deformational fabrics in crust and mantle lithosphere is not available. Fabrics in the mantle lithosphere can be studied geophysically by determining the birefringence produced in vertically incident shear waves by olivine lattice preferred orientations (LPOs) in the mantle lithosphere (Silver 1996; Savage 1999). This technique has been used to image mantle anisotropy beneath the San Andreas fault system (Teyssier & Tikoff 1998), the Alpine fault system of New Zealand (Klosko et al. 1999) and strike-slip fault systems in the Caribbean (Russo et al. 1996). Shear-wave splitting patterns in California and New Zealand differ in detail despite similar transform fault settings. From the assumption that fabrics in the asthenosphere are ultimately caused by translation, rotation and strain of lithospheric plates (i.e. a top-driven system), Savage et al. use models of mantle flow and strain for a number of strike-slip plate boundary types and compare them to observed shear-wave splitting patterns in New Zealand and California. The most successful models for the northern San Andreas fault system require rapid changes of viscosity with depth, implying a strong component of vertical decoupling. For southern California and for New Zealand, a stratified viscosity model accounts for many aspects of the splitting observed and clearly implies partial vertical coupling of the lithosphere in these regions. Whether systems are top- or bottom-driven is addressed by Tikoff et al. who point out that if lithospheric layers are separated by horizontal decollement zones or detachments as envisaged by Oldow et al. (1990), then this is an obstacle to understanding how lithospheric deformation relates to deeper Earth processes including mantle convection. However, patterns of upper-crustal deformation are similar to mantle deformation

3

patterns, as determined by shear-wave splitting in a variety of settings, implying that the layers are at least partially coupled. Tikoff et al. recognize and emphasize the difficulties in applying the plate tectonics concept to mountain belts. While it is straightforward to envisage rigid oceanic lithosphere moving over the ductile asthenosphere driven (primarily) by slab-pull, rigid lithospheric columns are less likely to exist in orogenic settings because of the inherent weaknesses of (over) thickened continental crust. As Oldow et al (1990) appreciated; in mountain belts it is much more difficult to envisage the source of body forces (other than those supplied by topography), that give rise to deformation. Tikoff et al. present the case that flow of the mantle drives deformation in continental orogenic zones by loading from below. A consequence of the bottom-driven idea is that mantle flow drives displacement of continental plates in orogenic settings while, in the oceans, slab-pull (topdown) systems dominate. Nevertheless, top-down systems can still be assumed for continental lithosphere and are quite capable of explaining the shear-wave splitting data (Savage et al).

Vertical axis rotations Important implications of vertical coupling between lithospheric layers concern the accommodation at depth of vertical axis block rotations recorded palaeomagnetically in the upper continental crust at wrench and subduction boundaries. Giorgis et al. describe an approach in which vertical axis rotations of upper crustal blocks are shown to be consistent with the orientation of olivine LPOs in a 'bottom-up' vertically-coupled system. The magnitude of vertical axis rotations determined palaeomagnetically is frequently much larger than can be reconciled with observed fault slip displacements. For example, in the Coastal Cordillera of northern Chile, clockwise vertical axis rotations typically approach 45°, although it is difficult to identify displacements on faults at an appropriate scale and magnitude to account for these as simply localised in situ block rotations (Randall et al 1996; Taylor et al 1998). This dilemma may be resolved in one of two ways: either a pattern of faults exists in space and time and at a range of scales, so that each fault set accounts for only an increment of the vertical axis rotation (Grocott & Taylor 2002); or rotation is achieved by distributed strain during (oroclinal) bending as described in this volume for the Japan arc by Fabbri, Iwamura, Matsunaga, Coromina & Kanaori.

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Whatever the principal cause of vertical axis rotation, the large scale of the blocks commonly found to have rotated at wrench or subduction boundaries begs the question of how the rotation is accommodated at depth. The blocks may, or may not be, bounded by faults that root deeply into the lithosphere. Where the blocks are bounded by regionally significant faults, such as in the Chilean Coastal Cordillera (Taylor et al 1998), there is little support from reflection geophysics that they vertically transect the whole of the crust or the lithosphere (Scheuber & Giese 1999; Adam & Reuther 2000). Instead, the magnitude of vertical axis rotations may imply that a mid-crustal layer exists upon which both the faults and the blocks rotate and are vertically coupled to, and driven by, deformation in the deeper lithosphere and mantle. Evidence that block rotation is indeed driven by basal shear across a low-angle attachment zone is presented by Legg et al. in their analysis of vertical axis rotations in the Californian continental borderland. The patterns of strain to be expected in attachment zones below an upper crust characterized by translating and/or rotating fault blocks were investigated by Teyssier et al (2002) for deformation characterized by a large vertical axis rotation. In this volume, Teyssier & Cruz determine strain gradients in transpressional to transtensional attachment zones below rigid blocks in the upper crust that were translated but not rotated. This approach is relevant because it seems that most of the vertical axis rotation component of deformation in the Japan arc (Fabbri, Iwamura, Matsunaga, Coromina & Kanaori) is related to the formation of a large-scale oroclinal bend, rather than to in situ block rotations accommodated by slip on steeply dipping faults. In such circumstances, the patterns of foliation and lineations predicted by Teyssier & Cruz should be observed as commonplace, when and if attachment zones are exposed.

Examples of vertical coupling from orogenic belts and exposed attachment zones McCaffrey et al. describe a sub-horizontal attachment zone from the forearc of the Palaeoproterozoic Ketilidian orogen in southern Greenland. Flat-lying, intense ductile fabrics with orogen-parallel stretching lineations and orogen-parallel transport of the upper plate are interpreted to underlie upper-crustal blocks that were translated parallel to the subduction boundary. Deformation was driven from below by coupling of the down-going plate with the

upper plate. In the Ketilidian, both upper crustal and lower crustal structural levels are exposed. Each had its own partitioning mode separated by an attachment zone that migrated to higher structural levels, overprinting the upper crustal domain. Vertical migration of the attachment zone was due to an increase in geothermal gradient caused by migration of the locus of arc magmatism into the earlier forearc. The attachment zone exposed in the Ketilidian fore-arc dips below the Julianehab arc batholith to re-emerge as the basal 'detachment' of a foreland-fold-thrust belt on the external side of the orogen. The trajectory of the attachment zone beneath the core of the orogen has been traced on seismic reflection profiles (Dahl-Jensen et al 1998) and in the field (McCaffrey et al.). Klepeis & Clarke contribute an extremely well-documented example of a mid-lower crustal attachment zone from Fiordland, New Zealand, showing how displacements were relayed through a 25 km-thick section of exhumed Lower Cretaceous middle and lower crust. Steep shear zones in the lower crust are shown to merge with a sub-horizontal attachment zone that coupled lower crustal deformation to deformation in a middle crustal fold-thrust zone. In both the Greenland and New Zealand studies, there are clear echoes of the orogenic float conceptualization of Oldow et al (1990). In an example from the North American Cordillera, McClelland & Oldow build directly on the ideas of Oldow et al (1990) and, although they promote the concept of a deep 'decollement' below the cordilleran orogen, they do not envisage this as a deep structural discontinuity. Rather, "a substantial degree of mechanical coupling is predicted from geodynamic models and is required to accommodate displacement and volumebalance considerations for lithospheric-scale deformation." An example from the Betic cordilleras of Spain (Fernandez-Fernandez etal), and analogue experiments based on the tectonic setting of southern Mexico during the Laramide orogeny (Cerca et al.) give further weight to the hypothesis that patterns of displacement, strain and rotation in the upper crust are underlain by a partial attachment horizon above a zone characterized by more distributed deformation.

Examples of vertical coupling from subduction boundaries: island arcs and marginal basins Sub-horizontal zones of deformation in the middle crust of the upper plate at active subduction boundaries have not been modified by collision

VERTICAL COUPLING AND DECOUPLING IN THE LITHOSPHERE

or terrain accretion. An example from the contemporary southern Japan arc (Fabbri, Monie & Fournier) shows just how important these structures are in the deformation and architecture of the lithosphere. These authors describe a system of regional thrust faults reactivated as low-angle normal faults that merge into a mid-crustal partial attachment zone. This zone provides the kinematic linkage between the lower crust characterized by ductile distributed deformation and the upper crust characterised by heterogeneous, distributed brittle deformation in a 600 km-long, across-strike section through the Japan arc. In a companion paper, Takeshita & Yagi describe deformation and metamorphism in a low-angle ductile shear zone of Late Cretaceous age in southern Japan, that may have formed the upper to lower crustal attachment zone of the ancestral Japan arc.

Channel flow: vertical decoupling in the lithosphere In the Himalayan-Tibetan orogenic system, topography is buffered by a sub-horizontal zone of partial melting located below Tibet which facilitates horizontal collapse of over-thickened crust (Hodges 2000; Jackson 2002). In effect, the lower crust is Theologically weakened by partial melting and flows away from where it underlies the highest topography in a thick, sub-horizontal channel (Bird 1991). To the south, in the Himalayan arc, the channel is exposed, the upper boundary is marked by north-dipping, low-angle, extensional faults of the South Tibetan fault system and the lower boundary by the Main Central thrust system. In effect, the rocks entrained in the channel are extruded southwards and simultaneously removed from the orogen by erosion. Extensive partially molten layers in the lower crust of orogenic belts carry the implication that deformation in the lithosphere may be comprehensively decoupled from mantle flow at least temporarily. The Andean and the Himalayan-Tibetan orogenic systems are characterized by high plateaux: the Altiplano of Peru and Bolivia, and the high plateau of Tibet. In their paper, Whitney, Paterson, Glazner & Kopf provide evidence that the North American Cordilleras were also characterized by high plateaux, reaching altitudes of 4-6 km, and thickened crust in Late Mesozoic-early Cenozoic time. In detail, there are strong differences in the exhumation style and magnitude from north to south in the cordilleran arcs, and this may reflect differences in the degree of coupling between the subducted

5

plate and the overriding plate. Differences in the degree of coupling may be internal depending, for example, on crustal rheology (Himalaya) or external, depending on a change in plate motion vectors such as an increase in the roll-back velocity. In a related contribution, Whitney, Teyssier & Fayon address the causes and consequences of partial melting of the continental lithosphere in thickened continental crust. They argue that partially molten layers 20-30 kmthick may be formed by decompression partial melting and that this will facilitate decoupling and horizontal channel flow. On the other hand, there is also potential for buoyancy-driven vertical flow of the partially molten layer and this may give rise to the ubiquitous mantled-gneiss domes in the internal zones of orogenic belts. Whitney, Teyssier & Fayon note that decompression typically accompanied emplacement of mantled gneiss domes (metamorphic core complexes) and may have been facilitated by isostatic uplift during extensional deformation and/or by diapiric ascent. Some aspects of the strain rate-dependent rheology of partially molten rocks are addressed by Vigneresse & Burg. Change in rheology attendant on partial melting may well facilitate a transition from an upper plate lithosphere coupled to mantle deformation at subduction boundaries, to a decoupled lithosphere subjected to severe crustal thickening during terrane accretion or continent-continent collision.

References ADAM, J. & REUTHER, C.-D. 2000. Crustal dynamics and active fault mechanics during subduction erosion. Application of frictional wedge analysis on to the North Chilean forearc. Tectonophysics, 321, 297-325. ALLMENDINGER, R.W., FARMER, H., MAUSER, B.C., SHARP, J., VON TISH, D., OLIVER, J., & KAUFMAN, S. 1986. Phanerozoic tectonics of the Basin and Range-Colorado Plateau transition from COCORP data and geologic data. In: BARAZANGI, M. & BROWN, L. (eds) Reflection Seismology: The Continental Crust. American Geophysical Union Geodynamics Series, 14, 257-268. ALLMENDINGER, R.W., HAUGE, T.A., HANSE, E.G., POTTER, J.C., KEMPER, S.L., NELSON, K.D., KNUEPER, P. & OLIVER, J. 1987. Overview of the COCORP 40° N transect, western United States—The fabric of an orogenic belt: Geological Society of America Bulletin, 98, 308-319. BALLY, A., GORDY, P.L. & STEWART, G.A. 1966. Structure, seismic data, and orogenic evolution of southern Canadian Rockies: Canadian Association of Petroleum Geologists Bulletin, 14, 337-381.

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J. GROCOTT £7AL.

BERTOTTI, G., SCHULMANN, K. & CLOETINGH, S. (eds) 2002. Continental Collision and the TectonSedimentary Evolution of Forelands. Stephan Mueller Special Publication Series, 1, European Geosciences Union. BIRD, P. 1991. Lateral extrusion of lower crust from under high topography, in the isostatic limit. Journal of Geophysical Research, 96, 10275-10286. BOURNE, S.J., ENGLAND, P.C. & PARSONS, B. 1998. The motion of crustal blocks driven by flow of the lower lithosphere and implications for slip rates of continental strike-slip faults. Nature, 391, 655-659. DAHLSTROM, C.D.A. 1969. Balanced cross sections. Canadian Journal of Earth Sciences, 6, 743—757. DAHL-JENSEN, T., THYBO, H., HOPPER, J. & ROSING, M. 1998. Crustal structure at the SE Greenland margin from wide-angle and normal incidence seismic data. Tectonophysics, 288, 191-198. DEWEY, J.F. 1982. Episodicity, sequence and style at convergent plate boundaries. In: STRANGEWAY, D.W. (ed.) The Continental Crust and its Mineral Deposits. Geological Association of Canada, Special Publications, 20, 553-576. GARDE, A.A., CHADWICK, B., GROCOTT, J., HAMILTON, M.A., MCCAFFREY, K.J.W. & SWAGER, C.P. 2002. Mid-crustal partitioning and attachment during oblique convergence in an arc system, Palaeoproterozoic Ketilidian orogen, southern Greenland. Journal of the Geological Society, London, 159, 247-261. GORDON, R.G. 1998. The plate tectonic approximation: Plate nonrigidity, diffuse plate boundaries, and global plate reconstructions. Annual Reviews of Earth and Planetary Sciences, 26, 615-642. GROCOTT, J. & TAYLOR, G.K. 2002. Magmatic arc fault systems, deformation partitioning and emplacement of granitic complexes in the Coastal Cordillera, north Chilean Andes (25°30'S to 27°00/S). Journal of the Geological Society, London. 159, 425-442. HODGES, K. 2000. Tectonics of the Himalaya and southern Tibet from two perspectives. Geological Society of America, Bulletin, 122, 324-350. JACKSON, J. 2002. Strength of the continental lithosphere: time to abandon the jelly sandwich? GSA Today, 12, 4-10. 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. KLOSKO, E.R., Wu, F.T., ANDERSON, H.J., EBERHARDTPHILIPS, D., MCEVILLY, T.V., AUDOINE, E., SAVAGE, M.K. & GLEDHILL, K.R. 1999. Upper mantle anisotropy in the New Zealand region. Geophysics Research Letters, 26, 1497-1500. MCCAFFREY, R. 1996. Estimates of modern arc-parallel strain rates in fore arcs. Geology, 24, 27-30. MOLNAR, P. 1988. Continental tectonics in the aftermath of plate tectonics. Nature, 335, 131-137. MOLNAR, P. 1992. Brace-Goetze strength-profiles, the partitioning of strike-slip and thrust faulting at zones of oblique convergence, and the stress-heat

flow paradox of the San Andreas fault. In: EVANS, B., & WONG, T.-F. (eds) Fault mechanics and transport properties of rocks, Academic Press, London, 435-459. MUELLER, ST., ANSORGE, J., EGLOFF, R. & KISSLING, E. 1980. A crustal shortening section along the Swiss geotraverse from the Rhine graben to the Po plain. Ecologae Geologicae Helvetiae, 73, 463-483. OLDOW, J.S., BALLY, A.W., AVE LALLEMANT, H.G. & LEEMAN, W.P. 1989. Phanerozoic evolution of the North American Cordillera: United States and Canada. In: BALLY, A.W. & PALMER, A.R. (eds) The Geology of North America - an Overview. Geological Society of America, Boulder, Colorado, 139-232. OLDOW, J.S., BALLY, A.W., AVE LALLEMANT, H.G. & LEEMAN, W.P. 1990. Transpression, orogenic float, and lithospheric balance. Geology, 18, 991-994. PRICE, R.A. 1981. The Cordilleran fold and thrust belt in the southern Canadian Rocky Mountains. In: Me CLAY & PRICE, N.J. (eds) Thrust and nappe tectonics. Geological Society, London, Special Publications, 9, 427-448. PRICE, R.A. 1986. The southeastern Canadian Cordillera: thrust faulting, tectonic wedging, and delamination of the lithosphere. Journal of Structural Geology, 8, 239-254. RANDALL, D., TAYLOR, G.K. & GROCOTT, J. 1996. Major crustal rotations in the Andean margin: palaeomagnetic results from the Coastal Cordillera of northern Chile. Journal of Geophysical Research, 101, 15783-15798. Russo, R.M., SILVER, P.G., FRANKE, M., AMBEH, W.G. & JAMES, D.E. 1996. Shear-wave splitting in northeast Venezuela, Trinidad and the eastern Caribbean. Physics of the Earth and Planetary Interiors, 95, 251-275. SAINT-BLANQUAT, M. (DE), TIKOFF, B., TEYSSIER, C. & VIGNERESSE, J.L. 1998. Transpressional kinematics and magmatic arcs. In: HOLDSWORTH, R.E., STRACHAN, R.A. & Dewey, J.F. (eds) Continental Transpressional and Transtensional Tectonics. Geological Society, London, Special Publications, 135, 327-340. SANDERSON, D. & MARCHINI, R.D. 1984. Transpression. Journal of Structural Geology, 6, 449-458. SAVAGE, M.K. 1999. Seismic anisotropy and mantle deformation: what have we learned from shear wave splitting? Reviews of Geophysics, 37, 65-106. SAVAGE, M.K., SVARC, J.L. & PRESCOTT, W.H. 1999. Geodetic estimates of fault slip rates in the San Francisco Bay area. Journal of Geophysical Research, 104, 4995-5002. SCHEUBER, E. & GIESE, P. 1999. Architecture of the central Andes - a compilation of geoscientific data along a transect at 21 °S. Journal of South American Earth Science, 12, 103—107. SILVER, P.G. 1996. Seismic anisotropy beneath the continents: probing the depths of Geology. Annual Review Earth and Planetary Sciences, 24, 385-432.

VERTICAL COUPLING AND DECOUPLING IN THE LITHOSPHERE TAYLOR, G.K., GROCOTT, J., POPE, A. & RANDALL, D.E. 1998. Mesozoic fault systems deformation and fault block rotation in the Andean forearc: a crustal scale strike-slip duplex in the Coastal Cordillera of northern Chile. Tectonophysics, 299,93-109. 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, 75-91. TEYSSIER, C., TIKOFF, B. & WEBER, J. 2002. Attachment between brittle and ductile crust at wrenching plate boundaries. In: BERTOTTI, G., SCHULMANN, K. & CLOETINGH, S. (eds) Continental Collision

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and the Tecton-Sedimentary Evolution of Forelands. European Geophysical Society, Special Publications, 1, 57—73. TIKOFF, B. TEYSSIER, C. & WATERS, C. 2002. Clutch tectonics and the partial attachment of lithospheric layers. In: BERTOTTI, G., SCHULMANN, K. & CLOETINGH, S. (eds) Continental Collision and the Tecton-sedimentary evolution of Forelands. European Geophysical Society, Special Publications, 1, 119-144. VAUCHEZ, A. & NICOLAS, A. 1991. Mountain-building strike-parallel motion and mantle anisotropy. Tectonophysics, 185, 183-211. VAUCHEZ, A., TOMMASI, A. & BARRUOL, G. 1998. Rheological heterogeneity, mechanical anisotropy, and tectonics of the lithosphere. Tectonophysics, 296,61-86.

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Strain modelling, seismic anisotropy and coupling at strike-slip boundaries: applications in New Zealand and the San Andreas fault M. K. SAVAGE1, K. M. FISCHER2 & C. E. HALL2'3 Institute of Geophysics, School of Earth Sciences, Victoria University of Wellington, New Zealand (e-mail: [email protected]) ^Department of Geological Sciences, Brown University, Providence, Rhode Island, USA 3 Present address: Seismological Laboratory, California Institute of Technology, Pasadena, California, USA 1

Abstract: Despite similar surface transform faulting behaviour, observed shear-wave splitting patterns in the California and New Zealand plate boundary regions are markedly different. To better understand the origin of the seismic anisotropy in these regions we model mantle flow and strain for a variety of strike-slip plate boundary scenarios. Simple relations between the flow or strain and elastic anisotropy are assumed to determine the integrated splitting in shear particle motion along teleseismic paths. Strain-controlled models fit the observations in New Zealand and California better than simplified flow-controlled models. Fast shear polarizations are progressively rotated toward the shear plane over time, and even a constant-viscosity model provides a good fit to the fast directions in New Zealand and southern California. The constant viscosity implies strong coupling between the surface and the deeper mantle. To fit the lack of decrease in delay times with distance from the fault, the relationship between delay time and strain must saturate at small strains. If this is the case, then strain in southern New Zealand and southern California may be small, equivalent to that achieved along an infinite fault by about 3-10 Ma of their present motion. Stratified viscosity allows more rapid rotation of fast directions toward faultparallel than occurs in isoviscous models, and can explain the nearly fault-parallel fast directions in the central South Island. Different aspects of the northern California results are fitted with different models, but a rapid change in viscosity with depth is needed to produce the full effects of the behaviour previously modelled as two layers of anisotropy suggesting vertical decoupling.

Seismic anisotropy measurements provide important clues to the extent of vertical coupling and decoupling in the mantle. SKS phases, which pass through the core as P waves and are converted to S waves at the core-mantle boundary almost directly beneath a station, have been used to determine anisotropy in the upper mantie asthenosphere and lithosphere through the method of shear-wave splitting or birefringence (Silver 1996; Savage 1999). The polarization of the first arriving wave, hereafter known as the 'fast direction' () to distinguish it from the polarization of the initial wave, is likely due to preferred orientation of olivine crystals, with , dt v. polarization do not fit N.SAF,S.SAFandNZ not well fitted S. SAP and S. AF: fit 0 v. fault distance for 3 Ma Central AF: fit for 800km displacement (23.5 Ma) Decrease in dt away from fault in model does not fit data N. SAP average fast directions fit, but 0, dt v. polarization do not fit Fit to SAP 0 v. fault distance reasonable; , dt v. polarization not fitted

Only central AF fitted Constant dt with distance from fault fits AF

Isoviscous

200 km & 300 km

4, 5a-c, 6, 7

200 km & 300 km

5d, e, 8a-c

WUS + SAP

As above

See Table 2

WUS only

As above

As above but with no weak zone

200km

8d, e

Oneside

As above

5 x 102U everywhere, except a factor of 81 lower from 0-80 km on right

200 km & 300 km

5g, h, 9a, c, d

Same as above Same as above

Same as above

WUS + PAC

As above

See Table 3

200km

5f, 9b, c, d

APM iso

Bottom: fixed Top: Fault-parallel: left (Pacific) side — 46 mm a"1, right (N.A.) 1 mm a"1 Fault-perpendicular: left 9.7 mm a"1, right —19A mm a"1

Isoviscous

200 km & 300 km

5i, j, 10, 11

APM WUS + SAP

As above

In Table 2

200 km & 300 km

11

APM WUS + PAC

As above

In Table 3

200 km & 300 km

11

AsthenEast

Bottom: 50 mm a 1 due E (—35 mm a"1 fault-parallel and +35 mm a"1 faultperpendicular on both sides) Top: Fault-parallel motion on left: — 11 mm a"1, right: +36 mm a""1 Across-strike component left: — 25.3 mm a"1, right: — 54.4 mm a"1 Bottom: Fault-parallel: 0 everywhere Faultnormal: +22 mm a"1 everywhere Top: Fault-parallel: left -57 mm a"1, right: —7 mm a"1 Fault-perpendicular: 0 everywhere

Isoviscous,

200 km & 300 km

Not shown

See Table 4

200km

Bot22mm

5k, 1, 12

Fit to SAP v. fault distance similar to WUS + SAP (f> gets more faultparallel further from the fault. If delay time does not depend on strain, does not change with distance from the fault - does not match data (/> is nearly faultperpendicular from 50-200 km from the fault - not a good fit Same as above

No viscosity models tested fit data well

(/>, dt v. polarization best fitted to twolayer model, cj) does not show quite as strong decrease with fault distance as data

Same as above 4> does not match measurements anywhere

approaches faultperpendicular at 200 km from fault. Does not fit well Flow at 200 km has reasonable fit to SAP 0 v. fault distance No viscosity models tested fit data well

Does not fit at all

22

M. K. SAVAGECTAL.

Fig. 6. Evolution of shear-wave splitting parameters over time, as a function of distance from the fault, calculated for 200 km to the isotropy/anisotropy boundary. Solid symbols with solid lines represent anisotropy calculations based on the numerical strain calculation from the analytic flow field. Open symbols with dashed lines represent corresponding anisotropy based on the finite element calculation of the flow field with a uniform viscosity. The solutions are symmetric across the fault, (a) Fast direction (f>. Values for dt < 0.4 s are not shown because their solution becomes unstable in the numerical codes, (b) Delay times also approximate the analytic solution. Symbols are same as in (a).

close to the fault (Fig. 5c) results in greater weight on the upper, fault-parallel layers. For the case of flow-controlled anisotropy, splitting parameters are constant at all distances across the fault and do not evolve with time. Fast directions line up with the shear direction, as shown by the line at 0 degrees for 'infinite time' in Fig. 6a. The constant delay times are caused by the constant anisotropy used in relating flow velocity to anisotropy. For the 11.8 Ma constant delay time-strain case, the results are similar to the variable delay time calculation near the fault. However, the delay times become larger and the fast directions become more fault-parallel further away

from the fault. This is due both to the rotation of the a-axes toward fault-parallel and to the decrease in plunge of the a-axes away from the fault, as shown in Fig. 5b. The independence of delay times on strain in these models means that splitting from the deep regions is not downweighted in the net splitting calculation as it is in the other cases. Note that the maximum anisotropy in a hexagonal symmetry based on olivine a-axes rotations is for propagation close to 45° to the vertical (e.g., fig. 4 in Savage 1999). The observation of variations in shear-wave splitting parameters as a function of incoming polarization at certain stations close to the San Andreas fault represent another test for the flow models examined here. Figure 7 shows predicted fast directions and delay times as a function of incoming polarization for the isoviscous models discussed above for a station 5 km from the fault. (Predicted splittings for a more complex viscosity structure are also shown and will be discussed later.) For comparison, theoretical curves for a two-layer anisotropic system and the data from station BKS on the San Andreas fault (Ozalaybey & Savage 1995) are included, rotated into the coordinate system of the model. The smooth rotation in finite strain and/or flow direction in the isoviscous models does not produce a strong variation in predicted fast direction as a function of polarization, although some aspects of the predicted delay times are reminiscent of the observed delay time trends. Because the fast direction predictions are significantly more robust and less dependent on modelling assumptions than the predicted delay times, we conclude that the isoviscous models do not provide a good fit to the observed polarization dependence. Effect of viscosity structure The Earth has a variable viscosity, and in particular it has been suggested that the San Andreas fault is very weak, with low viscosity (e.g. Flesch et al. 2000). Including variable viscosity has the effect of concentrating the strain in lower-viscosity regions (e.g. Turcotte & Schubert 1982). For very large strains in our models, anisotropy is relatively insensitive to viscosity structure because strain ellipsoids approach fault-parallel everywhere. For intermediate strains, the low-viscosity region tends to yield fault-parallel anisotropy and higher strains, and the higher-viscosity region yields fast directions at angles up to 45° to the fault and lower strains. We tested several forward

SEISMIC ANISOTROPY AND COUPLING

23

Fig. 7. Predicted splitting measurements for fast direction and delay time as a function of the polarization of the incoming wave (the back-azimuth in the case of SKS phases). Vertical incidence, stations at 5 km distance from the fault, and time steps equivalent to 11.8 Ma since the inception of San Andreas fault motion (corresponding to 400 km displacement across the fault) are used. Solid lines correspond to the two-layer solution presented in Figure 2 for station BKS at Berkeley, California, rotated so that the coordinate system is the one shown in Figure 3. Circles and stars with error bars are SKS and S measurements, respectively, from station BKS (Ozalaybey & Savage 1995) translated into the coordinate system of the fault. Patterns are periodic with a period of 90° in the incoming polarization direction. Different viscosity structures and anisotropic thicknesses lead to slightly different patterns (see text and Table 2). The delay time variations are similar to those of the BKS data, while the fast directions exhibit less variation in these finite element models than are observed in the data.

models of the San Andreas fault where we assumed simple but variable viscosity structures. Symmetric weak fault model (WUS + SAF). To estimate the viscosity structure near the San Andreas fault (SAF), we started with the vertical viscosity structure of the Western United States (WUS) constrained by lake level rebound in the nearby Basin and Range (Bills et al. 1994) (Table 2). In addition, a region of very low viscosity was imposed along the San Andreas fault, following the results of Flesch et al. (2000), by dividing the viscosity in the model by 27 at all depths for a region within 50 km of the fault (Fig. 8a, Table 2).

The biggest velocity gradients, and hence the biggest strain, are concentrated within the lowviscosity regions (Figs 8 & 5e). The surface strain is particularly concentrated along the fault, caused by the narrow lateral zone of low viscosity. This is exemplified in plots of the eigenvalue ratios and the dips and plunges of the resultant olivine a-axes (Figs 8 & 5d, e), calculated at 11.8 Ma. In this region (out to about 27 km from the fault), the azimuths of the maximum extension, and hence the a-axis azimuths, are close to fault-parallel and their plunges are nearly horizontal. Outside the low-viscosity fault region, the largest strain is at depths of 70-120 km. In general, regions of low strain correspond to

24

M. K. SAVAGE ET AL. Table 2. WUS + SAF viscosity model used Depth range of layer (km)

Viscosity away from fault (Pa s)

15-40 40-70 70-125 125-150 150-300 300-800

5 x 1020 5 x io20-4 x 1017 4 x 1017 4 x 1017-2 x 1020 2 x 1020 1 x 1019

regions with the fast direction at —45° (135°) to the fault with variable plunges, while high-strain regions have a-axes closer to fault-parallel and plunges approaching horizontal. However, the details of the strain field become quite complicated at shallow depths. The fast directions predicted for this model are nearly fault-parallel both far from the fault and close to the fault. But between 0 and 37 km from the fault, however, the fast directions become progressively further from fault-parallel for the 11.8 and 6 Ma cases, reaching —25° at 40 km for 6 Ma (200 km displacement; Fig. 8). The case with hexagonal anisotropy and delay times independent of strain also becomes more fault-parallel further from the fault, similar to what occurred with the above i so viscous case. This can be explained by the behaviour of the azimuths of the maximum extension direction (Fig. 5d). For stations within the lateral bounds of the low-viscosity region, anisotropy in the uppermost regions of the model contribute the most to the predicted cumulative splitting parameters, both inherently (e.g. Saltzer et al.

Viscosity next to fault (Pa s)

1.9 1.9 4.8 4.8 9.5 4.8

x x x x x x

1019 1019-4.8 x 1015 1015 1015-9.5 x 1018 1018 1017

Lateral distance to fault boundary (km)

50 50 50 50 50 50

2000), and also because strains are largest at shallow depths. The obvious exception is the test case where anisotropy is not scaled with strain. In the low-viscosity region, the azimuths change from —45° at depths of 150km to fault-parallel at the surface, but because strains are larger at or near the surface, the resulting fast directions lie close to fault-parallel. Past 57 km from the fault, the strains are largest between 60 and 120km depth, where the azimuths are fault-parallel, leading to fast directions within 5° of fault-parallel. The delay times are largest in the low-viscosity region near the fault because the strains are largest there. Even though there is a sharp change in fast direction with depth, the resulting splitting parameters do not show a strong two-layer effect in the fast directions, largely because the strain favours the near-surf ace layers so much that they contribute more to the splitting and do not yield two-layer type behaviour (Fig. 7). For comparison, we also show a case for 23.5 Ma of strain (800 km displacement) with the WUS viscosity but no weak zone near the

Fig. 8. (a) Contours of the velocity field as in Figure 4 for the WUS + SAP viscosity model of the San Andreas fault region (Table 2). (b,c) Time evolution of shear-wave splitting predicted for the WUS + SAP model, as in Figure 6. (d,e) Time evolution of shear-wave splitting predicted for the WUS model without a weak zone near the fault.

SEISMIC ANISOTROPY AND COUPLING

25

Time Evolution WUS+SAF Model 300 km to anisot/isot

Fig. 8. Continued.

fault (Fig. 8d, e). It is similar to the isoviscous case and the weak-fault case just discussed. However, in the strain-controlled case, the fast directions near the fault are closer to faultparallel than for the same time evolution as in the other two models. The case where delay times are independent of strain yields similar delay times and fast directions to the corresponding case in the other two models.

Asymmetric viscosity models. The assumption of symmetric viscosity structures across a fault is unlikely to be met in many instances. The oceanic Pacific Plate (PAC) may have higher viscosity in shallow regions than the continental North American Plate because of the weaker materials in continental crust. Teyssier and Tikoff (1998) have suggested that the plate boundary shearing extends into the asthenosphere

26

M. K. SAVAGEETAL.

Fig. 8. Continued.

in a narrow region on the North American side to yield the observed splitting measurements. We have considered two cases of asymmetric viscosity. In the first model (oneside), a uniform viscosity base is reduced by a factor of 81 if it is shallower than 90 km and within 70 km of the fault on the eastern side (Fig. 9, Table 1). The second model (WUS -h PAC) uses the viscosity structure discussed above on the North American

plate (Bills et al. 1994), and an approximation to viscosity profiles for oceanic lithosphere at 20 Ma age (Hirth & Kohlstedt 1996) for a mantle that is depleted of water content due to melting at a mid-ocean ridge (Table 3). The two models both yield velocity fields with velocity gradients, and hence strain, that are largest near the fault (Figs 9 & 5h). The strain in both models is similar to that in the WUS + SAP

SEISMIC ANISOTROPY AND COUPLING

27

Fig. 9. As in Figure 8, but for viscosity structures that differ across the fault. Contours are constant velocity. (a) A simplified viscosity structure, with a base viscosity of 5 x 1020 Pa s. The viscosity is a factor of 81 lower in a block on the right side of the fault, extending from the surface to a depth of 90 km and extending 70 km in width, (b) The material on the right side of the fault uses the viscosity structure determined by Bills et al. (1994) (Table 2) for the western USA, while the material on the left of the fault uses an approximation to the Hirth & Kohlstedt (1996) viscosity profiles (Table 3) to match the Pacific plate parameters, (c) The resultant fast directions on a profile across the model in parts (a) and (b). (d) Corresponding plots of fast direction and delay time as a function of the incoming polarization direction, as in Figure 7, for the Oneside model using an anisotropic thickness of 300 km.

model just discussed: close to the fault (within about 30km), the strain increases rapidly toward the surface, while far from the fault (past about 60 km) the strain either peaks between 120 and 70km depth (WUS + PAC), or is constant near a value of 1 (Oneside). The WUS + PAC model (Fig. 5f) has a very similar pattern of a-axis orientation with depth to the WUS + SAF model (Fig. 5d). The Oneside model shows less dependence on distance, with a > —90° in dextral oblique convergence, which causes homogeneous transpression in the deforming zone. The shear plane parallels the xz plane, (c) a = 0° for dextral wrenching (+180° for sinistral wrenching), which causes homogeneous simple shear in the deforming zone. The shear plane parallels the xz plane, (d) 0° < a < 90° for dextral oblique divergence, which causes homogeneous transtension in the deforming zone, (e) a = 90° for divergence. Divergence causes vertical thinning and homogeneous pure shear in the deforming zone. Flow is confined to the yz plane.

86

S. GIORGISCTAL.

Fig. 3. Rotation of material lines in the horizontal (xy) plane during deformation, based on work by Ramberg (1975), Fossen et al (1994) and Passchier (1997). (a) In convergent pure shear, the repulsor flow apophysis parallels the y-axis, and the attractor flow apophysis parallels the ;c-axis. Half the material lines rotate clockwise, and half rotate counterclockwise, (b) In transpression, the repulsor parallels the direction of convergence, and the attractor parallels the x-axis. Most, but not all, lines rotate clockwise, (c) In simple shear, the z-axis acts as both an attractor and a repulsor, and all lines rotate clockwise, (d) In transtension, the repulsor parallels the z-axis, and the attractor parallels the direction of divergence. Most, but not all, lines rotate clockwise, (e) In divergent pure shear, the repulsor parallels the Jt-axis, and the attractor parallels the y-axis. Half the material lines rotate clockwise, and half rotate counterclockwise.

not rotate: those that are initially parallel to the xaxis and those that are initially parallel to the direction of convergence (at 90° in Fig. 3a and 135° in Fig. 3b). Other material lines rotate towards the jc-axis and away from the direction of convergence. The jc-axis is an 'attractor' flow apophysis, and the direction of convergence is a 'repulsor' flow apophysis (Ramberg 1975; Fossen et al. 1994; Passchier 1997). During simple shear (Fig. 3c), lines that are initially parallel to the *-axis (the shear plane) do not rotate, and all other horizontal lines rotate in the same direction (clockwise in the example shown here) towards this orientation. The ;t-axis is therefore both an attractor and a repulsor. During oblique divergence (Fig. 3d) and divergent pure shear (Fig. 3e), two orientations of lines again do not rotate: those that are initially parallel to the *-axis and those that are initially parallel to the direction of divergence (at 45° in Fig. 3d and 90° in Fig. 3e). In this case the direction of divergence acts as an attractor and the Jt-axis acts as a repulsor. In all deformations except for simple shear, the sense of rotation of a material line (clockwise or counterclockwise) depends on the initial orientation of the line. Rigid objects embedded in viscous, deforming media are also sensitive to fabric attractors and repulsors (Jeffery 1922). Rotation rates for rigid ellipses are well established in simple shear and pure shear (Ghosh & Ramberg 1976; Willis

1977), and through sparse work on other kinematic regimes such as transpression and transtension (Freeman 1985; Jezek et al 1994, 1996). The mathematics used in this study to calculate the orientations of rigid clasts in transpression (and transtension) are described in Giorgis & Tikoff (2004). The behaviour of rigid clasts with low aspect ratios (< 10:1) differs from that of material lines in that there are no orientations in which rigid ellipses do not rotate (Ghosh & Ramberg 1976). Figure 4 shows two examples in transpression for an ellipse whose aspect ratio is 1.75:1. Figures 2b and 3b are useful for comparison because they explore the same boundary condition: transpression related to oblique convergence with a = — 4 5 ° . For material lines, the sense of rotation depends on the initial orientation (Fig. 3b). For example, a material line initially oriented at 20° from the positive x-axis rotates clockwise, but one at 160° rotates counterclockwise. But for the lowaspect-ratio rigid ellipses that we consider in dextral transpression, the sense of rotation is always clockwise (Fig. 4b). This observation is not true for all low-aspect-ratio rigid objects in transpression. For example, ellipses of this aspect ratio show both senses of rotation during transpression related to higher angles of convergence (a closer to —90°). The aspect ratio of a rigid object also exerts a first-order control on its rotation rate and stable positions (Jeffery 1922; Ghosh & Ramberg

CRUSTAL BLOCK ROTATION BY MANTLE FLOW

Fig. 4. Rotation of a rigid ellipse with an aspect ratio 1.75:1, in the horizontal plane during transpression, based on work by Ghosh & Ramberg (1976). (a) In highly oblique transpression (a — — 15°), ellipses of any initial orientation rotate clockwise. An ellipse oriented parallel to the fabric attractor (0° or 180°) or repulsor (165°) rotates more slowly than others, (b) In less oblique transpression (a = -45°), ellipses of any initial orientation rotate clockwise. An ellipse oriented parallel to the fabric attractor (0° or 180°) or repulsor (135°) rotates more slowly than others.

1976). Equant rigid objects rotate at a rate dependent only on the vorticity and strain rate of the viscously deforming medium. The rotation rates of low-aspect-ratio rigid objects (10:1) rotate similarly to material lines. For example, higher-aspect-ratio rigid objects than the one considered in Figure 3b show both senses of rotation during transpression related to the same boundary conditions (a = —45°). Experimental apparatus and design Our experimental apparatus (Fig. 5) is described in Venkat-Ramani & Tikoff (2002). These experiments use a rubber sheet and silicone gel to distribute simple shear, pure shear, transpression or transtension homogeneously in a basal rubber sheet. The rubber sheet stretches to approximately twice its original width without buckling or breaking. In these experiments, we covered the rubber sheet with an even thickness of silicone (RD-20; Rhone Poulenc), which behaves in a Newtonian fashion (viscosity = 1 x 104 Pa s)

87

and sticks to the underlying sheet. The initial thickness of this layer of silicone was approximately 2 cm. We embedded two different kinds of rigid objects into the surface of the silicone and gathered a variety of data from video stills in order to document the experiments. First, we performed a series of experiments involving matchsticks, which have an aspect ratio of ^25:1 (Fig. 6). On each video still, we measured matchstick orientation (/3) and the position of the moving plate. Comparing the initial width (w) of the deforming zone in the y direction, the change in position (dy) of the moving plate in the y direction and the change in position (dx) of the moving plate in the x direction, we calculated a dimensionless measure of plate displacement:

From 100 Ma cooling ages and moderate-P rocks with 55 km) and density of the eastern zone could have supported average surface elevations of >4km over a broad region (Schmidt 2000). In contrast, the dense, relatively little denuded, and presently anomalously thick western zone probably could not have supported more than 1 km of average surface elevation during the Cretaceous. We therefore infer that the eastern PRB represented a zone of high elevation in the Late Cretaceous. Comparison with the modern Andean continental arc is useful for assessing the relationship between crustal thickening and topography in the North American Cordilleran arcs. In the Andes, the ~4km high, 350-400 km wide Altiplano and Puna Plateaux were created by thickening of lithosphere during subduction.

Crustal thickness varies, but is generally 5070 km (Zandt et al 1994). The plateau region is bounded on the eastern side by a thrust belt, and on the west by the arc. The Altiplano-Puna Plateau comprises varied and complex basement geology, as does the North American Cordillera, and has been uplifted across a wide area due to the thermo-mechanical effects of contraction of hot, thermally weakened lithosphere. Regional uplift of the plateau regions has been attributed to an increase in plate convergence rate, which led to a decrease in the angle of subduction, and thinning/weakening of the South American lithosphere (Allmendinger et al 1997). In the Cordillera, the CMC and Sierran belts likely comprised the western margins of broad plateau regions, similar to the relation between the Andean arc and the Altiplano-Puna Plateau. The PRB may have represented the southern end of the plateau, analogous to the narrow, southern end of the Altiplano-Puna Plateau. It is difficult to identify the northern end of the plateau because of lack of data and effects of faulting in the northern Cordillera. These observations and their similarities to inferences about the Mesozoic-Early Cenozoic North American arcs support the idea that plateau formation commonly accompanies construction of continental arcs. Unroofing mechanisms The unroofing of the North American Cordilleran arcs and inboard regions can be described in terms of the collapse of the plateau region (east of the

CORDILLERAN ARC AND OROGENIC PLATEAUX

arcs) and the unroofing of the arcs. The former is characterized by metamorphic core complex development and, in the central and southern Cordillera, by Basin and Range style extension. Unroofing of the arcs occurred by a combination of erosion and syn- to post-convergence extension (transpression, then transtension). The balance between these processes varied in the northern v. the central/southern Cordilleran arcs. The Omineca-Sevier belt may have collapsed due to the presence of hot, thickened crust (Livaccari 1991) weakened either by magmatism (Armstrong & Ward 1991) or by partial melting at depth (Vanderhaeghe & Teyssier 2001). The exhumation of deep middle crust (~20-50km depth in a thickened orogen) may also have been influenced by the presence of lowerdensity rocks at depth, due to the emplacement of denser terranes on felsic basement, and/or to the generation of partial melting at depth during heating and decompression (Teyssier & Whitney 2002). These are all internal processes that are consequences of crustal thickening. The arcs also consisted of hot crust of similar thickness to the internal zone, but only the northern belt experienced major collapse and extension. The Sierra and PRB arcs may have been comparable to the CMC in terms of crustal thickening and magmatic history, but they did not experience major extension, and only underwent about half the magnitude of exhumation. This observation is incompatible with a model of gravitational collapse driving unroofing of the arcs, because collapse of hot, thickened crust did not occur on a large scale in the central/southern Cordilleran arcs. This suggests that external mechanisms for driving exhumation may have been important in the north but not in the central and southern regions. External mechanisms include a change in plate motion factors in the Eocene; for example, plate velocity, trajectory and/or number of small plates in the northern Pacific. Gravitational collapse of thickened crust and changes in Pacific plate motions have been proposed to account for extension in the northern Cordillera. For example, rearrangement of Pacific plates to more oblique convergence relative to North America, consumption of the Kula plate and/or a slowing of plate convergence rate in the Eocene (Engebretsen et al 1985; Stock & Molnar 1988) may have caused extension throughout the northern Cordillera. Plate boundary kinematics resulted in a NW-SE minimum principal stress component, consistent with the orientation of syn-extensional structures in the metamorphic terrains of this region. A change in plate boundary stresses (e.g. interaction of the

173

North American margin with the Pacific-Kula— Farallon triple junction) has also long been proposed as a mechanism for Eocene magmatism in the northern Cordillera (Ewing 1980). These observations are consistent with geodynamic coupling between the subducting plate and the continental lithosphere, but the nature of the link cannot be inferred from the observations.

Cordilleran crustal thickening, deep crustal flow and unroofing Previous North American Cordilleran studies (e.g. Coney & Harms 1984; Dilek & Moores 1999) have used exposures of metamorphic rocks to model palaeo-crustal thickness. In these models, regions with exposures of highgrade metamorphic rocks (Omineca-Sevier belt) are depicted as having thicker Late Cretaceous crust than regions with few or no exposures of mid-crustal rocks. For example, these studies show thinner crust (5000 km2) of exhumed Mesozoic lower crustal granulites. These exposures not only provide us with an important natural example of the variability of lower crustal strength profiles, but also allow us to examine directly the effects of this variability on the evolution of lower crustal fabrics and shear zones. Through work related to this and our previous studies (Klepeis et al 1999, 2001, 2003; Clarke et al 2000; Daczko et al 2001<s, b, 2002), we identified a continuous lower crustal section in Fiordland representing Early Cretaceous paltoodepths of 25-50 km. This exposure, excellent geochronological control and regional-scale crosscutting relationships enabled us to determine how the lower crustal section evolved structurally within a well-constrained, ~25 Ma period of time (130-105 Ma).

Geological setting The geology of SW New Zealand records a history of magmatism, high-grade metamorphism and convergence that accompanied the evolution of an Early Mesozoic magmatic arc along the ancient margin of Gondwana (J.D. Bradshaw 1989; Muir et al 1994; Daczko et al 20010). Part of this arc is thought to have formed initially outboard of the Gondwana margin during Early Triassic-Early Cretaceous (247-131 Ma) sub-

duction-related magmatism (Kimbrough et al 1994; Muir et al 1998; Tulloch & Kimbrough 2003). In Fiordland (Fig. Ib), a north- and NNEstriking belt of plutonic, volcanic and sedimentary rocks called the Median Tectonic Zone (Tulloch & Kimbrough 1989; Muir et al 1994) or the Median Batholith (Mortimer et al 1999) represents this outboard part of the arc. On the eastern side of the Median Tectonic Zone (MTZ), Triassic plutons intruded Permo-Triassic volcanoclastic rocks of the Brook Street volcanics and the Maitai terrane (Fig. 2). These latter units represent arc-derived sedimentary rocks and accretionary complexes that also lay outboard of Gondwana during the Early Mesozoic (J.D. Bradshaw 1989; Mortimer et al 1999). Following its initial construction, the outboard part of the Mesozoic arc (the MTZ) accreted onto the Gondwana margin during convergence. This stage of arc evolution resulted in the emplacement of 126-105 Ma intrusive rocks (ages from Kimbrough et al 1984; Mattinson et al 1986; McCulloch et al 1987; J.Y. Bradshaw 1989; Tulloch & Kimbrough 2003; Klepeis et al 2004) into crust composed of both MTZ and Gondwana margin rocks (Mortimer et al 1999). Mafic-felsic rocks of the Western Fiordland Orthogneiss and the Separation Point Suite are included in this event (Fig. Ib). In addition, the ages of mafic dykes that intrude Lower Paleaozoic host gneiss in the Arthur River complex (Fig. 2) suggest that the amalgamation of the outboard arc with Gondwana may have occurred as early as ~ 136 Ma and probably by -129 Ma (Hollis et al 2002). In Fiordland, rocks of the Gondwana margin are represented by metasedimentary rock and granitic orthogneiss of Cambrian-Permian age (Fig. Ib). Intense contractional deformation and crustal thickening accompanied and followed magmatism that resulted in the emplacement of the Western Piordland Orthogneiss (Bradshaw 1990; Muir etal 1995, 1998; Daczko etal 2001(3, 2002). On the basis of geochronological and geochemical data, Muir et al (1995, 1998) suggested that the sudden appearance of large volumes of Na-rich magma within the arc at —126 Ma, including the Western Fiordland Orthogneiss, was triggered tectonically by the underthrusting and subsequent melting of MTZ rocks beneath western Fiordland. Daczko et al (2001a, 2002) presented structural data that support this interpretation. However, the origin of this contraction is controversial. The deformation may reflect the collision of the outboard part of the arc (the MTZ) with Gondwana following emplacement of the Western Fiordland Orthogneiss (Bradshaw 1990; Bradshaw &

MID-LOWER CRUSTAL ATTACHMENT ZONE

199

Fig. 1. (a) Location of Fiordland on the South Island of New Zealand, (b) Geological map of Fiordland showing regional subdivisions. Pressures representing the peak of Early Cretaceous granulite facies metamorphism at ~120 Ma show tilted lower crustal section constructed using metamorphic data in Table 1 and, for the Doubtful Sound area, in Gibson & Ireland (1995). (c) Schematic cross-section showing a convergent setting between the Gondwana continental margin and an Early Mesozoic arc (MTZ). WFO, Western Fiordland Orthogneiss batholith.

Kimbrough 1989; Daczko et al 2001^, 2002). Alternatively, the deformation could reflect the collision of a basaltic plateau with the subduction zone on the outboard side of the arc (Sutherland & Hollis 2001; Tulloch & Kimbrough 2003). In addition to the deformation, crustal thickening is reflected in the occurence of high-pressure granulite facies and upper amphibolite facies metamorphism that recrystallized the Western Fiordland Orthogneiss and its host rock (Blattner 1978; J.Y. Bradshaw 1989; Clarke et al 2000;

Daczko et al. 2001/?). Garnet-pyroxeneplagioclase-bearing assemblages have yielded peak metamorphic pressures of P = 1216kbar and temperatures of r > 7 5 0 ° C (Table 1). These assemblages occur within the Western Fiordland Orthogneiss and its host rocks between Doubtful Sound and Milford Sound (Fig. 1). The Early Cretaceous age of this metamorphism is constrained by three relationships: (1) crystallization ages (126120 Ma) of the Western Fiordland Orthogneiss

200

K. A. KLEPEIS ETAL

Fig. 2. Geological map of north-central Fiordland showing the major lithological subdivisions. Map was constructed using data from this study and from Bradshaw (1989, 1990), Blattner (1991), Turnbull (2000), Daczko et al. (2002) and Claypool (2002). Plotted dates (see also Table 2) are from U-Pb analyses of zircon cores and rims reported by Hollis et al. (2002). Data show the distribution of Early Cretaceous crystallization (zircon core) ages and ~120 Ma metamorphic (rim) ages from the Arthur River complex. Also note distribution of migmatite (m) below the batholith. MD, Mount Daniel; ME, Mount Edgar; P, Pembroke Valley; CO, Camp Oven Creek; ARC, Arthur River complex; MTZ, Median Tectonic Zone. Profiles A-A', B-B', C-C, D-D', E-E' and F-F are shown in Fig. 3.

(Table 2); (2) metamorphic ages from the Arthur River complex (Fig. 2) and adjacent units; and (3) crosscutting relationships between highgrade fabrics and dykes of known age (Mattinson et al. 1986; McCulloch et al. 1987; Gibson & Ireland 1995; Tulloch et al. 2000; Hollis et al. 2002; Klepeis et al. 2004). The ages of specific fabrics and rock units discussed in this chapter are presented in more detail below. By ~ 105 Ma, widespread extension affected parts of the Fiordland belt and adjacent areas (ID. Bradshaw 1989; Tulloch & Kimbrough

1989; Gibson & Ireland 1995). The Doubtful Sound shear zone (Fig. Ib) is the dominant structural expression of this extension in Fiordland. This shear zone is composed of granulite facies and upper amphibolite facies fabrics that cut the Western Fiordland Orthogneiss. Kinematic data reported by Gibson et al. (1988) and Claypool (2002) indicate that this shear zone records NE-SW stretching and a dominantly NE-directed sense of shear that is consistent with extension directions in other parts of New Zealand during the Mid-Cretaceous (Tulloch & Kimbrough

Table 1. P-T data from northern Fiordland between Caswell and Milford Sounds Location

Lithological unit*

Assemblage

Calculated P (kbar)

Calculated T (°C)

P-T method*

Data source

Interpretation

g-bi-hbl-pl-q

4-5

500-600

6,8,9

Daczko et al (2002)

Palaeozoic tectonism

Pembroke Valley

Country rock outside WFO contact aureole

ARC

opx-cpx-hbl

750

see source

Early stages of

George Sound

WFO contact aureole Country rock rafts in WFO

g-bi-hbl-pl

8.7 ± 2.1

470, 482, 604,

6,8,9

g-bi-pl-ky-q

11-12

704-735

see source

g-bi-ksp-pl

7-8

>700

6,8,9

Charles Sound

WFO and contact aureole

WFO

g-cpx-pl-q

8-10

Mount Daniel & adjacent areas

WFO and ARC

g-cpx-pl-ksp-rt-q

12-13

650-700

see source

Milford Sound

ARC

to

2-cpx-pl-q-kv r r M J

13-15

730-870

1, 2, 3, 4, 5, 7

Milford Sound

ARC

g-cpx-pl-q

14, 16

715-812

1, 2, 3, 4, 5

Milford Sound

ARC

g-cpx-hbl-pl-q

13-15

800

1, 2, 3, 4, 5

Mount Daniel Milford Sound

ARC and WFO ARC

g-cpx-ky-pl-hbl-q g-cpx-pl-q

13.9 11.8-13.7

700-800 624-775

1,2,5 1, 2, 3, 4, 5

Poison Bay

ARC

g-cpx-hbl-pl-q

12.4-15.6

768-820

1, 2, 3, 4, 5

Pembroke Valley

ARC

w g-cpx-pl-q to .f -^ F i

14.7-16.6

662-700

1, 2, 3, 4, 5, 6

Mount Daniel & adjacent areas Poison Bay

ARC and WFO

g-cpx-bi-pl

11-12

662-766

see source

WFO

g-cpx-pl-q

12.6 ± 1.9

718+ 115

8

Clarke et al (2000) Daczko et al. (2002) Bradshaw (1985, 1989) Daczko et al. (2002) Bradshaw (1985) Bradshaw (1985, 1989) Clarke et al. (2000) Clarke et al. (2000) Clarke et al. (2000) Daczko (2001) Clarke et al. (2000) Clarke et al. (2000) Clarke et al. (2000) Bradshaw (1989) Clarke et al. (2000)

Caswell Sound

George Sound & Mount Daniel areas Caswell Sound

686

see source

arc

Stages 1 and 2 Stages 1 and 2 Stages 2 and 3 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 Stage 2 (continued}

Table 1. Continued Location

Lithological unit*

Assemblage1

Calculated P (kbar)

Milford Sound

ARC

g-hbl-pl-q-ky

13-16

Milford Sound

ARC

15.8 ± 2.6

Pembroke Valley

ARC

Pembroke Valley

ARC

g-hbl-pl-bi-epky-q g-cpx-pl-rt-hblcz-q g-bi-pl-rt-hbl-cz-q

Pembroke Valley & Mount Daniel Pembroke Valley & Mount Daniel Pembroke Valley & Mount Daniel Poison Bay

ARC

Poison Bay

Calculated T (°Q 674

P-T method* 6, 7

Data source

Interpretation

14.0 ± 1.3

676 + 34

1, 3, 4, 6, 8, 11

14.1 ± 1.2

674 + 36

1, 3,4, 6, 8, 11

g-hbl-cz-pl-q-ky

11.7-13.3

677 + 64

6,7, 8, 11

Clarke et al. (2000) Clarke et al (2000) Daczko et al (2001«) Daczko et al. (20010) Daczko (2001)

ARC

g-hbl-pa-pl-cz-q

11.4, 11.1, 13.2 ± 1.7

687-694, 703-736

6, 8, 11

Daczko (2001)

Stage 3, cooling

ARC

g-hbl-mu-pl-cz-q

10.9, 11.9+ 1.8

704, 629 ± 56

6, 8, 11

Daczko (2001)

Stage 3, cooling

ARC

g-pl-bi-hbl-rt-cz

11.9+ 1.1

581 + 34

8

Exhumation

ARC

g-hbl-pl-bi-ti-cz

8.7 + 1.2

587 + 42

8

Klepeis et al. (1999) Klepeis et al. (1999)

8

Stage 3 Stage 3 Stage 3, duplex Stage 3, duplex Stage 3, cooling

Exhumation

*ARC, Arthur River complex; WFO, Western Fiordland Orthogneiss; results organized according to stages discussed in the text (oldest to youngest). f g, garnet; cpx, clinopyroxene; opx, orthopyroxene; pi, plagioclase; q, quartz; ky, kyanite; bi, biotite; ep, epidote; hbl, hornblende; rt, rutile, cz, clinozoisite; ksp, potassium feldspar; ti, titanite; pa, paragonite; mu, muscovite. *1, Ellis & Green (1979); 2, Powell (1985); 3, Krogh (1988); 4, Eckert et al. (1991); 5, Newton & Perkins (1982); 6, Graham & Powell (1984); 7, Newton & Haselton (1981); 8, Powell & Holland (1988); 9, Ferry & Spear (1978); 10, Hodges & Spear (1982); 11, Kohn & Spear (1990).

Table 2. Geochronological data from intrusive rocks in northern Fiordland* Lithological unit1

Rock type*

Method*

Mineral

Corrected age (Ma)

Dyke in Milford Gneiss Milford Gneiss

pegmatite

U-Pb ion probe

zircon

81.8 ± 1.8

metadiorite

K-Ar

hornblende

90.3 ± 2.6

Milford Gneiss

metadiorite

K-Ar

hornblende

91 ± 0.4

Milford Gneiss

diorite gneiss

K-Ar

hornblende

92.3 ± 2.8

WFO

metadiorite

U-Pb SID

apatite

91 ±2

WFO

amphibolite

K-Ar

hornblende

92.6 ± 0.9

WFO

metadiorite

U-Pb SID

apatite

92.6 ± 2

Milford Gneiss

hbl gneiss

K-Ar

hornblende

115-74

Milford Gneiss Milford Gneiss

hbl gneiss hbl gneiss

K-Ar K-Ar

hornblende hornblende

105 105 ± 4

Milford Gneiss

gneiss

K-Ar

hornblende

106.4 ± 0.82

WFO

diorite gneiss

U-Pb ion probe

zircon

107.5 ± 2.8

Milford Gneiss

hbl granulite

K-Ar

hornblende

111 ± 2

Milford Gneiss

dioritic granulite

K-Ar

plagioclase

113.4 ± 1.8

Darran complex

gabbronorite

U-Pb SID

apatite

113 ±0.4

MTZ

leucogranite

K-Ar

biotite

114.8 + 5.1

Interpretation"

Source Hoflis et al (2002) Nathan et al. (2000) Nathan et al. (2000) Nathan et al. (2000) Mattinson et al. (1986) Gibson et al. (1988) Mattinson et al. (1986) Blattner(1991) Blattner (1978) Nathan et al. (2000) Nathan et al. (2000) Gibson & Ireland (1995) Nathan et al. (2000) Nathan et al (2000) Mattinson et al. (1986) Williams & Harper (1978)

C or inherited

Exhumation phase

Cooling

Exhumation phase

Cooling

Exhumation phase

Cooling

Exhumation phase Exhumation phase

Cooling

Exhumation phase Exhumation phase

M/cooling M M

End Stage 3 exhumation End Stage 3 End Stage 3

M

End Stage 3

M

Stage 3

M

Stage 3

M

Stage 3 Stage 3

C

Stage 3

(continued}

Table 2. Continued Lithological unitf

Rock type*

Method§

Mineral

Corrected age (Ma)

MTZ

trondhjemite

K-Ar

biotite

115.8 ± 3.1

MTZ

trondhjemite

K-Ar

muscovite

115.8 ±4.1

Harrison Gneiss

orthogneiss

U-Pb ion probe

zircon rim

120 ± 2

Palaeozoic gneiss in ARC1 Selwyn Creek Gneiss11 Dyke in Milford Gneiss1 WFO

migmatitic gneiss

U-Pb ion probe

zircon rim

120

dioritic orthogneiss felsic dyke (02DA)

U-Pb ion probe

zircon rim

120

U-Pb ion probe

zircon rim

123

amphibolite

U-Pb ion probe

zircon

119 ± 5

WFO

orthogneiss

U-Pb SID

zircon

120-130

WFO

granulite

Rb-Sr

whole rock

120 ± 15

WFO

orthognesis

U-Pb ion probe

zircon

125.9 ± 1.9

WFO

granulitic metabasite

U-Pb ion probe

zircon

126 + 3

Harrison Gneiss

orthogneiss

SHRIMP

zircon

134 + 2

Pembroke Granulite Mount Edgar Diorite1

mafic granulite

K-Ar

hornblende

130

dioritic orthogneiss

U-Pb ion probe

zircon

128.8 ± 2.4

Source Williams & Harper (1978) Williams & Harper (1978) Tulloch et al (2000) Hollis et al (2002) Hollis et al (2002) Hollis et al (2002) Gibson & Ireland (1995) Mattinson et al (1986) McCulloch etal (1987) Muir et al (1998) Gibson & Ireland (1995) Tulloch et al (2000) Blattner(1991) Hollis et al (2002)

Interpretation" C

Stage 3

C

Stage 3

M

Stages 1 and 2

M

Stages 1 and 2

M

Stages 1 and 2

M

Stages 1 and 2

C

Stages 1 and 2

C

Stages 1 and 2

C

Stages 1 and 2

C

Stages 1 and 2

C

Stages 1 and 2

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

Dyke in Milford Gneiss11 Milford Gneiss1

felsic dyke (02DA)

U-Pb ion probe

zircon core

133.4 ± 2.1

U-Pb ion probe

zircon

131.3 ± 2.9

U-Pb ion probe

zircon

132.6 ± 2.9

Dyke in Milford Gneiss1 Milford Gneiss

gabbroic orthogneiss dioritic orthogneiss diorite dyke

U-Pb ion probe

zircon

135.9 ± 1.8

granulite

K-Ar

hornblende

140.1 ± 1.02

Milford Gneiss

granulite

K-Ar

hornblende

138 ± 4

Milford Gneiss

hbl schist

K-Ar

hornblende

149 + 4

Milford Gneiss

amphibolite

K-Ar

hornblende

164 ± 4

Selwyn Creek Gneiss1 Darran complex

dioritic orthogneiss diorite

U-Pb ion probe

zircon core

154.4 ± 3.6

K-Ar

biotite

134 + 4

Darran complex

gabbronorite

SID

apatite

135 ± 0.4

Darran complex

diorite

K-Ar

biotite

135 + 4

Darran complex

U-Pb SID

apatite

135.6 ± 0.4

Darran complex

olivine gabbronorite felsic dyke

U-Pb ion probe

zircon

136.8 + 1.9

Darran complex

gabbronorite

U-Pb SID

zircon

137 ±4

Darran complex

gabbronorite

U-Pb SID

zircon

137 ± 1

Darran complex

diorite

U-Pb ion probe

zircon

138 ± 2.9

Milford Gneiss11

Hollis et al (2002) Hollis et al (2002) Hollis et al. (2002) Hollis et al (2002) Nathan et al (2000) Nathan et al (2000) Nathan et al (2000) Nathan et al (2000) Hollis et al (2000) Nathan et al (2000) Mattinson et al (1986) Nathan et al (2000) Mattinson et al (1986) Muir et al (1998) Kimbrough et al (1994) Mattinson et al (1986) Muir et al. (1998)

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Stages 1 and 2

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc (continued}

Table 2. Continueaf Lithological unit1

Rock type*

Met!i0df

Mineral

Corrected age (Ma)

Darran complex

leucogabbro

K-Ar

biotite

140 ±2

Darran complex

monzite

U-Pb SID

zircon

141 ± 4

Darran complex

diorite

U-Pb SID

zircon

142 ± 5

Darran complex

diorite

K-Ar

hornblende

142 ±4

Darran complex

dioritic pegmatite

K-Ar

hornblende

142 + 6

Darran complex Darran complex

dioritic pegmatite diorite

K-Ar K-Ar

hornblende biotite

142.8 + 1.8 143.82 ± 1.2

Darran complex

dioritic pegmatite

K-Ar

hornblende

145.2 ± 3

Darran complex

diorite

K-Ar

biotite

149 ± 2

Darran complex

diorite pegmatite

K-Ar

hornblende

153 ± 8

Darran complex

hbl dioritic pegmatite hbl dioritic pegmatite

K-Ar

hornblende

152.6 ± 2.1

K-Ar

hornblende

156.4 + 4.1

Darran complex

Source Nathan et al. (2000) Kimbrough et al. (1994) Kimbrough etal. (1994) Nathan et al (2000) Nathan et al. (2000) Adams (1975) Nathan et al. (2000) Nathan et al. (2000) Nathan et al. (2000) Nathan et al. (2000) Adams (1975) Nathan et al. (2000)

*Data are organized according to stages discussed in the text (from youngest to oldest) Additional unpublished ages cited in the text. f ARC, Arthur River complex; WFO, Western Fiordland Orthogneiss; MTZ, Median Tectonic Zone. *hbl, hornblende, § SID, standard isotope dilution method. "C, crystallization age; M, metamorphic age. " Sample location indicated on Figure 2.

Interpretation" C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C C

Early phase of arc Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

C

Early phase of arc

MID-LOWER CRUSTAL ATTACHMENT ZONE

1989). The crystallization age of the Western Fiordland Orthogneiss and K-Ar cooling ages on hornblende from the upper amphibolite facies fabrics indicate a Mid-Cretaceous (~ 108 Ma) age for this deformation (Gibson et al 1988). The Anita shear zone (Figs Ib & 2) in northern Fiordland also formed during or after this period (Hill 1995; Klepeis et al 1999). Both of the Doubtful Sound and Anita shear zones record cooling, decompression and exhumation of the granulite belt following crustal thickening and arc-related magmatism (Gibson & Ireland 1995; Klepeis et al 1999). Crustal structure and geochronology In this section we define the boundaries and main lithological divisions of Fiordland's highpressure metamorphic belt (7-16kbar, Fig. Ib, Table 1). These rocks represent the deformed and metamorphosed roots of the Early Mesozoic arc. We also review the ages of the major rock units that make up the section (Fig. 2, Table 2) and describe structural relationships (Figs 3 & 4) that reflect a heterogeneous history of Early Cretaceous magmatism and convergence. Boundaries of the high-grade metamorphic belt Rocks of contrasting tectonic affinity surround Fiordland's high-pressure granulite and upper amphibolite facies gneisses. To the east of the high-grade rocks, the Darran Suite is composed of mostly unmetamorphosed gabbro and diorite that represent the upper crustal part of the Early Mesozoic arc (the MTZ). Most of the Darran Suite is only weakly deformed. However, its western margin is highly deformed and shares a steep, upper amphibolite facies foliation that forms the dominant structural grain north and NE of Milford Sound (Figs 3b & 4). This foliation represents part of a steep, 10-15 km wide shear zone (Figs 3b & 4) defined and discussed in detail later in this chapter. SW of the Darran Suite, weakly foliated tonalite and quartz diorite of the Roxburgh Suite (Fig. 2) form the boundary between the high-grade gneisses to the west and low-grade plutonic rocks and Tertiary sedimentary rocks of the MTZ to the east. The Roxburgh Suite is thought to be mostly Palaeozoic in age (Turnbull 2000). The western boundary of the high-grade gneisses coincides with the 4-5 km wide, nearvertical Anita shear zone (Fig. 2). The Anita shear zone separates intensely deformed granulite and upper amphibolite facies rocks of the

207

Arthur River complex from weakly deformed Lower Palaeozoic metasedimentary rocks and granite of the Gondwana margin. The Greenland Group, the Thurso Gneiss and the Saint Anne Gneiss (Fig. 2) represent parts of this Lower Palaeozoic succession. Upper amphibolite and greenschist facies mineral assemblages within the Anita shear zone record peak metamorphic pressures of — 12kbar and decompression that accompanied exhumation of the Fiordland granulites after -108-105 Ma (Hill 1995; Klepeis et al 1999). The Alpine fault and other Late Cenozoic faults (Fig. 2) reactivate the steep margins of the Anita shear zone and truncate the exposures of high-grade gneisses north of Milford Sound (Claypool et al 2002). The southern boundary of the metamorphic belt occurs south of Doubtful Sound (Fig. Ib) and coincides with a large Tertiary fault that forms the northern margin of the Cambrian-Permian Southwest Fiordland terrane (Fig. Ib). Lithological divisions of the magmatic arc West of the Darran Suite near Milford Sound, the Arthur River complex (ARC, Fig. 2) is composed of gabbroic and dioritic gneiss that is subdivided into four lithologically distinctive units (Fig. 2). The Milford Gneiss is composed mostly of garnet- and hornblende-bearing metagabbro with minor migmatitic metadiorite (m, Fig. 2). Mafic dykes of the Milford Gneiss intrude rafts of Palaeozoic rock at Camp Oven Creek south of Milford Sound (CO, Fig. 2). The Mount Edgar Diorite (Fig. 2) is a pile of dioritic sheets that intruded into the Milford Gneiss, cutting across older gneissic layering at low angles. Structural relationships within these units are described in more detail in the next section. North of Milford Sound, the Pembroke Granulite forms a large pod that is aligned parallel to the NNE strike of the dominant foliation near Milford Sound (Fig. 4). This unit includes two pyroxene- and hornblende-bearing metagabbros and migmatitic metadiorite. East of the Pembroke Granulite, the Harrison Gneiss is composed of banded metadiorite with discontinuous sheets of mafic to felsic compositions in gradational contact with the more mafic Milford Gneiss. The zone located between the Arthur River complex and the Darran Suite is composed of discontinuous gabbroic units intruded by variably deformed quartz diorite gneiss and felsic dykes. Bradshaw (1990) divided this zone into the Indecision Creek and Mount Anau igneous complexes. For simplicity, we refer to all of the rocks in this area as the Indecision Creek complex (Fig. 2). The northern contact of the

208

K. A. KLEPEIS ETAL.

Fig. 3. (a) Composite cross-section constructed using structural data from Caswell, Charles and George Sounds (locations shown in Fig. 2). Profile shows the geometry of structures above and below the upper contact of the Western Fiordland Orthogneiss batholith. Shaded areas represent Palaeozoic paragneiss representative of the ancient Gondwana margin. Dashed lines near C-C represent gneissic layering that predated batholith emplacement. (b) Composite cross-section constructed using data from the region between the Pembroke Valley and Mount Daniel (for locations see Fig. 2). Sections show the geometry of structures above and below the lowermost contact of the Western Fiordland Orthogneiss. Patterns are the same as in Figure 2.

Indecision Creek complex is gradational with the Harrison Gneiss; its eastern and western contacts are gradational with dioritic rocks of the weakly deformed Darran Suite and the highly deformed Western Fiordland Orthogneiss, respectively. This unit is chemically similar to the Western Fiordland Orthogneiss (Fig. 2) and may contain less Palaeozoic inheritance than the subunits of the Arthur River complex (Turnbull 2000). In the central part of the high-grade metamorphic belt, the Western Fiordland Orthogneiss is a major batholith that intruded the Arthur

River complex. The batholith is composed mostly of layered sequences of gabbro and diorite with some ultramafic dykes and pods. The lower (northern) contact of this unit is well exposed at Mount Daniel (MD, Fig. 2); its upper (southern) contact with Palaeozoic metasedimentary rocks is well exposed at Caswell Sound. Some areas of this batholith preserve primary igneous layering that escaped the intense recrystallization observed in most of the rock units of northern Fiordland (described below). However, in areas such as the eastern end of George Sound and along its eastern

MID-LOWER CRUSTAL ATTACHMENT ZONE

209

Fig. 4. (a) Map showing structural data from northern Fiordland. Structural data are from this study and data from Bradshaw (1989, 1990), Blattner (1991), Klepeis et al (1999), Daczko et al (2002) and Claypool et al (2002). Foliation trajectories represent the interpolation of structural trends constructed using plotted data and reconnaissance mapping. Shaded areas show the Indecision Creek shear zone (eastern side) and the George Sound shear zone (western side), Lower hemisphere. Equal-area stereoplots in (b), (c), (d) and (e) show poles to foliations (open circles) and mineral lineations (black squares) for the four structural domains discussed in the text. Black lines in stereoplots represent best-fit planes to foliation clusters.

210

K. A. KLEPEIS ETAL.

contact, these rocks were intensely deformed and recrystallized. Rafts of Palaeozoic metasedimentary rocks occur within and near the contacts of the intrusion east of Mount Daniel and at George Sound. These metasedimentary rocks are migmatitic (m, Fig. 2) near the contacts with the batholith.

Ages of magmatism, crustal melting and high-grade metamorphism Published U-Pb ages from northern Fiordland (Fig. 2, Table 2) reflect a history of Early Cretaceous magmatism and Early Cretaceous metamorphism and melting. Tulloch et al (2000) analysed zircon cores from the Arthur River complex near Milford Sound and obtained both Early Carboniferous (—360 Ma) and Early Cretaceous (— 134 Ma) U-Pb ages. Hollis et al (2002) obtained 136-129 Ma U-Pb ages from the cores of magmatic zircon in the Arthur River complex and showed that these ages most likely reflect the crystallization of Early Cretaceous dykes and plutons. These data, combined with evidence of an intrusive relationship between the mafic dykes of the Milford Gneiss and migmatitic Palaeozoic host rocks at Camp Oven Creek (CO, Fig. 2), suggest that the Arthur River complex was emplaced into crust composed of a mixture of MTZ arc rocks and Palaeozoic rocks of Gondwana. A crystallization age of — 154 Ma for the main dioritic part of the Selwyn Creek Gneiss (Fig. 2, Table 2) suggests that these rocks also were emplaced during an early phase of MTZ magmatism (Hollis et al. 2002) and likely represent deformed parts of the Darran Suite. Following emplacement of the Arthur River complex, the Mount Edgar Diorite (Fig. 2) was emplaced at - 129 Ma (Hollis et al 2002). This age is consistent with field relationships indicating that the Mount Edgar Diorite intruded and cuts across gneissic layering in the Milford Gneiss at low angles. The dioritic parts of the Indecision Creek complex also intruded older metagabbroic crust. A few unpublished U-Pb dates suggest that parts of it intruded between — 135 and — 122 Ma (reported in Bradshaw 1985, 1990). These ages, combined with a history of some Palaeozoic inheritance (although less than the Arthur River complex), indicate that it was emplaced following the amalgamation of the MTZ with the Gondwana margin. This relationship suggests that it probably has a similar origin as the Mount Edgar Diorite and the Western Fiordland Orthogneiss.

U-Pb geochronological data also indicate that high-grade metamorphism accompanied Early Cretaceous magmatism and crustal melting. Thin, low-U metamorphic rims around magmatic zircon cores of both Palaeozoic and Early Cretaceous occur in the Milford Gneiss, the Harrison Gneiss, the Selwyn Creek Gneiss and Palaeozoic gneiss at Camp Oven Creek (Tulloch et al 2000; Hollis et al 2002; Klepeis et al 2004). These ages are interpreted to reflect the influx of heat that accompanied the emplacement of the mafic-intermediate Western Fiordland Orthogneiss (Tulloch et al 2000; Hollis et al 2002). This interpretation is consistent with the range of crystallization ages from the Western Fiordland Orthogneiss (126-120 Ma, Table 2) and with the spatial distribution of zircon displaying metamorphic rims in migmatite and granulite facies rocks below the batholith. The age of lower crustal melting (migmatite localities shown in Fig. 2) also is interpreted to have coincided with this widespread thermal pulse. Zircon ages (81.8 ± 1.8 Ma) from a posttectonic dyke located in the Pembroke Valley (P, Fig. 2) north of Milford Sound indicate that high-grade granulite and upper amphibolite facies metamorphism in the Arthur River complex terminated in the Mid-Cretaceous (Hollis et al 2002). In addition, K-Ar ages on hornblende (Gibson et al 1988; Nathan et al 2000) and U-Pb dates on apatite (Mattinson et al 1986) indicate that the Western Fiordland Orthogneiss and Arthur River complex had cooled to 300-400C by -90 Ma (Table 2). KAr amphibole and biotite cooling ages of —93 and —77 Ma, respectively, also support a MidCretaceous age for the extensional Doubtful Sound shear zone (Gibson et al 1988). Structural relationships Two distinctive structural domains characterize the zone between the Anita shear zone and the Darran Suite (Fig. 4a-c). The western part of this zone is dominated by layered intrusions (e.g. the Mount Edgar Diorite) and gneissic foliations that mostly dip moderately to the south, SW and west (Fig. 4b). Hornblende and biotite mineral lineations on foliation planes plunge moderately to the west and SW. This western domain is narrowest (>10km) near Milford Sound and widest (>30km) near Caswell Sound. The eastern domain contains penetrative upper amphibolite facies foliations that are mostly steep to subvertical and strike to the NNE (Figs 3b & 4c). These latter foliations

MID-LOWER CRUSTAL ATTACHMENT ZONE

define a 10-15 km wide zone of penetrative ductile deformation that we define in this chapter as a steep shear zone. Hornblende and biotite mineral lineations in this domain mostly plunge moderately to the south and SW with some localities also displaying near-vertical plunges (Fig. 4c). We introduce the new name Indecision Creek shear zone for this zone of intense deformation. All foliations and lithological contacts in both domains are cut by mylonitic foliations of the Anita shear zone (Fig. 4a). The structure of the western domain constitutes a tilted middle and lower crustal section, with the deepest palaeodepths occurring in the north near Milford Sound and the shallowest palaeodepths occurring in the south near Caswell and George Sounds (Fig. Ib, Table 1). The dominant features in this section include the diorite intrusions that make up the Mount Edgar Diorite and the Western Fiordland Orthogneiss (Figs 2 & 3c). At Mount Daniel (MD, Fig. 2) the lower contact of the Western Fiordland Orthogneiss is mostly concordant with gneissic foliations in the Milford Gneiss. This relationship suggests that magma exploited gneissic layering in the older rocks during batholith emplacement. Elsewhere, such as at the northwestern end of George Sound, the contacts of the Western Fiordland Orthogneiss cut obliquely across gneissic layering in host rocks (Fig. 3a). In some areas, rocks of the Western Fiordland Orthogneiss (WFO) and Mount Edgar Diorite record a heterogeneous, tectonic overprint. The eastern side of the Mount Edgar Diorite is folded (Fig. 2). At Mount Daniel the lower contact of the WFO dips to the west and SW and also is folded (Fig. 3b). However, most of the original intrusive relationships between different phases of the batholith are well preserved at this locality. In areas where this deformation is intense, such as at the SE end of George Sound and near the upper and eastern contacts of the Western Fiordland Orthogneiss, these rocks contain a strong upper amphibolite facies tectonic foliation that transposes all primary igneous layering in the batholith. Thermobarometric data (Table 1) derived from garnet granulite and upper amphibolite facies mineral assemblages from the Western Fiordland Orthogneiss and its contact aureoles constrain the range of palaeodepths represented in the tilted section. Pressures reflecting the peak of Early Cretaceous metamorphism range from 7-9kbar at Caswell and George Sounds, to 10-13kbar near the base of the Western Fiordland Orthogneiss, and 13-16 kbar in the deepest part of the section north of Milford Sound (Fig. la). These

211

data indicate palaeodepths of 25-50 km from south to north (Fig. 3a, b). The dominant NNE-striking foliations of the eastern domain everywhere cut and transpose the west- and SW-dipping foliations and igneous layering of the western domain. Crosscutting relationships between these two groups of foliations are best preserved inside a ~15 km wide transitional zone that includes the eastern contacts of the Milford Gneiss (from east of Mount Edgar, ME, to the Pembroke Valley, P, Fig. 2) and the Western Fiordland Orthogneiss. In this transitional zone the dominant SW- and west-dipping foliations of the western domain are tightly folded into a series of overturned south-plunging folds (Figs 3b & 4a). Fold tightness (interlimb angle) increases from NW to SE across this zone and fold axial planes progressively steepen toward the SE. Where these folds are tight to isoclinal and upright, such as occurs inside the Indecision Creek complex, the NNE-striking foliation of the eastern domain parallels the axial planes of the folds. In zones of the highest strain, the steep NNE-striking foliations transpose all folds and intrusive contacts, including the eastern contacts of the Western Fiordland Orthogneiss and Mount Edgar Diorite. Mylonitic foliations also occur locally. These crosscutting relationships and variations in the intensity of folding and transposition form the basis of our interpretation that this eastern domain represents a zone of high strain. The kinematic significance of this zone is discussed in a later section (Stage 3). Between Caswell Sound and Mount Daniel, a third structural domain occurs near the southern end of George Sound (Fig. 4a). This domain exhibits structural relationships that are similar to those of the eastern domain. At George Sound, steep upper amphibolite facies foliations that strike to the NNE transpose older west- and SW-dipping gneissic foliations and dykes (Fig. 4d) inside the Western Fiordland Orthogneiss. These steep foliations parallel the axial planes of tight, upright folds of the dykes and the older gneiss. The boundaries of this zone approximately parallel those of the eastern domain (Fig. 4a). These relationships indicate that the structures of this third area, like those of the eastern domain, are younger than the dominant gneissic foliations of the titled section in western Fiordland. The tightness of folding and degree of transposition also indicate that this area represents a zone of high strain. We introduce the name George Sound shear zone to describe this zone of high strain. Finally, the fourth structural domain occurs at Caswell Sound (Fig. 4a, e). Here, a series of

212

K. A. KLEPEIS ETAL.

upper amphibolite and garnet granulite facies thrust zones cut and transpose the upper contact of the Western Fiordland Orthogneiss. The zone consists of both west- and east-verging ductile thrusts that separate a central domain that contains open to tight folds of gneissic foliation in Palaeozoic metasedimentary rocks (Fig. 3a). Fold tightness increases towards the contacts of the Western Fiordland Orthogneiss. The most intense zone of thrusts occurs at the eastern end of the sound. Here, thrust splays and fold axial surfaces curve downward to sole into a 1 km thick basal shear zone that occurs within the Western Fiordland Orthogneiss (Fig. 4e). Hornblende, clinozoisite and biotite mineral lineations on foliation surfaces in the dipping thrust zones plunge moderately and gently to the west (Fig. 4e). This basal shear zone is exposed at Charles Sound and at the eastern end of Caswell Sound. The steep upper amphibolite facies fabric of the George Sound shear zone merges into parallelism with this gently dipping basal shear zone (Figs 3a & 4a).

Space-time correlation of high-grade fabrics Crosscutting relationships, published U-Pb ages and similarities in metamorphic grade and structural style allowed us to correlate magmatic, metamorphic and deformational events across northern Fiordland. These correlations helped us to reconstruct the Cretaceous structural and magmatic evolution of the middle and lower crustal section. A key marker horizon in the section is the 126-120 Ma Western Fiordland Orthogneiss. The well-documented age and regional extent of this unit allowed us to divide the evolution of rock fabrics in the section into three time intervals corresponding to events that occurred before (>126Ma), during (126-120 Ma) and after (500 m) they contain chlorite-epidote amphibolite facies assemblages. Daczko et al. (2002) described these mineral assemblages in detail

MID-LOWER CRUSTAL ATTACHMENT ZONE

and showed that they reflect a temperature gradient of 700-800 °C within the contact aureole and 550-600 °C outside it. These observations suggest that the thrust zones initially formed within the thermally softened aureole of the batholith and continued to evolve as batholith cooled. The high pressures (7-9 kbar, Fig. Ib) and temperatures recorded by the mineral assemblages of the Caswell thrust zone and their contractional style indicate that they formed prior to the onset of regional extension at ~ 108 -105 Ma (Daczko et al 2002). Using this relationship and the fact that the thrust zone deforms the margins of the Western Fiordland Orthogneiss, the Caswell thrust zone must have evolved during the same time interval (116-105 Ma) as the George Sound shear zone and the Indecision Creek complex shear zones (ages from Klepeis et al. 2004). These features exhibit identical metamorphic grade and identical crosscutting relationships with respect to the Western Fiordland Orthogneiss. This interpretation also is consistent with the gradual merging of the George Sound shear zone with the basal shear zone of the Caswell thrust zone. The structural relationships we have outlined show a remarkable variability with depth in the crustal section (Fig. 3a, b). These relationships combined with good age control at the regional scale allowed us to investigate variations in the style of strain partitioning with depth for different time periods. Here, we define three stages in the magmatic and structural evolution of the Fiordland section: 1

the injection of mafic-intermediate magma of the Western Fiordland Orthogneiss batholith into a lower crust during the interval 126120 Ma, the partial melting of mafic host rocks below the batholith and its aureole (m, Fig. 2) accompanying this magmatism; 2 the mobilization and extraction of lower crustal melts during deformation that occurred as the batholith cooled to subsolidus temperatures (120-116 Ma); 3 the development of the steep upper amphibolite facies fabrics that define the Indecision Creek and George Sound shear zones and the Caswell thrust zone (116-105 Ma). The second and third stages are bracketed by the age of the Western Fiordland Orthogneiss and the youngest age of lower crustal fabrics that formed prior to the Mid-Cretaceous exhumation of the belt (~ 108-105 Ma). These age intervals allow for diachronous activity within the belt.

213

Stage 1: Mafic-intermediate magmatism and the partial melting of lower crust During the interval 126-120 Ma the Western Fiordland Orthogneiss was emplaced into lower crust composed of older (>126Ma) rocks of the Median Tectonic Zone and Palaeozoic margin of Gondwana. This magmatism resulted in a > 3000 km2 batholith that was at least 10km thick. The first phases of the batholith were gabbroic with minor ultramafic compositions; later phases were dominated by coarsegrained diorite. In some places, diffuse, undulate contacts between the gabbro and slightly younger diorite dykes suggest that these two phases mingled while still in a semi-molten state (Fig. 5a). The lower contact of the batholith at Mount Daniel (MD, Fig. 2) preserves a 200-500 m thick banded igneous complex. In this zone, sheets of tonalite, trondhjemite and hornblende gabbro display mutually crosscutting relationships and are complexly interfolded (Fig. 5b). Thin ( 750 °C. Piston-cylinder experiments conducted at the University of Vermont (Antignano et al. 2001; Antignano 2002) on a natural unmelted sample

215

of metadiorite from the Pembroke Granulite confirmed this relationship. These experiments produced two results that are of importance to the study of partial melting in the Arthur River complex during batholith emplacement. First, in accordance with field observations, they showed that at T = 850 °C biotite undergoes melting in the absence of free water followed by the reaction of hornblende and clinozoisite to form garnet plus melt as reaction products (Antignano et al 2001; Antignano 2002). Second, they showed that the melt fractions produced during melting remained low (10km thick batholith composed of mafic-intermediate magma was emplaced into the lower crust section. During magma emplacement, layer-parallel, melt-enhanced shear zones formed at the upper and lower boundaries of the batholith where strength contrasts were produced by magma and steep temperature gradients. Magmatism was accompanied by granulite facies metamorphism and the formation of migmatite up to 10km below the intruding batholith. Fluid-absent melting of mafic-intermediate gneiss was controlled mostly by the decomposition of hornblende + clinozoisite. Positive volume changes and high melt fluid pressures during melt production resulted in fracture networks that aided the segregation and transfer of melt through and out of the lower crust during the period 120116 Ma. Ductile deformation in steep shear zones also was important for moving partial melts vertically and horizontally through the lower crust. During the period 116-105 Ma, the batholith and rocks located below it cooled to temperatures of T < 800 °C, and differential shortening of the crust created a heterogeneous network of subvertical and subhorizontal shear zones. During the earliest stages of this deformation, these shear zones exploited the layered architecture of the middle and lower crust and separated it into structural domains exhibiting widely different structural styles. A ductile fold-and-thrust system formed in the middle crust (25-30 km palaeodepths), with garnet granulite and upper amphibolite facies thrust splays rooting down into a 1 km thick, nearly horizontal shear zone that localized at the uppermost contact of the batholith. This flat shear zone physically connected the mid-crustal fold-and-thrust zone to subvertical shear zones in the lower crust that are up to 15 km wide. The subvertical foliations in these steep shear zones cut across the lower contact of the batholith and merge smoothly with the subhorizontal shear zone in the middle crust. Fabric correlations based on regionalscale crosscutting relationships and U-Pb data strongly suggest that these systems evolved simultaneously or nearly so during the same 116-105 Ma interval. Simultaneous movement

above and below the mid-crustal fold-andthrust zone also is required to accommodate shortening at different crustal depths. This crustal structure and evidence of kinematic links above and below the flat shear zone at the top of the batholith define a midlower crustal attachment zone that accommodated differential displacements at different crustal depths. Together, this network of flat and steep shear zones accommodated mostly subhorizontal east-west, arc-normal shortening in the fold-and-thrust zone at mid-crustal depths and oblique sinistral displacements on steep, complexly deforming shear zones at lower crustal depths. The oblique-sinistral displacements also were accompanied by a component of vertical thickening and subhorizontal arc-parallel displacements on NNE-striking surfaces. These displacement patterns are consistent with oblique convergence along an Early Cretacous plate boundary located outboard of the early Mesozoic arc. We are grateful to Cheryl Waters, Olivier Vanderhaeghe and Basil Tikoff for suggestions on ways to improve the original manuscript. We also thank N. Daczko, I. Turnbull, N. Mortimer and A. Tulloch for helpful discussions during this study. J. Hollis, A. Claypool, S. Marcotte, W.C. Simonson and G. Mora-Klepeis also provided valuable assistance. We thank the Department of Land Conservation in Te Anau for permission to visit and sample localities mentioned in the text. This study was funded by a National Science Foundation grant to KAK (EAR0087323), an Australian Research Council grant to KAK and GLC (ARC-A10009053), and grants from the Geological Society of America and the University of Vermont.

References ADAMS, C.J. 1975. New Zealand potassium argon age list, 2. New Zealand Journal of Geology and Geophysics, 18, 443-467. ANTIGNANO, A., IV, 2002. Experimental constraints on granitoid compositions in convergent regimes: a geochemical study. MSc Thesis, University of Vermont. ANTIGNANO, A., IV, RUSHMER, T, DACZKO, N.R., CLARK, G.L., COLLINS, WJ. & KLEPEIS, K.A. 2001. Partial melting of a hornblende-biotiteclinozoisite bearing metadiorite: applications to the deep crust, Fiordland, New Zealand. Geological Society of America Abstracts with Programs, 33(6), 211. AXEN, G.J., SELVERSTONE, J, BYRNE, T & FLETCHER, J.M. 1998. If the strong crust leads, will the weak crust follow? GSA Today, 8, 1-8. BARKER, AJ. 1990. Introduction to Metamorphic Textures and Micro structures. Chapman and Hall, New York. BEAUMONT, C, JAMIESON, R.A., Nguyen, M.H. & LEE, B. 2001. Himalayan tectonics explained by

MID-LOWER CRUSTAL ATTACHMENT ZONE extrusion of a low-viscosity crustal channel coupled to focused surface denudation. Nature, 414, 738-742. BLATTNER, P. 1976. Replacement of hornblende by garnet in granulite facies assemblages near Milford Sound, New Zealand. Contributions to Mineralogy and Petrology, 55, 181-190. BLATTNER, P. 1978. Geology of the crystalline basement between Milford Sound and the Hollyford Valley, New Zealand. New Zealand Journal of Geology and Geophysics, 21, 33-47. BLATTNER, P. 1991. The North Fiordland transcurrent convergence. New Zealand Journal of Geology and Geophysics, 34, 543-553. BRADSHAW, J.D. 1989. Cretaceous geotectonic patterns in the New Zealand region. Tectonics, 8, 803-820. BRADSHAW, J.Y. 1985. Geology of the northern Franklin Mountains, northern Fiordland, New Zealand, with emphasis on the origin and evolution of Fiordland granulites. Unpublished PhD Thesis, University of Otago, Dunedin, New Zealand. BRADSHAW, J.Y. 1989. Origin and metamorphic history of an Early Cretaceous polybaric granulite terrain, Fiordland, southwest New Zealand, Contributions to Mineralogy and Petrology, 103, 346-360. BRADSHAW, J.Y. 1990. Geology of crystalline rocks of northern Fiordland: details of the granulite facies Western Fiordland Orthogneiss and associated rock units. New Zealand Journal of Geology and Geophysics, 33, 465-484. BRADSHAW, J.Y. & KIMBROUGH, D.L. 1989. Comment: Age constraints on metamorphism and the development of a metamorphic core complex in Fiordland, southern New Zealand. Geology, 17, 380-381. CLARKE, G.L., KLEPEIS, K.A. & DACZKO, N.R. 2000. Cretaceous high-P granulites at Milford Sound, New Zealand: metamorphic history and emplacement in a convergent margin setting. Journal of Metamorphic Geology, 18, 359-374. CLAYPOOL, A., KLEPEIS, K., DOCKRILL, B., CLARKE, G., ZWINGMANN, H. & Tulloch, A. 2002. Structure and kinematics of oblique continental convergence in northern Fiordland, New Zealand. Technophysics, 359, 329-358. CLEMENS, J.D. & MAWER, C.K. 1992. Granitic magma transport by fracture propagation. Tectonophysics, 204, 339-360. CONNOLLY, J.A.D., HOLNESS, M.B., RUBIE, D.C. & RUSHMER, T. 1997. Reaction-induced micro-cracking: an experimental investigation of a mechanism for enhancing anatectic melt extraction. Geology, 25,591-594. DACZKO, N.R. 2001. The structural and metamorphic evolution of Cretaceous high-P granulites, Fiordland, New Zealand. Unpublished PhD Thesis, University of Sydney, Australia. DACZKO, N.R., KLEPEIS, K.A. & CLARKE, G.L. 2001 1850 Ma) on the continental margin was followed by an extended sinistral transpression from 1850 to 1730 Ma now separated into three episodes or peaks of activity. The first episode was focused on the back-arc region and was followed by the main arc construction phase during which transpression was partitioned into strike-slip and contraction components. Despite the longevity of this active margin system, individual tectonic events took place rapidly, e.g. development of fore-arc Dj-D 3 and accompanying high-temperature, low-pressure metamorphism took place over c. 12 Ma. We explain the fore-arc and batholith evolution by the upward migration of an underlying attachment structure through the upper crustal partitioned blocks. This migration may be attributed to an increase in the geothermal gradient accompanied by, or followed by, exhumation of the mid-crust. The partially molten, hence weak, attachment zone solidified and strengthened during cooling before emplacement of the post-orogenic rapakivi suite during the third distinct phase of mild sinistral transpression.

Oblique plate convergence forms orogenic belts that display complex three-dimensional transpressional architectures as shown by many field studies, predicted by analytical descriptions and exhibited by numerical and analogue experiments (Dewey 1975; Mackenzie & Jackson 1983; Richard & Cobbold 1990; Teyssier & Tikoff 1998). In the upper crust, partitioning of transpression produces coexisting orogennormal shortening and orogen-parallel strikeslip deformation (Fitch 1972; McCaffrey 1992; Molnar 1992; Tikoff & Teyssier 1994; Dewey et al 1998). The nature of this deformation partitioning is likely to change with depth due to the differing rheologies exhibited by the upper and lower crust, and the lithospheric mantle (Molnar 1992; Kohlstedt et al 1995; Roy den 1996; Mackenzie & Jackson 2002). The mid- to lower crust is typically composed of amphibolite to granulite facies rocks. In many exhumed orogenic belts these rocks display originally subhorizontal peak metamorphic fabrics

and evidence for high shear strains, e.g. sheath folds (Strachan et al. 1992; Northup & Burchfiel 1996). These shear zones have been termed 'basal detachments' or 'decollements' (Richard & Cobbold 1990; Oldow et al. 1990; Teyssier & Tikoff 1998), with implied decoupling between the upper crust and the mantle lithosphere. Recently, Tikoff etal. (2001) and Teyssier et al. (2002) have proposed the term 'attachment' for a subhorizontal shear zone across which there is coupling between lithospheric layers that display different rheological properties. Tikoff et al (2001) further suggest that these attachment zones are important for transmitting driving forces vertically within an orogenic belt and question the importance of horizontal or side-acting tectonic driving forces in orogenic belts. Attachment zones should be recognizable by 3D architectures and kinematic frameworks that are compatible with an obliquely convergent orogen, taking into account contrasting rheologies

From: GROCOTT, JL, MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) 2004. Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227, 231-248. 0305-8719/04/$15 © The Geological Society of London 2004.

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typical of different lithospheric levels. In this chapter, we examine an obliquely convergent orogenic belt, exposed at mid- to lower crustal levels, in order to assess whether its overall architecture is compatible with that of a partitioned transpressional orogen with an attachment structure. The Ketilidian orogen is dramatically exposed in the alpine topography of south Greenland, and a range of metamorphic grades suggest that different crustal levels are exposed. We provide a summary of the architecture and tectonic history of the Ketilidian orogen (Fig. 1), a type example of Palaeoproterozoic (c. 1800 Ma) continental growth by arc magmatism. It formed in response to convergence between an oceanic plate and an Archaean craton (Chadwick & Garde 1996; Garde et al 20020). Unlike many other Palaeoproterozoic orogens, primary structure is well preserved because the orogen escaped modification by collisional processes. The evolution of the Ketilidian orogen has been outlined elsewhere (e.g. Garde et al. 20020, b), but here the aim is to focus on the overall tectonic architecture of the orogen and to discuss the nature and timing of the main tectonothermal events that affected its component parts in the context of a preserved attachment structure.

(Berthelsen & Henriksen 1975, and references therein) (Fig. 1). In the northwestern Border Zone (NWBZ), supracrustal rocks in Midternaes and Grasnseland collectively form a minimum 5 km thick stratigraphic sequence that overlies peneplained Archaean basement and includes siliciclastic and calcareous rocks with chertpodded iron formation and overlying c. 4 km thick Sortis Group basic metavolcanic rocks. In the northeastern part of the orogen, the northeastern Border Zone (NEBZ) is exposed between Mogens Heinesen Fjord and Napasorsuaq Fjord (Garde et al. 1999) (Fig. 1). Isolated occurrences of low metamorphic grade conglomerates, feldspathic and quartz arenites assumed to have been deposited unconformably on an Archaean basement are exposed on nunataks west and NW of the peninsula Puisortoq (Fig. 1). South of Puisortoq, on either side of Tunua sound (Fig. 1), a 5 km thick supracrustal sequence metamorphosed to upper amphibolite facies is present. Psammites, meta-andesites, minor metamorphosed conglomerates, siltstones, mudstones, calcareous rocks and thin horizons of amphibole- and diopside-bearing intermediate metavolcanic rocks are tentatively correlated with the basal part of the Vallen Group of the NWBZ. Equivalents of the Sortis Group basic metavolcanic rocks are absent. The Ketilidian orogen - major Garde et al. (20020) reported a preliminary ion probe study of zircons from supracrustal rocks of components the NEBZ. The results showed a complex distriThe core of the orogen (Fig. 1) is a 100-200 km bution with a maximum age of c. 1850 Ma resultwide, NE-trending continental magmatic arc, the ing from (i) detrital zircon cores derived from the Julianehab batholith (JB), constructed by juvenile Julianehab batholith, (ii) recrystallization and calc-alkaline magmas emplaced during north- overgrowth at c. 1800-1780 Ma and (iii) variable directed oblique subduction of an oceanic plate lead loss during a subsequent heating episode. (Chadwick & Garde 1996; Garde et al. 20020). The Julianehab batholith is separated from the Archaean foreland by the c. 50 km wide Border Julianehab batholith Zone exposed in the northwestern and northeastern The Julianehab batholith, exposed over a parts of the orogen; however, these two regions are c. 30 000 km2 wedge-shaped area of south Greenseparated by the Greenland Ice Cap (Fig. 1). To the land (Fig. 1), was emplaced between 1850 and SE beyond the outboard margin of the batholith, 1800 Ma (Hamilton et al. 1996; Hamilton 1997; the Psammite and Pelite Zones have been collec- Garde et al. 20020) and is clearly younger than tively interpreted as the relics of the fore-arc the Border Zone supracrustal sequence that it basin (Chadwick & Garde 1996; Garde et al. intrudes. Intrusive ages for the entire batholith 2002a). A prominent late- to post-orogenic suite cluster into two groups centred on c. 1840 and of rapakivi granites sensu Into and related noritic c. 1810 Ma (Garde et al. 20020). Granodiorite plutons are also present in this part of the orogen. and granite sensu stricto predominate along with quartz monzodiorite, tonalite and quartz syenite. Dioritic and more basic lithologies comprise at Border Zone least 5% of the batholith, and form kilometreThe Border Zone (Fig. 1) is defined as the scale distinct intrusions, numerous swarms of synregion of reworked Archaean basement, plutonic mafic dykes and centimetre- to decideformed and recrystallized Palaeoproterozoic metre-sized, fine- to medium-grained dioritic basic Iggavik dykes (c. 2130 Ma) and the enclaves in leucocratic plutons (Chadwick et al. overlying Palaeoproterozoic supracrustal rocks 1994; Chadwick & Garde 1996). Biotite and horn-

Fig. 1. Geological map of the Ketilidian orogen in south Greenland (modified from Garde et al 2002(3). The inset map of Greenland shows the location of the Ketilidian orogen (Ket), the Archaean block of southern Greenland (AB), and the Ammassalik (Am) and Nagssugtoqidian (Nag) orogenic belts. Locations of cross-sections A-A', B-B', C-C' and D-D'-D" are indicated on main map. A, Archaean; S, supracrustal rocks exposed in nunataks.

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blende are the main ferromagnesian minerals, but ortho- and clinopyroxene and locally olivine occur in the mafic intrusions. The batholith has a distinct calc-alkaline geochemical character (strong negative Nb, Ta, P and Ti anomalies), rare-earth element (REE) patterns show distinct LREE enrichment with small positive or negative Eu anomalies, and Sr, Pb and Nd isotopic data indicate its juvenile nature (Garde et al 20020). Psammite and Pelite Zones The Psammite Zone and Pelite Zone occupy a minimum width of c. 100 km and comprise deformed and metamorphosed feldspathic arenites (psammites), siltstones and mudstones (pelites), minor limestones and volcanic rocks with an estimated pre-erosional thickness of over 15 km (Garde et al. 20020). Psammitic rocks predominate in the NW close to the batholith (the Psammite Zone), whereas pelitic rocks are widespread in the SE (the Pelite Zone). An excellently preserved volcano-sedimentary sequence was deposited unconformably on the Julianehab batholith at Kangerluluk (Mueller et al. 2000). The Psammite Zone around N0rrearm also contains broadly concordant sheets of hornblende granodiorite, (e.g. sample 412056 dated at 1792 + 1 Ma by Hamilton 1997), diorite, gabbro and gabbro-anorthosite, which are lithologically indistinguishable from intrusive suites in the Julianehab batholith (Garde et al. 20020). The largest of these, the Stendalen complex west of N0rrearm (Fig. 1), is a c. 8 km wide and minimum 1300 m thick, saucer-shaped layered intrusion of gabbro, leucogabbro and anorthosite. Both the Psammite and Pelite Zones have been intensely deformed and have undergone metamorphism to migmatite grade (Garde et al. 20020). In the Pelite Zone, migmatized paragneiss predominates south of Lindenow Fjord; however, anatectic garnetiferous granites and components of the rapakivi granite suite underlie much of the region from Kap Ivar Huitfeldt to Kap Farvel (Fig. 1). Intrusions of the rapakivi suite collectively cover c. 3000km2 and were emplaced mainly into supracrustal rocks of the Psammite and Pelite Zones between 1755 and 1723 Ma (Fig. 1; van Breemen et al 1974; Gulson & Krogh 1975; Hamilton 1997; Garde etal 20020). Border Zone structure Border Zone-foreland boundary relationships In the northwestern Border Zone, Palaeoproterozoic reworking of the Archaean margin is shown

by the NE-trending Iggavik dyke swarm of the foreland becoming increasingly reorientated, deformed and recrystallized southwards (Berthelsen & Henriksen 1975) (Fig. 2a). Henriksen (1969) noted that north of Arsuk Brae the dykes are bent and open folded whereas to the south on the Ivittuut peninsula they are boudinaged into trains of amphibolite lenses and almost conformable to the NNW-trending gneissic foliation. The overlying basal unconformity and Ketilidian cover rock sequence also become increasingly modified by ductile deformation southwards over a distance of c. 100 km (Bondesen & Henriksen 1965). To the south of Arsuk Brae the angular nature of the unconformity is absent as the moderately SE-dipping foliation in the supracrustal rocks is parallel to that in the underlying gneissic rocks (Henriksen 1969). On the east coast, a narrow transition zone approximately 10 km wide between flat-lying, low-grade sedimentary rocks, exposed on nunataks to the north, and intensely deformed and migmatized, steeply SE-dipping supracrustal rocks located south of Puisortoq, takes the form of a NE-trending, composite synform (Fig. 2b). Border Zone structural history The large-scale architecture of the NWBZ is controlled by sinistral transpression partitioned into zones of crustal shortening (folds and thrusts) and zones of sinistral strike-slip deformation (Garde et al. 1998). This main deformation phase produced a major basement-cored, WSW-WNEtrending, anticlinal arch in the region to the west of Graenseland (Fig. 2a). A synclinal structure has also been mapped in the cover rocks at Midternaes to the north. To the south of the main arch, at Arsuk 0, an ENE-trending syncline and anticline pair have been mapped. In the area between Graenseland and Kobberminebugt, sinistral shear produced prominent ENE-plunging, S-shaped folds, which are parasitic on the southern limb of the large anticlinal arch (Fig. 2a). These formed synchronously with intrusion of the Qornoq augen granite (1848 + 2 Ma, Garde et al. 20020). At Kobberminebugt, the emplacement of the Borg Havn granite at 1845±lMa (Garde et al. 20020) post-dated an early fabric and was synchronous with sinistral non-coaxial deformation. Sinistral deformation within the Kobberminebugt shear zone was locally extremely intense and produced steeply plunging, tight folds and refolds of earlier fabrics, sinistral S-C, C' structures and asymmetric boudins. The sinistral transpressional regime was succeeded by a phase of extremely localized dextral non-coaxial ductile deformation dated by the intrusion of the 1805+ 2 Ma, Satukujoq

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Fig. 2. Geological maps of Border Zone (modified from Garde et al. 1999, 2002a). (a) Northwestern Border Zone. B, Borgs Havn granite; S, Satukujoq granite, (b) Northeastern Border Zone. Inset map shows location of more detailed maps.

granite at Kobberminebugt (McCaffrey et al 1998). Structures that predate the main sinistral transpression have been recognized locally in the Border Zone. There are north- and NNE-trending upright, gently plunging folds and NW-directed thrusts in Midternaes and Graenseland and a prominent top-to-NW transport thrust with contemporaneous tight overturned folds developed in the underlying and intercalated sedimentary rocks of the Sortis Group. The earliest recognized structures preserved in low-strain windows in the Kobberminebugt shear zone are steeply plunging folds with dextral shears on fold limbs (McCaffrey et al. 1998) (Fig. 2a). From these observations, we suggest that an early (pre-1850 Ma) phase of top-to-NW thrust-

ing and margin-parallel dextral shear predated the main sinistral transpression. In the NEBZ, migmatitic feldspathic psammites and deformed plutonic rocks related to the Julianehab batholith generally display intense, steeply SE-dipping LS fabrics with a margin-parallel (NE - SW-trending) stretching lineation and kinematic indicators that indicate sinistral shear. In western Napasorsuaq Fjord, a minimum age for the supracrustal rocks and the intense sinistral deformation is provided by a discordant, but folded, 30 cm thick granite dyke at 1800+1 Ma (Garde et al 2002^). Late, kilometre-scale NE-SW- to ENE-WSWtrending open, upright or steeply north-inclined antiforms and synforms fold this main fabric. To the east of Tunua Sound the southeastern

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limb of the late NE-trending, steeply inclined synform (Fig. 2b) displays top-to-NE kinematics. This style of deformation is similar to the main deformation observed in the batholith and forearc to the south on the east coast (see below). To summarize, the Border Zone experienced a complex tectonic history with (i) pre-1850Ma thrusting and dextral shear, (ii) c. 1845 Ma sinistral shear and plutonism (NWBZ), (iii) c. 18001780 Ma sinistral and top-to-NE shear coeval with plutonism, and high-grade metamorphism (NEBZ), and (iv) local dextral shear and plutonism (NWBZ) at c. 1800 Ma. The available evidence suggests either that the main sinistral transpressional deformation migrated away from the northwestern Border Zone towards the east or that there were two episodes of sinistral deformation, one at 1845 Ma and one at c.l 800 Ma. Julianehab batholith structure Border Zone-Julianehab batholith boundary The boundary between the Border Zone and the Julianehab batholith to its south is defined on the west coast by the ENE-trending, vertical Kobberminebugt shear zone (Fig. 2a), which comprises syn- and post-kinematic batholithrelated granitic plutons intruded into upper amphibolite facies supracrustal rocks (Watterson 1968). On the east coast, the equivalent boundary has been mapped along Napasorsuaq Fjord (Fig. 2b), where it represents the southernmost exposures of the supracrustal rocks that have been intruded by Ketilidian plutonic rocks. Here the boundary is a partially reworked intrusive contact. Julianehab batholith, west coast structure Despite being lithologically similar throughout, the Julianehab batholith displays important differences in structural geometry between the west and east coasts. Most of the western parts of the Julianehab batholith were emplaced between c. 1850 and 1800 Ma ago synchronously with deformation in an ENE-WNW-trending sinistral transpressive tectonic setting (Chadwick et al 1994; Chadwick & Garde 1996). South of Qaqortoq (Fig. 1), granitic, intermediate to mafic dykes were emplaced into a system of steep to vertical NNE- to NE-trending, sinistral and subordinate ESE-trending dextral synplutonic shear zones (Watterson 1968; Chadwick et al 1994; Chadwick & Garde 1996). Detailed contact relationships for larger plutons are

unknown except in the NW part of the batholith near Nunarrsuit (Fig. 1) where Pulvertaft (1977) has mapped steep NE-SW-trending contacts between kilometre-sized granitic bodies. A steep NE-SW-trending fabric is present throughout most of the batholith on the west coast. Mineral textures, such as the parallel alignment of subhedral, oscillatory zoned plagioclase phenocrysts set in a finer-grained matrix of quartz and feldspar with undulatory extinction indicate magmatic state deformation (Chadwick et al. 1994). These magmatic state fabrics are overprinted by coplanar LS fabrics with subhorizontal stretching lineations and sinistral S-C fabrics indicative of solid-state strain. Solid-state strain intensifies towards kilometre-scale shear zones, which are mylonitic to ultramylonitic in their central parts. The largest of these is the 1.5km wide Sardloq shear zone (Fig. 1) in the southwestern part of the batholith, which has been dated as being active from 1810 to 1820 Ma (Garde et al 20020). East coast batholith structure In contrast to the steep magmatic and tectonic fabrics that predominate in the northwestern and southwestern parts of the batholith, intrusions over large parts of the batholith in its eastern part were emplaced with moderately inclined contacts (Garde et al. 1999, 2002a). These flat-lying or gently inclined contacts along with a locally intensely developed solidstate fabric have been folded by kilometre-scale late-stage folds into moderately to steeply NWor SE-dipping panels. Unfolding the late folds restores intrusive contacts and magmatic fabrics to subhorizontal (see below). Swarms of gently inclined, undeformed intermediate to mafic silllike sheets are also present in this part of the batholith. The main solid-state foliations throughout the batholith on the east coast are defined by crystallographic and shape preferred orientations of ferromagnesian minerals and plagioclase. These fabrics are moderately to steeply dipping due to the effect of later folding (see below) but generally strike NE-SW, displaying lineations that have moderate to shallow plunges. The main kinematic indicators in the granitic rocks are cr-shaped plagioclase porphyroclasts and a crystallographic preferred orientation in quartz. These show that the earliest pervasive solidstate fabrics were mostly consistent with a transport direction of top to the NE when the effects of later upright folding are removed. At Igutsaat Fjord and Kangerluluk in the SE (Fig. 3), there is clear evidence for at least two

ATTACHMENT FORMATION DURING PARTITIONING

Fig. 3. Geological map of eastern exposures of the Ketilidian orogen showing the positions of late fold and shear zone structures. The boundaries between the Border Zone, Julianehab batholith, Psammite and Pelite Zones are indicated (long thin dashes).

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phases of pluton intrusion and solid-state deformation. This has resulted in early granodiorite orthogneiss fabrics (protoliths dated at c. 1845 Ma) being intruded by gently inclined, granitic sheets at c. 1794 Ma (Garde et al. 2002a). The late granitic sheets display geometrically similar but less intensely developed magmatic and solid-state fabrics. A series of kilometre-scale, ENE-trending open fold structures have reoriented the earlier magmatic solid-state fabrics into NW- and SEdipping panels (Fig. 3). These antiforms and synforms are spatially associated with zones of intense solid-state strain in a number of NEtrending high-strain zones. They are commonly cut by thin granite sheets or dykes indicating that the high-strain zones were developed before the magmatic activity had finally ceased. Stretching lineations within the latest highstrain zones have oblique plunges, which combined with reverse slip indicators show that these shear zones accommodated an additional component of NW-SE-directed crustal shortening. The late-tectonic Anorituup shear zone (Fig. 3) is a 5 km wide east-west-trending vertical shear zone that displays abundant kinematic indicators, e.g. 10kbar of decompression (>33km) (Fig. 1). In some terranes, decompression may be preceded or accompanied by heating. High-temperature decompression paths have been demonstrated for migmatite terranes (e.g. Jones & Brown 1990; Whitney 1992; Audren & Triboulet 1993), including migmatite-cored gneiss domes (Norlander et al. 2002), and other high-grade rocks, including some ultrahigh-pressure terranes (e.g. Su-Lu, China: Wang et al 1993; Western Gneiss Region, Norway: Dunn & Medaris 1989). P-T paths estimated for many exhumed high-grade silicic melt that has segregated on a centimetre metamorphic rocks as well as those calculated scale but has not drained far from its source has by forward modelling for the thermal evolution fuelled a long debate about the connection, or of thickened crust (England & Thompson 1984) lack thereof, between anatectic migmatites and remain within ^50-80 °C of rmax during a subgranitoids (Brown 1994, and references stantial portion of the decompression path. therein). Much attention has been given to deter- These temperatures are within the uncertainty mining small-scale melt-segregation mechan- range of most metamorphic temperature calcuisms and discussing how silicic melt, once lations (e.g. geothermometry based on cation segregated from its source, might flow, coalesce exchange), so it is not possible to discern details and pool to form larger-scale magma bodies, of the paths other than that they remain at high such as orogenic leucogranites (Le Fort et al. temperature during decompression. 1987; Inger & Harris 1993). Possible mechanisms For average values of thermal conductivity, contributing to segregation and transport include mantle heat flux, crustal heat production and other compaction (McKenzie 1984) and flow through factors related to deformation and metamordykes/fractures (Clemens & Mawer 1992), phic reactions, isothermal decompression is not

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predicted by thermal models for slowly eroding ( 10 km (and, in some cases, tens of kilometres), such as commonly observed in migmatite complexes, gneiss domes and ultrahigh-pressure metamorphic terranes, so that we can better understand thermomechanical links and feedback relationships among orogenic processes operating at the surface, at different levels within the continental crust and within the mantle. Erosion Unroofing via erosion may account for largemagnitude isothermal decompression if the erosional products are removed from the system or if there is a large lag time (>10Ma) between crustal thickening (deep burial) and the initiation of denudation. In the latter case, high-temperature decompression and partial melting may occur in part because of the added crustal heat production from buried radiogenic crust, following crustal thickening (England & Thompson 1986). Rapid erosion (>l-5 mm a"1) in concert with tectonism (uplift, faulting; Fig. 2a) may account for localized decompression, such as beneath alpine glaciers and deeply incised rivers (e.g. Zeitler et al 2001), although, in the latter case, many questions remain about how rivers or river systems evolve (move) through time and influence exhumation over a region much greater than that of the main incised channel. For regionalscale, large-magnitude decompression of deep crust, it is likely that other mechanisms are responsible for near-isothermal decompression, particularly in orogens in which there was no lag time between crustal thickening and thinning/ decompression. Low-angle normal faults Detachment systems can lead to exhumation/ decompression of deep-seated rocks by crustal

thinning during symmetric or asymmetric extension (Lister & Davis 1989). For example, asymmetric detachment systems may create a locus of lower crust upwelling that is offset relative to the normal fault break-away zone. Melting during this initial decompression is likely; however, given the low angle of the detachment, rocks cool as they are exhumed in the footwall of the detachment system, and therefore do not experience isothermal decompression (Fayon et al. 2002). Therefore, detachment faults are unlikely to account for the magnitude of decompression observed in gneiss domes and other high-pressure terranes unless the faults were originally higher angle (Buck 1988) or the rate of motion on detachments is unusually fast. Crustal thinning/collapse Regional crustal thinning may drive decompression of the deep crust during collapse of thickened crust (Rey 1993) (Fig. 2b), especially if thinning is localized in a narrow region. Thinning may involve the upper crust, lower crust, or both. Decompression by thinning of thickened crust and lateral flow of deep crust is slow if thinning is restricted to the lower crust, but can be significantly faster if the upper crust is extended, thinned and/or rapidly eroded (Teyssier & Whitney 2002). Because the rate of bulk thinning likely decays with time, this mechanism may initiate decompression, but acting alone it is unlikely to maintain near-isothermal conditions during substantial decompression. Crustal thinning combined with buoyancy may, however, produce the observed and modelled pressure-temperaturetime (P-T-t) paths. Folding /buckling Folding of the lithosphere under compression has been shown to occur in oceanic lithosphere (Wiessel et al. 1980; Gerbault 2000), and possibly also in continental lithosphere (Martinod & Davy 1994; Burg et al. 1994). Buckling theory predicts the amplification of folds of a particular wavelength, given appropriate viscosity contrasts and layer thicknesses. As an antiformal buckle develops (Fig. 2c), a positive feedback relation between folding and erosion may cause an acceleration of fold amplification, resulting in effective, localized lithospheric uplift. This principle has been applied to the exhumation of metamorphic rocks in the Namche-Barwa syntaxis of the eastern Himalayas (Burg et al 1997). The combined effect of fold amplification and erosion may allow exhumation under nearisothermal conditions, such as the P-T path

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Fig. 2. Some of the decompression mechanisms described in the text (a) erosion of uplifted rocks (may be coupled with local faulting/crustal shortening); (b) crustal thinning associated with normal faulting (also shown are schematic 400 and 700 °C isotherms and a partially molten mid-crustal zone); (c) crustal buckling (e.g. at erogenic syntaxes); (d) exhumation of subducted continental crust; and (e) upwelling of a partially molten diapir with associated downflow of denser rocks (upwelling may be driven by density inversion or triggered by removal of the upper crust or thinning of the deep crust).

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inferred for metamorphic rocks at the NamcheBarwa and Nanga-Parbat syntaxes (Burg & Podladchikov 2000). Buckling may therefore be an effective mechanism of localization of lithospheric uplift. If this lithospheric uplift is accompanied by fast glacial and/or fluvial erosion, heat will be advected upward and a positive feedback relation between crustal softening and strain localization will occur and lead to localized, near-isothermal decompression. Exhumation of ultrahigh-pressure rocks (continental subduction) Continental subduction provides a means of burying and heating a large volume of fertile material that is continually being replenished as long as subduction is active. This mechanism can produce large volumes of melt in orogens, and results in the most dramatic examples of large-magnitude decompression of continental material. Ultrahigh-pressure (UHP) terranes represent continental material that has been subducted to depths of > 100-150 km and exhumed, as shown by the presence of coesite and diamond in metamorphosed supracrustal rocks (Chopin 1984; Coleman & Wang 1995). The largest-magnitude isothermal decompression in continental orogens is recorded by UHP terranes (e.g. Zhang et al 1991 \ Nakamura & Hirajama 2000). UHP terranes now exposed at the surface typically represent crustal slices exhumed at/near the suture zone, and some are in fault contact with blueschist and serpentinite complexes (Dora-Maira, Western Alps; Dabie Shan, China). The subducted crust that returns to the surface as an identifiable UHP terrane may have remained close to the subduction zone and been exhumed via buoyancy-driven

exhumation of a slice bounded below by a thrust and above by a normal fault (Chemenda et al 1995) (Fig. 2d). Subducted continental crust that does not experience a return path near the suture may remain accreted to the base of the overriding plate as slices of continental crust and lithospheric mantle, or may be added to the base of orogenic crust during melting of buoyantly ascending crustal material. Highly oblique subduction, such as proposed for the subducting continental lithosphere under Tibet (Tapponnier et al. 2001), may facilitate prolonged burial/heating and delayed exhumation, creating the right thermal conditions for melting during ascent (Fig. 3). Active collisional orogens provide information about the scale of continental subduction. For example, geophysical studies of subcrustal structure beneath the Himalayan orogen show that a zone of low-density material resides beneath Tibet as well as the western Himalayas (Nelson et al 1996; Van der Voo et al 1999). This zone comprises in part continental crust from the subducted Indian plate (Zhao et al 1993), as well as subducted continental material from the Eurasian plate (Tapponnier et al 2001). In the western Himalayas-Hindu Kush-Pamir zone of central Asia, geophysical evidence documents ongoing steep subduction of continental lithosphere, including low-density crustal rocks (the Indian continental margin). In this region, earthquake hypocentres (Searle et al 2001) and tomographic images based on P-wave velocities (Van der Voo et al 1999) show a steeply dipping region of continental lithosphere that has descended into the mantle to ~300 km depth. To the east, under southern Tibet, Indian lithosphere is currently subducting at a shallow angle (Zhao et al 1993; Owens & Zandt 1997). Evidence from active and ancient orogens

Fig. 3. Tectonic model for continental subduction, decompression and partial melting. Advanced subduction of continental material along multiple, sequential subduction zones (now sutured). Modified from the model of Tapponnier et al (2001). The oldest subducted continental material has risen buoyantly, melted and contributed to crustal thickening of the Tibetan Plateau.

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suggests that continental subduction is a major process in the evolution of orogens, and that exhumation of subducted continental rocks may be isothermal at high temperatures. The role of melting in this process remains to be discussed (see following section on 'Partial melting in orogens'). Buoyancy / diapirism Lower-density rocks underlying higher-density rocks may rise buoyantly if rheological properties allow flow. During orogeny, a regionalscale density inversion can be achieved in several ways: (1) tectonic stacking of terranes/ thrust slices that are more dense than underlying felsic basement; (2) subduction of continental crust into the mantle; and (3) reduction in bulk density of rocks by addition of felsic magma, with no large-scale segregation of melt. One or more of these situations, combined with the fact that rocks are generally weak at depth owing to elevated temperature and the presence of even small amounts of melt, creates a condition for gravitational instabilities to develop and grow (Fig. 2e). For the case of schist/ gneiss overlying granitic basement, the density contrast (~0.1-0.3 g cm~ 3 ) is sufficient for solid-state diapirism (Fletcher 1972; Soulaet al. 2001). The magnitude of decompression of rising diapirs may be significant, as flow is along subvertical trajectories for a significant portion of the ascent. Once initiated, diapirism may be self-sustained by a positive feedback relation between decompression and melting. The presence of melt lowers the bulk density within the diapir relative to its surroundings, enhancing upward flow and decompression. Rise of partially molten crust is accompanied by downflow of surrounding rocks, which may be transformed to granulite and eclogite that accumulate in the lower crust. The rate of decompression of deep crustal rocks by buoyancy-driven flow will be higher if diapirism is coupled with thinning of thickened crust, particularly if the upper crust is removed by extension, tectonic denudation and/or erosion. The rising diapir may localize upper crustal extension, or the removal of upper crust may drive the buoyant rise of partially molten crust. It is likely that these processes are coupled because they involve a positive feedback relation through the generation of melt (Teyssier & Whitney 2002). The rise of diapirs comprising partially molten crust can account for largemagnitude, near-isothermal decompression if the rate of ascent is fast enough to maintain

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elevated temperatures and some melt remains in the system.

Partial melting in orogens: mechanisms and consequences Two major processes that characterize collisional orogens are crustal melting and the exhumation of formerly deep continental material. These processes may be genetically linked, as demonstrated by the P-T-t histories of migmatite complexes, and in particular migmatite-cored gneiss domes that are observed in exhumed orogens worldwide. Migmatite diapirs and gneiss domes In orogens, the geological expression of partially molten crust that has experienced high-temperature, near-isothermal decompression is commonly a migmatite-cored gneiss dome. Gneiss domes occur in orogens ranging in age from Archaean through Cenozoic and in tectonic setting from wide (hundreds to thousands of kilometres wide) orogens to narrow (< 10 km wide) shear zones. Well-documented examples occur in the North American Cordillera, Himalayas (Zanskar, Karakorum, Garwhal, South Tibet), Pamirs, Alps (central Alps, Aegean region, Anatolia), Iberian and French Variscides, the Bering Sea region (Alaska, Russia), and the Appalachians, as well as numerous domes in Precambrian terranes. Typically, more than one dome is present in a region, they are elongate and aligned parallel to the strike of the orogen, and they are characterized by a core of anatectic migmatites, orthogneiss and/ or granitoids surrounded by high-grade metasedimentary rocks. In some orogens, there is a characteristic spacing between gneiss domes, e.g. 40-50 km in the northern Cordillera, and 25 + 5 km in the northern Appalachians (Fletcher 1972). The short dimension of gneiss domes ranges from ~4km (Naxos, Greece; Baltimore, USA) to 60 km (Velay, Massif Central, France), but most domes have a short diameter of 15-25 km. The origin of gneiss domes has been debated for more than 50 years (Eskola 1949). Their origin has been ascribed to diapirism (Berner et al 1972; Ramberg 1980; Calvert et al 1999), crustal shortening (Ramsay 1967; Burg et al 1984; Rolland et al 2001), extension (Chen et al 1990; Brun & Van Den Driessche 1994; Escuder Viruete et al 2000), or more complicated models invoking both contraction and extension (Lee et al 2000),

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crustal flow/extrusion (Beaumont et al. 2001), or extension-controlled upwelling of partially molten crust (Vanderhaeghe et al. 1999). In recent years, diapirism has been rejected as a general mechanism for the generation of gneiss domes because in many cases the structural doming event is believed to post-date partial melting/magmatism (e.g. Lee et al. 2000; Holland et al 2001). In most orogens, the crustal structure beneath gneiss domes is not known, but in southern Tibet, seismic profiles suggest that the partially molten (high conductivity, low seismic velocity) mid-crust extends to the base of the Kangmar dome (Nelson et al. 1996). Recent papers have suggested links between the deep crust of the southern Tibet and the Himalayan wedge, including flow of ductile Tibetan crust into the Himalayan wedge (Beaumont et al. 2001; Grujic et al. 2002). An example of the relationship between decompression and doming is found in the Shuswap metamorphic core complex, British Columbia, where a series of elongate gneiss domes are aligned along the strike of the belt near its eastern margin. These domes are approximately 15 km across their short axis, are located ~40-50km apart, and include the Frenchman's Cap, Thor-Odin, Pinnacles and Valhalla domes. The high-grade core of the Thor-Odin dome experienced near-isothermal decompression from P > 10 kbartoP < 4 kbar (Norlanderefa/. 2002). The dome is defined in part by the transition from migmatitic rocks that retain a coherent metamorphic layering, to migmatites dominated by the granitic fraction and characterized by lack of a coherent solid framework (Vanderhaeghe etal. 1999). Isothermal decompression of migmatite domes The advection of partially molten material from >30km depth towards the Earth's surface is a significant agent of heat transfer during orogeny. The temperature-time history at shallower levels depends on the final emplacement depth of the diapir and the geothermal gradient of the upper crust: these are controlled by surficial processes and upper crustal deformation (e.g. extension, erosion). Thermal modelling can be used to evaluate the conditions required for near-isothermal decompression and to examine the thermal effects of migmatite diapirs following ascent to the upper crust. We used time-dependent thermal modelling to quantify the relationship between isothermal

decompression of a diapir and cooling rates (Fig. 4). The advection-diffusion equation is solved in two dimensions using an explicit finitedifference method (e.g. Noye 1982). These models illustrate the thermal response of a diapir rising from the deep crust to the middle crust and do not address the mechanical issues of emplacement of diapirs into the rigid upper crust. Despite the mechanical limitations, important information regarding the relationship between pressure-temperature (P-T) and temperature-time (T-f) paths is obtained through this analysis. The predicted T-t paths (Fig. 5) can be compared to observed paths (Fig. 1) to evaluate the exhumation history. The diapir is modelled as a piston 15 km wide with an initial temperature of 775 °C (Fig. 4a). The diapir ascends from a depth of 30 km at a constant rate to mid-crustal levels. Points within the diapir are advected vertically; exhumation rates investigated by the model were 2, 5, 10, 15 and 20 km Ma"1. The more rapid rates are consistent with rates recorded by rocks exposed in migmatite-cored domes (e.g. Brown & Dallmeyer 1996; Calvert et al. 1999). The initial geothermal gradient is non-linear, with an increase in temperature from 320 °C at 15 km to 550 °C at 16 km. Below this interval, which represents the base of a rigid upper crustal lid, the geothermal gradient changes to 10 °C km"1. The results shown in Fig. 4 illustrate the effects of the diapir stopping at the base of the upper crustal lid (Fig. 4b) and piercing the upper crustal lid (Fig. 4c). Pressure-temperature-time paths calculated for various points within the diapir, according to the second model (the diapir pierces the upper crust) and an exhumation rate of 20 km Ma"1, are illustrated in Fig. 5a. These results show that, even at extreme exhumation rates, a rock at the top of the diapir (initial depth = 30 km) loses heat to the surroundings during decompression, and therefore records cooling during decompression. In contrast, rocks within the diapir (initial depth — 35 or 39 km) retain heat for a longer period of time during decompression, resulting in a component of near-isothermal decompression. The numerical experiments therefore predict that rocks within a diapir initiating within the deep crust maintain significantly high T (>700 °C) during decompression to shallow mid-crustal levels. Once the diapir is emplaced at a shallow level, the associated cooling paths predict a range of cooling rates from 24 °C Ma~ ! to as high as 400°CMa~ 1 (Fig. 5b). Both models shown in Fig. 4b and c predict greatest cooling rates of 400°CMa~ 1 during exhumation for rocks at

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Fig. 4. (a) Initial conditions for time-dependent numerical experiments. The initial geothermal gradient is determined using a heat production value of 9.6 x 10~4 jjiW kg~ ! for the upper crust and 5 x 10~5 jjiW kg~ for the lower crust. The boundary between the upper and lower crust is located at 15 km (arrow). The diapir is modelled as a piston, the top of which is initially at 30 km depth and moves vertically at a given rate. rinitial for points within the piston is 775 °C. Boundary conditions are constant basal heat flux and constant surface temperature, Tsurf = 20 °C. (b) Diapir is exhumed to the base of the upper crust at 20 km Ma"1. Dashed contours show position of isotherms after advection (1 Ma); black contours represent thermal structure after 5 Ma of cooling at t = 6 Ma. Contour interval = 200 °C. (c) Diapir is exhumed to a position within the upper crust at a rate of 20 km Ma'1. Contours are the same as in (b). In both models, rocks near the top of the diapir record the greatest cooling, but the magnitude of cooling is greater in the case where the diapir pierces the upper crust.

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Fig. 5. (a) Depth (pressure)-temperature diagram showing model results for a particle point located at the top of the diapir (initial depth, Zj = 30 km), a point 5 km below the top of the diapir (Zj = 35 km) and a point 9 km below the top (Zi = 39 km). The decompression rate for all three cases is 20 km Ma"1. The top of the diapir cools continuously during decompression, but the points modelled within the diapir experience near-isothermal decompression until the diapir pierces the upper crust, at which point the diapir experiences isobaric cooling. The grey shaded region illustrates the approximate location of dehydration melting reactions involving biotite. (b) The T—t plot corresponding to the same conditions as in (a) with similar labels and line styles. This T—t diagram illustrates the rapid cooling rates of the modelled processes.

the top of the diapir. Following exhumation, the cooling rate slows to 24 °C Ma"1. The average cooling rate recorded for a rock at the top of the diapir is 56 °C Ma"1. A rock within the diapir, however, records a slower average cooling rate of 40 °C Ma"1. These cooling rates are consistent with observed cooling rates determined by thermochronology of migmatite-cored domes (e.g. Calvert et al 1999; Vanderhaeghe et al 1999). Observed P-T-t paths for migmatite-cored domes suggest rapid isothermal decompression of the deep crust. Calculated P-T-t paths further support buoyancy-driven flow as a viable mechanism for rapid exhumation of these deep rocks. To understand how these domes develop in orogens requires addressing the general question of partial melting during orogenesis. In this chapter thus far, we have emphasized the importance of decompression specifically isothermal decompression - and the tectonic/thermal fate of partially molten crust. We have focused on relatively small(crustal-) scale observations, but will now consider larger- (lithosphere-) scale processes that might contribute to melting of continental crust. By doing so, we can address the fundamental question of how and why the deep crust of orogens attains such high melt fractions over such a great crustal thickness (Nelson et al. 1996; Schilling & Partzsch 2001).

Buoyant return of subducted continental crust The contribution of continental subduction to melting in orogens requires further consideration. In the mid- to deep crust of the southern Tibetan Plateau, it is difficult to explain the presence of the amount of melt inferred (20 vol.%) by simple heating during crustal shortening, as the mantle is not unusually hot beneath this part of Tibet (Nelson et al 1996; Chen et al. 1996). In addition, the rapid rate of underthrusting (subduction) of Indian crust beneath Eurasia (55-60 mm a"1; Patriat & Achache 1984; Le Pichon et al. 1992) and the estimated low temperature at the Indian plate moho at the modern suture (inferred from heat flow data; Gupta 1993) suggest that the lower crust of the subducting Indian continent is at moderate temperatures (800°C). The UHP terranes identified at the Earth's surface represent a small fraction of the total volume of subducted continental material. For example, the amount of Indian continental crust that is estimated to have been buried/subducted beneath the Eurasian plate cannot be accounted for by the amount of crustal thickening in the Himalayas. Some of the missing crust has likely flowed under southern Tibet. Although some subducted continental crust may become dense enough to sink into the mantle as eclogite (Le Pichon et al. 1992), some crust may remain less dense than the surrounding mantle rocks (Hermann 2002) and can rise buoyantly. The question of the fate of subducted continental crust is related to the present discussion because the buoyant rise of subducted continental crust may contribute to partial melting in collisional orogens, by providing source rocks for partial melting during decompression and/or by influencing decompression mechanisms in the overlying (non-subducting) crust. It is unclear at present whether exhumed UHP rocks were partially melted during ultrahighpressure metamorphism. Some UHP terranes contain abundant migmatites (e.g. Dabie Shan), but the general view is that the partial melting occurred in a later event, unrelated to UHP metamorphism (Coleman & Wang 1995). Furthermore, geochemical studies in some exhumed ultrahigh-pressure terranes suggest that the continental material was 'old, cold, and dry' (Coleman & Wang 1995) prior to subduction. Sharp et al. (1993), however, proposed that the stable isotope values of UHP schists from the Dora Maira Massif indicated equilibration with a melt, possibly represented by kyanite + jadeite + garnet + quartz layers in the schists (Schreyer et al 1987). We believe the relationships among UHP metamorphism, exhumation of UHP rocks and partial melting deserve further investigation, but, whether or not exhumed UHP terranes experienced partial melting, it is reasonable to propose that some subducted continental material has sufficient fertility and reaches P-T conditions appropriate for partial melting during deep subduction and/or decompression.

Summary Large-magnitude (>10km), regional-scale decompression that is nearly isothermal at elevated temperatures (>700°C) may be driven

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by a variety of mechanisms, including surficial and deep crust/mantle processes that may influence each other through feedback relationships. If decompression is a necessary condition for large-scale melting of continental crust, then a significant source of material and/or a major driving force for melting in collisional orogens may be subducted continental material. The transfer of continental material, including melt, from the subducting plate to the non-subducting plate may result in decompression-driven partial melting, with the melt accumulating in the nonsubducting plate. The large volume of melt in a layer within the thickened crust fundamentally changes the balance between tectonic and buoyancy forces in a collisional orogen and results in mechanical decoupling of continental crust and lithospheric mantle. Within continental crust, upper and deep levels may initially be mechanically coupled (e.g. if decompression is driven by surficial or other upper crustal processes), but these too become decoupled through time as partial melting proceeds and the deep crust flows laterally (channel flow) and/or vertically (diapirism). The rise of partially molten crust and associated downflow of denser country rocks signifies decoupling of orogenic crust from the mantle lithosphere and of deep crust from upper crust. We thank Olivier Vanderhaeghe for his comments on the paper, and acknowledge reviews by Mike Brown and J.-L. Vigneresse. In particular, the helpful suggestions of Mike Brown improved the discussion of migmatites and melting relationships. This work was partially supported by NSF grant EAR-9814669.

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Strain-rate-dependent rheology of partially molten rocks J. L. VIGNERESSE1 & J. P. BURG2 1 CREGU, UMR CNRS 7566 G2R, BP 23, F-54501 Vandoeuvre Cedex, France (e-mail: jean-louis. vigneresse@g2r. uhp-nancy.fr) 2 Geologisches Imtitut, ETH-Zentrum, Sonneggstrasse 5, CH-8006 Zurich, Switzerland (e-mail: [email protected]) Abstract: Coupling or decoupling in the lithosphere is often related to the absence or the presence of a layer of partially molten rocks, although the rheology of such rocks remains unsolved. Theoretical arguments correlated with structural observations provide new insights into the rheology of partially molten rocks, namely migmatites and magma. These rocks are simplified to two-phase pseudo-fluids constituted of a quasi-solid matrix and a variable amount of melt. Previous experiments indicate that the matrix deforms plastically according to a power law. The melt is Newtonian and weakens at high shear strain rates. Because of the heterogeneous distribution of matrix and melt phases, their rheologies cannot be averaged to obtain the rock rheology. Four behaviours are identified, (i) At high stress and strain rates, the viscosity contrast between melt and matrix is lowest. Both phases can accommodate strain at a comparable rate, allowing migmatite and magma bodies to deform as quasi-solid units, (ii) At low strain rates, the viscosity contrast between melt and matrix is highest. Melt deforms and relaxes much faster than the matrix. The simultaneous coexistence of a weak and a strong phase is expressed in a 3D viscosity-strain rate-melt fraction diagram, in which a cusp-shaped surface represents viscosity. The cusp graphically shows that the viscosity of partially molten rocks may jump several orders of magnitude. These jumps, leading to sudden melt segregation, are temporally erratic, (iii) At low strain rates, strain partitioning may lead to internal instabilities and segregation between melt and restitic phases, as observed in the leucosome/melanosome separation, (iv) Cyclic processes follow hysteresis loops and trigger strain localization.

Vertical or horizontal decoupling in the lithosphere requires that there is an interface whose rheology restricts stress transmission. This may occur along a discontinuity, as a fault in the brittle crust, or as weak shear zones in the ductile crust. Weakness is commonly attributed to a compositional change, including soft minerals (Jordan 1987), or to the onset of melting, the molten rocks providing the decoupling interface (Dewey 1988; Block & Royden 1990; see recent review by Vanderhaeghe & Teyssier 2001). At a smaller scale, coupling and decoupling are coeval in migmatites, which were partially molten crustal rocks (Mehnert 1968; Ashworth 1985; Brown 1994). The presence of melt induces strain partitioning and decoupling between the melt and its matrix (Vigneresse & Tikoff 1999). In contrast, large-scale structures of a migmatite body are concordant with those of the surrounding rocks, manifesting some kind of coupling. The rheology of partially molten rocks (PMR) has been approached from a theoretical point of view, importing into Earth sciences observations

and experiments from other disciplines such as soil science, food engineering, polymer rheology and chemical engineering. Crystallizing magma, and migmatites also, consist of solid crystals suspended in a fluid melt. They can both be considered Theologically as two-phase materials at the time of crystallization and melting, respectively. The present chapter identifies rheological behaviours and physical effects to be taken into account for investigating magma segregation and extraction. We first review results concerning the rheology of two-phase materials. PMR are considered to be viscous, with viscosity values rapidly changing from that of the strong to that of the weak phase (Arzi 1978; Barboza & Bergantz 1998; Renner et al 2000). With such considerations, melting and crystallization present symmetrical rheological properties, which is not supported by microstructural observations (Vigneresse et al 1996; Rosenberg 2001). To explore the many aspects of PMR rheology, we extend a previous model that included non-linear interactions between melting, strain partitioning and rheology (Burg & Vigneresse

From: GROCOTT, J., MCCAFFREY, K. J. W., TAYLOR, G. & TIKOFF, B. (eds) 2004. Vertical Coupling and Decoupling in the Lithosphere. Geological Society, London, Special Publications, 227, 327-336. 0305-8719/04/$15 © The Geological Society of London 2004.

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2002) and obtain catastrophic events and memory effects due to the hysteretic behaviour of PMR. We first present the intrinsic rheological responses of each phase. We then briefly discuss the limits of the averaging formulation. We attempt to reach a global formulation, which is highly strain-rate-dependent. The ensuing metastable states describe local instabilities. We finally discuss geological implications, but point out that these findings are only in a preliminary, mostly conceptual, stage.

Rheology of two-phase materials Basic rheological laws Rheology is the science of deformation and flow (e.g. Adler et al 1990). In elastic materials, strain disappears as stress is removed. In viscous materials, the strain rate (e) depends in a complex manner on the applied stress ( 10~4 5 s"1 (Webb & Dingwell 1990; Dingwell et al. 1996). In our model, the maximum strain rate is set to 1 s"1, equivalent to seismic strain, an estimated upper bound for strain-rate values (Scholz 1990; Reimold 1995). In all cases, strain concentrates into the melt (Vigneresse & Tikoff 1999). Complex nonlinear interactions occur, inducing feedback loops between melt production, strain response

and material rheology, and instabilities are a function of the melt fraction (Burg & Vigneresse 2002). In order to investigate in more detail the rheology of melt plus matrix systems, we map the stress-strain rate field in which PMR plot. Equations (1) and (2) are better formulated in logarithmic coordinates, thus allowing direct comparison of the two laws: for the melt and

for the matrix In this formulation, log rj — 1 corresponds to 107 Pa s, the Newtonian viscosity of a granitic melt (Spera et al. 1988; Clemens & Petford 1999). The C(7) term in equation (4) includes the coefficient A of equation (2) and a temperature-dependent term that includes the activation energy Q. We use experimental data with log A = -4.89 and Q = 243 kJ moF1 (Wilks & Carter 1990). We take 800 °C as the approximate temperature for the onset of dehydration melting of biotite (Patino Douce & Beard 1995). C(T) is then set to a constant value of — 32. In the range of strain rates 1Q-4.5 < g < io°s~ 1 , we let viscosity decrease with strain rate (Dingwell et al. 1996). In that range of strain rates, a power law in which n — 2 best matches experimental data. Those values adopted must be considered as average assessments. Departures due to changes in melt composition or temperature (±100°C) do not strongly change the parameters. They vary by about ± 1 logarithmic unit in viscosity, a factor of 10. In all cases this does not alter the shape of the stress-strain rate diagram (Fig. 1). In Figure 1, the melt and the matrix curves intersect at a point with coordinates (12.5, 5.5) if no shear weakening occurs. With shear weakening, the point is located at (15, 12). Both intersections occur at stress and strain-rate magnitudes that are not significant in geological conditions. However, intersection means that, for the corresponding parameters, the viscosity of the matrix is similar to that of the melt. Bracketing geological stress between 0.1 and lOOMPa (5 to 8 in log coordinates), the two lines do not coincide and delimit an irregular pentagon, which is the region that contains the rheology of natural PMR in a log-log diagram. In Cartesian coordinates, this region turns into a triangle because the strain rate remains very low (10~17 to 10° s"1) compared to the stress range (105 to 108 Pa). Within this triangle, the couple (cr, e) varies non-linearly with the melt content.

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Adopting equation (5) or equation (6) is not trivial and has implications for the theoretical behaviour of each phase, because the two descriptions correspond to specific boundary conditions. The arithmetic description corresponds to the same strain rate applied to both phases. In an elasto-viscous PMR, the arithmetic mean viscosity cannot fit the fluid response if the elastic component is significant. It should be selected for a solid-like behaviour in which small deformations are damped. This is the case in a crystallizing magma. Conversely, the geometric formulation supposes that the same stress applies to both phases, which respond as a fluid to a steady or stationary state. The fluidlike behaviour allows the accommodation of large strain with partial elastic recovery. It applies to melting migmatites. The rheological equation of a crystallizing magma can thus be expressed as Fig. 1. Log-log stress-strain rate (cr-e) diagram displaying the relation (3) for a Newtonian melt (n = 1), which turns to shear thinning for e > 10~45 s"1, and the power law (4) (n = 3) for the matrix (see references in the text). Grey lines indicate the corresponding viscosity values.

Hence, we must now investigate the effects of the melt fraction. Rheology of a two-phase pseudo-fluid The rheology of PMR implies introducing a proportion of solid phase into the melt, i.e. solid fraction () relative to melt; O ranges from 0 in the fluid-like case to 1 in the solid-like case. The simplest formulation of bulk rheology consists in estimating either the average strain rate or the average stress that simultaneously applies to both phases. They form the extreme ideal formulation of a mixture. Owing to the large viscosity contrast, averaging is not particularly useful for PMR. The arithmetic and geometric averaging equations are the simplest formulations (Gueguen & Palciauskas 1992). A system composed of two materials with viscosities 77! and 172 has an arithmetic (or Voigt) mean viscosity j]a given by

and a geometric (or Reuss) mean viscosity j]g given by

which can also be written as

This is the general Herschel-Bulkley equation (Adler et al 1990). For melting migmatites, the rheological equation can be expressed as

or alternatively as

This is the general equation for thixotropy (Adler et al. 1990; Barnes 1997). Suffixes 1 and 2 correspond to melt and matrix, respectively. The corresponding viscosity values are plotted with respect to the solid fraction (Fig. 2). The viscosity contrast between the two phases is varied from 101 to 109 Pa s to emphasize differences between the two averaging methods. The arithmetic mean favours the largest values of the average at the expense of small ones. Conversely, the geometric average is damped by the small values of the average. In other words, the arithmetic average favours the strong phase, whilst the geometrical average gives much more weight to the weak phase. We conclude that there is no correct way to address the problem of a heterogeneously

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Fig. 2. Viscosity contrast between the liquid and solid phases as a function of the solid fraction (4>) for a PMR. Different values of the viscosity contrast (102, 105, 108, 1011 bottom to top) are presented. For each, two curves correspond to the arithmetic (r/a) and geometric (rjg) averages, respectively. They define a system that develops at constant stress and constant strain rate, respectively. The geometric average does not show many differences between a contrast of 102 and 1011 in the end-member values.

varying phase proportion. Any model for averaging rheological properties, whether arithmetic or geometric, fails to represent adequately the rheology of PMR. Averaging formulations involve either stress or strain-rate boundary conditions that apply simultaneously to both phases. Hence, a unique formulation cannot indifferently describe melting and crystallization that proceed from different boundary conditions. A second consequence of averaging methods is that the resulting equations for bulk flow involve both melt and matrix flow. The flowing material is a pseudo-fluid. This may be the case when the viscosity contrast is restricted to 1-3 orders of magnitude. In the case of PMR, the viscosity contrast is huge, up to 15 orders of magnitude, and varies with the imposed strain rate. A limit to the flow would be a Darcy flow, in which the matrix is fixed, which is obviously not a correct model for PMR. Therefore, we suggest another way to examine the rheology of a two-phase material. Rheology of a two-phase material Representing the PMR rheology implies introducing an amount of solid fraction as the third dimension in a (cr, s, ) diagram. The applied stress is difficult to estimate. Therefore, we prefer to display the rheology in a diagram scaled with viscosity (Fig. 3). Several considerations are prerequisites.

1 The PMR viscosity is constrained by experimental measurements (Lejeune & Richet 1995; Renner et al 2000) and numerical applications (Vigneresse et al. 1996; Barboza & Bergantz 1998). They demonstrate the increase of viscosity with particle concentration, and thresholds that apply. 2 A 3D diagram must take into account the non-linear effects discussed earlier (Burg & Vigneresse 2002). 3 The bulk melt content is unsatisfactory to define a PMR. We must consider the coexistence of two phases with contrasting rheology. 4 The 3D representation should also coincide with the known viscosity dependence on the solid fraction (Lejeune & Richet 1995) and the solid-like behaviour of concentrated suspensions (Rogers et al. 1994). 5 The model should simultaneously describe the small- and large-scale behaviour of PMR, which correspond to short and long relaxation times of melt and matrix, respectively. The irregular pentagon of Fig. 1 turns, in the 3D diagram, into a cusped viscosity surface integrating the coexistence of the two phases (Fig. 3). In consequence, the strain response to stress, or viscosity determination, is not a single parameter. It varies strongly with applied stress, shear rate or solid fraction. This first-order conclusion strongly contrasts with the equivalent viscosity that would apply for a given solid fraction, independently of strain rate (Arzi 1978). Rheologies for low (^> < 10%) and high ( > 60%) solid fractions are both easy to estimate. Such PMR behave almost exactly as the solid and fluid end-members, with non-linear modifications of the effective viscosity according to the Einstein-Roscoe model of suspensions (reviews in Adler et al. 1990; Liu & Masliyah 1996; Vigneresse et al. 1996). For intermediate solid fractions (10-60%), complexity due to non-linear interactions occurs (Burg & Vigneresse 2002). We first examine the case of low strain rates. The viscosity contrast between melt and matrix is highest (Fig. 3). The relaxation time of the melt is much shorter than for the matrix. This difference generates an instability expressed by the cusp shape of the surface representing viscosity (Fig. 3). The coexistence of the two phases over a range of percentages (0.20 < < 0.75) results in a metastable state. For a given solid fraction, three viscosity values are theoretically possible. The highest and the lowest correspond to stable states and reflect

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Fig. 3. Diagram for strain rate-viscosity-solid fraction (e-Tj-O) that relates strain rate and viscosity according to the proportion of one phase.

the viscosity of the matrix and the melt, respectively. The third, intermediate one is unstable. It reflects the response of a PMR to stress and may fluctuate from the viscosity of the matrix to that of the melt. The fluctuation is rapid, similar to catastrophes as formulated in mathematics (Zeeman 1976; Thorn 1990; Wassermann et al 1992). At higher strain rates, the viscosity contrast between melt and matrix is a few orders of magnitude (Fig. 1). The relaxation times of both phases allow accommodation of their relative motions and there is no instability. The difference in relaxation times between melt and matrix has a second consequence expressed by hysteresis in a stress-strain ratesolid fraction diagram (Fig. 3). The metastable state develops over 0.20 < < 0.75, as stated above. However, it also develops for a fixed percentage of one phase, but at varying strain-rate values. For example, at 70% of melt, and low strain rate (Fig. 3), the sequentially loaded and unloaded PMR may suffer cyclic deformation, but the strong phase has no time to relax completely. Strain adds up at each cycle, resulting in large localization of strain in the weak phase, like it has been experimented on water-saturated sand and clay (Bardet 1995, 1996). In a similar way, experiments on partially saturated sediments,

which have a non-linear rheology, report a drastic increase of strain amplitude while the material is submitted to low-frequency stress cycling (Guyer et al 1995; Tutuncu et al 1998).

Geological implications The levels of stress and strain rate differ when considering the bulk behaviour of a PMR or only local melt segregation. In the following, we discuss some geological implications of our model. Bulk response of PMR to tectonic stresses (high strain rate) PMR is of regional extent (103-105 m) and the ambient stresses are tectonic. We first address melting metatexites, which are low-melt (1020%) migmatites. They commonly form under 400-800 MPa and 750-800 °C (Brown 1994). Metamorphic pressures correspond to the 1025 km depth of the brittle-ductile transition where fracturing occurs at 0.8 to 4 times the vertical load (Byerlee 1978): 100-400 MPa at 1012 km depth, depending on whether tectonics are extensive or compressive. Accordingly, we may assume an ambient differential stress of about 108 Pa (8 in the log-log diagram, Fig. 1). The

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corresponding strain rate is in the range of 1 to —3 in log scale, with viscosity of 10 7 Pas for the melt and 1010 Pa s for the matrix. The corresponding viscosity contrast is 103 (Fig. 4). The response to ambient tectonic stress is thus mostly the solid-state behaviour of the matrix. One may thus expect to observe structures that are similar to those of the surrounding nonmolten rocks (Nzenti et al 1988; Brun & Bale 1990). High tectonic stresses may impose correspondingly high strain rates onto a granitic body with 40-50% crystals. Although the magma is still rich in melt, it may react as a solid, up to fracturing (Dingwell 1997). At moderate strain rate, plastic deformation locally re-orientates the crystalline fabrics, inducing proto-faults as described in the Mono Creek massif (Saint Blanquat et al 1998). The bulk reaction of a two-phase system to high stress level is common in geology. At very fast strain rates, as during earthquakes, saturated sediments that are ordinarily very soft and without cohesion behave en masse during soil liquefaction (Ishihara 1993).

Bulk response of PMR to low strain rates PMR has similar regional extent as previously (103-105m), but the deformation rate of 10^ 16 s -1 is below the usual strain rates in geology (Pfiffner & Ramsay 1982). This situation applies to the lower continental crust, in which some weakness may form, either from grain size reduction, leading to strain localization, or from partial melting, that may induce magmatism. Such situations are inferred in lower crustal sections presently brought to the surface as in Kohistan (Arbaret et al. 2000), or deduced from geophysical studies as under the Andes (Schilling et al. 1997), the Pyrenees (Partzsch et al. 2000) and the Tibetan Plateau (Hauck et al. 1998). The amount of weak phase does not need to be high. Estimates of 4% melting are provided in the Pyrenees and the Tibetan Plateau (Partzsch et al. 2000). Nevertheless, connection of any melt is assumed from electrical conductivity, which suffices for strain partitioning. In such cases, decoupling can be assumed to develop at very low strain rate.

Fig. 4. Same diagram as Figure 3, but on which we have located: the specific cusp shape instability for vein segregation; the quasi-continuous variation of viscosity that rules the deformation of the bulk massif at low strain rate; the instability due to viscosity variation leading to melanosome segregation from the leucosome; and finally, the zone where hysteresis takes place.

RHEOLOGY OF A TWO-PHASE MATERIAL

Melt segregation at outcrop scale At the mesoscopic scale (1-10 m), low stress is added to the background field and results from local readjustments after en masse motion of melt and matrix. The large viscosity contrast between melt (107 Pa s) and matrix (10 13 Pas) implies a large strain-rate contrast. Strain partitions into the melt and relaxes much faster than in the surrounding matrix. In consequence, the rheological contrast between melt and matrix accentuates at low strain rate, leading to instability, whatever the melt percentage. If stress cycles occur while melt and matrix deform with the same strain rate, sudden jumps between the liquid-like and solid-like rheology, or vice versa, correspond to sudden expulsion of the melt. The low relaxation rate of the matrix makes the periods of melt segregation irregular and discontinuous in time. Irregular bursts of melt segregation have been obtained in numerical (Vigneresse & Burg 2000; Rabinowicz et al 2001) and analogue (Rosenberg & Handy 2001; Barraud et al. 2001) modelling. Conversely, deformation involving the same stress on melt and matrix leads to strain localization (Fig. 5). Repeated loading and unloading increases and amplifies strain localization, resulting in thin shear bands in which a small amount of melt may have focused deformation. With further deformation, the weaker shear bands may be reactivated, showing very large finite strain (see photos in Crawford et al. 1998). Melt segregation within a vein At smaller scale (10~2-1 m), strain partitioning and strain rate are high in the melt because its relaxation time is short. We do not wish to

Fig. 5. Instability developing at quasi-constant stress and leading to strain localization.

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enter the controversy about the actual proportion of molten and restitic phases and the effects of backreactions changing this proportion (Kriegsman 2001). We simply assume that the flowing melt is a suspension of crystals in a fluid. In suspensions, the velocity gradient and strain partitioning produce a plug-like concentration of solid particles in the centre of a channelled flow, known as the Bagnold effect (Bagnold 1954; Komar 1976). It is common in dykes and in crystal-rich lava flows (Correa-Gomes et al. 2001). In contrast, a concentration of restitic minerals (melanosome) on both sides of channelled melt (leucosome) is commonly described in migmatites. The flat shape of restitic biotites is generally parallel to leucosomes and flow direction (McLellan 1988). This orientation suggests that those minerals have passively rotated and side melanosomes seem to correspond to outward concentration of solid particles, more or less symmetric with respect to the central flow plane. Segregation towards the borders of the flow, as suggested by melanosome-bounded leucosomes in migmatites, thus differs from solid-phase segregation in dykes. It is due to both strain partitioning and chemical effects (Olmsted 1999; Tanaka 2000). Strain partitioning concentrates each phase along the flow direction and perpendicular to vorticity (Fig. 6). Exchange of elements between biotite and microcline (Olsen 1977) added to the mechanical effect could explain why biotites concentrate along each side of the leucosome, with their flat face aligned parallel to the flow direction (McLellan 1983). The concentration of restitic minerals on

Fig. 6. Catastrophic jumps in viscosity that take place at constant low strain rate, between places of different melt content. They lead to segregation of the melanosome either side of leucosomes, as occurs during spinodal decomposition.

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both sides of the leucosome would therefore be a consequence of instability under a common strain rate. Viscosity determination for PMR A direct consequence of time- or strain-ratedependent rheology is the non-unique determination of an equivalent viscosity for a PMR, given its proportion of melt. More generally, a two-phase material cannot be represented by one single equivalent viscosity. Extending the discussion a step further would be to link the viscosity determination to an underlying mechanism. Power-law rheology classically corresponds to dislocation creep (Kirby & Kronenberg 1987). Diffusion creep corresponds to Newtonian flow. However, recent experiments show that a transition from dislocation creep to diffusion creep may develop when a small fraction of melt is added to an aggregate of solid grains (Kohlstedt et al. 2000). A specific 'granular flow' has been coined for this situation (Paterson 2001). It manifests by a prompt switch from power-law (n — 3) to linear (n = 1) rheology (Kohlstedt et al 2000), which is the situation our model describes. Conclusions We imported models from other disciplines to approach the rheology of partially molten rocks (melting migmatites and crystallizing magmas). We point out that PMR is a material in which it is impossible to assign a given phase proportion at any point. Therefore, average laws fail to represent PMR rheology correctly. To alleviate the problem, we first mapped the PMR rheology within a stress-strain rate diagram in which we could incorporate the phase content. Results show hysteresis and catastrophic jumps of viscosity due to the coexistence of a solid and a fluid phase. Several responses may develop, depending on whether the system responds under high or low strain rate or stress level. We infer that a migmatite body and plutons deform as solid-like mechanical units under high tectonic stress level. At a smaller scale, and under low stress level, instabilities develop under steady strain rate or quasi-constant stress and produce strain localization. Since instabilities involve hysteresis, cyclic loading and unloading may lead to large strain concentrated in small regions. If the same strain rate is imposed on fluid and solid phases, melt segregation and increased vorticity cause veins transecting layering and flow-induced mineral

separation such as restitic minerals aside leucosomes. This paper resulted from numerous and enriching discussions with many colleagues. It also came out from several discussion meetings at Zurich, for which CREGU founding is acknowledged. C. Rosenberg and F. Spera provided quite thoughtful comments.

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Index Page numbers in italics, e.g. 59, refer to figures. Page numbers in bold, e.g. 169, signify entries in tables. absolute plate motion (APM) compared to relative plate motion (RPM) 58-60, 59 model 29-30, 29, 30 Acapulco 118, 119 Acatlan 779, 722 Agua Blanca fault (ABF) 67, 172 Aguilas 250 Alboran Sea 250 Alicante 250 Almeria 250 Alpine fault 96, 199, 200, 209 Amami Sub-basin 300 Amami-Kagoshima Tectonic Line (AKTL) 299 Ammassalik erogenic belt (Am) 233 Animal Basin 67 Anita shear zone 200, 208 Anorituup Kangerlua 233 Arguello fault zone 67 Arrollo Taibena Basin 255 Arsuk 233, 235 Arsuk Brae 235 Arsuk Fjord 233, 235 Arthur River complex (ARC) 200 Asemi, River 287, 283, 285, 287 Aso, Mount 282 caldera 299 Atenango 722 attachment see coupling of limospheric layers Avaqqat Kangerluat 233, 237 Awatere fault 96 Balsas Basin 722 basal traction driven rotation 69 Baza 250 Beartooth Mountains (BT) 779 Beppu Bay 299, 305 Besshi nappe 287, 283, 285, 287 Betic Cordillera, Internal-External boundary hanging wall deformation 249-250, 273-275 evolution of structures and implications for coupling and decoupling 273, 274 geological map 250, 255-256, 260 geological setting 250-253 External Zones cross-section 252 pattern of vertical-axis rotations 272-273 regional constraints 271-272 rock successions 253-259 stratigraphy 254 structures 259-271 cross-sections 257-258, 261, 270 poles of bedding 265

Rambla Seca Basin 262 striations and fault planes 264 Bitterroot extensional complex (BC) 184, 186-187 Borgs Havn granite 235 brittle-ductile arc-parallel extension 279, 293-294 brittle deformation in Sambagawa 288 extreme ductile layer, normal thinning and arc-parallel stretching 288-290 exhumation scenario for Sambagawa 292-293, 293 normal fault development 289, 290 spacial distribution of recrystallized quartz grain sizes 291 strike-slip displacements 290 uniform shear sense and reversal by late-stage folding and faulting 290-291 variable strain geometry of exhuming rocks 292 five possible exhumation mechanisms 280 Butsuzo Tectonic Line (BTL) 282, 299 Cadiz 250 California see also San Andreas fault absolute plate motion (APM) and relative plate motion (RPM) 50, 59 Continental Borderland 65-66, 79-80 bending fault termination 76-78, 76 Pacific-North America tranform plate boundary 66-68, 67, 68 straight fault termination 69-76, 70 strike-slip fault termination styles 66, 68-78 vertical coupling and decoupling along WTR boundary 78-79 seismic anisotropy applicability of models 36-37 comparison between north and south California 35-36 model results 36 northern California 35 shear wave splitting measurements 70 southern California 32-34 shear-wave splitting 50 California, Gulf of 67 Camp Oven Creek (CO) 200 Capas Rojas Formation 266 Cartagena 250 Caswell Sound 799, 200, 209 paragneiss 208 P-T data 201 Cauta 250 Charles Sound 799, 200, 209

INDEX

338

Cheviot Hills 74 Chichibu metamorphic belt 281, 282 Chilapa 722 Chilpancingo 118, 119, 122 Chilpancingo Basin 722 Chino Hills 68, 75 Chugoku 143 Clarence fault 96 clutch tectonics 41-42, 51-52, 60 bottom-driven systems 53 implications 53 convergence 54, 57 divergence 54, 56-57 strike-slip partitioning and homogeneous mantle deformation 58, 58 transcurrent boundaries 53-56, 54 relation between crustal and mantle deformation 52 top-driven systems 52-53 Coacoyula 722 Coast Mountains-Cascades (CMC) batholith belt 169,

170 Coast shear zone (CSZ) 186-187 Cocula 722 Coeur d'Alene 184 Collnet Basin 67 Columbia River embayment (CRE) 779, 186-187 continental crust 313-314, 323 melting 314 melt segregation mechanisms and scales 314-315 P-T paths 315 near-isothermal decompression paths and mechanisms 315—316, 377 buoyancy/diapirism 317, 319 crustal thinning/collapse 316, 377 erosion 316, 377 exhumation of ultrahigh-pressure rocks 377, 318-319,378 folding/buckling 316-318, 377 low-angle normal faults 316 partial melting in orogens 319 buoyant return of subducted continental crust 322-323 isothermal decompression and migmatite domes 320-322, 327, 322 migmatite diapirs and gneiss domes 319-320 continental tectonics 1 Copalillo 722 Copalillo Basin 722 Cordoba 779, 250 coupling of lithospheric layers 1, 2 attachment formation during partitioning 231-232, 246 development of attachment zone 245-246 Psammite and Pelite Zones 244, 244 analogue modelling model construction 124-126, 725, 126

model results 129-132, 730, 737, 733, 734 model rheological structure and analogue materials 126-127,727 scaling of models 127-129, 128 qualitative comparison of model results with geology 134-136,735 model limitations 132-134 vertical coupling and decoupling 136 Cuautla 779 Cucamonga fault zone 68 Cuernavaca 118, 119 Cuevas del Ambrosio 268-269 Danell Fjord 233, 237, 239 Daniel, Mount (MD) 200 P-Tdata 201, 202 decollements 1-2 displacement transfer 177-178, 191 kinematic model 189-190, 189 decoupling of lithospheric layers 1, 2 diapirism 377, 319 dip-slip fault systems 1 Doubtful Sound 799 Dozan, River 287, 283, 285, 287 Eastern Gabar Basin 255 Eastern Transverse Ranges, California 92-94, 93, 94 Edgar, Mount (ME) 200 edge driven rotation 69 Egger 233 Elsinore fault 67 Elysian Park 74 Embalse de Valdeinfierno 262 fast direction polarization 9 Ferrelo fault 67 finite strain-controlled anisotropy 15-16 Fiordland, crustal attachment zone evolution 197-198, 223-226, 225 crustal structure and geochronology 207 age of magmatism, crustal melting and high-grade metamorphism 210 boundaries of high-grade metamorphic belt 207, 208, 209 geochronological data 203-206 lithological divisions of magmatic arc 207-210 structural relationships 210-212 evolution stage 1 - mafic-intermediate magmatism and partial melting of lower crust 213-215,274 evolution stage 2 - melt segregation and transfer mechanisms 215-218 (a) melt-induced fracture propagation 215-218, 277

(b) melt accumulation in ductile shear zones 218, 279 evolution stage 3 - evolving styles of deformation following magmatism and crustal melting 218

INDEX (a) steeply dipping sinistral and dextral shear zones 218-219 (b) gently dipping, layer-parallel shear zones 219, 227

(c) steep Indecision Creek and George Sound shear zones 220-223, 222, 223 geological map 200 geological setting 198-207 location map 799 P-T data 201-202 space-time correlation of high-grade fabrics 212-213 Foreland batholith belt 168 Fraser-Straight Creek fault (FSC) 186-187 Gabar 255-256 Garlockfault777 George Sound 799, 200, 209 P-T data 201 steep shear zone 220-223 Geurrero Morelos Platform 722, 123 Gibraltar 250 Goto Islands 299 Goto Sub-basin 299 Graenseland 233, 235 Granada 250 Guadalquivir Basin 250 Guadalupe fault zone 67 Guadalupe Microplate 67 Guadalupe Rift 67 Guadix 250 Hikimi 746, 750, 757 fault-slip data 155 Hiroshima City 148, 150 Hitoyoshi Basin 299, 305, 306 Hokkaido 143 Honshu 143, 299 Hope fault 96 Hosgri fault 67 Huajuapan 779 Huiziltepec 722 Igutsaat Fjord 233, 237 Ikermit 235 Indecision Creek 799 steep shear zone 220-223, 222, 223 Intermontane batholith belt 168 Ippatit 233 Ippatit Valley 239 Itoigawa-Shizuoka Tectonic Line (ISTL) 143 Izu Peninsula 143 Izu-Bonin Arc 143 Japan see also Sambagawa block rotation and intracrustal vertical decoupling 141-142, 158, 160

339

fault kinematics 152 age of recent inversion of motion sense 155-156 fault-slip data 754, 755 Plio-Quaternary to present-day kinematics 152-153 pre-Plio-Quaternary kinematics 154, 756 general geology 142-145 tectonic framework 143 SW island arc 297, 310 diffuse extension across south Kyushu 301-307 geodynamical and geological outline 297-301, 299 Okinawa-Kyushu junction area evolution model 307-310 Western Chugoku fault system 145-146, 148 age 152 analogue model similarities 156-158, 757 earthquake focal spheres 753 field occurrence 150-151, 757 formation model 159-160, 759 geometry of fault system 147-150 Kake-Himini area 750, 757 northern boundary 147, 146-147 Yamaguchi area 749, 757 Yoshiwa-Kake area 146 Julianehab batholith 232-234, 233 Border Zone-Julianehab batholith boundary 236, 237 construction and deformation 244 east coast structure 236-238 evolution 243 Julianehab batholith-Psammite Zone boundary 238 west coast structure 236 Kake 746, 750, 757 fault-slip data 755 Kamio, River 257, 283, 255, 287 Kangerluaraq 233, 237 Kangerluk 233, 237 Kangerluluk 233, 237 Kanoya Plain 299 Kap Farvel 233, 237 Kap Ivar Huitfeldt 233, 237 Ketilidian orogen, oblique convergence and attachment formation 231-232, 246 Border Zone structure Border Zone-foreland boundary relationships 234, 235 history 234-236 development of attachment zone 245-246 Julianehab batholith structure Border Zone-Julianehab batholith boundary 236, 237 east coast structure 236-238 west coast structure 236 major components 232, 233 Border Zone 232 Julianehab batholith 232-234

340

INDEX

Psammite and Pelite Zones 234 Psammite and Pelite Zones Julianehab batholith- Psammite Zone boundary 238 nature and timing of structural and metamorphic events 240 rapakivi suite 240-241 structure 238-240, 239 tectonic evolution 241, 242 construction and deformation of Julianehab batholith 244 development of mid-crustal attachment structure in Psammite and Pelite Zones 244, 244 rapakivi granite intrusion 245 structural histories 241, 243 Kettle extensional complex (KC) 779 Kikai Caldera 299 Kobberminebugt 233, 235, 243 Koshiki Islands 299 Kumamoto 252 Kyushu 143, 299 diffuse extension across southern regions 301, 305-307 Beppu region 305 Hitoyoshi-Ichifusa region 305, 306 Osumi region 301-305, 302, 303, 304, 305 geological structure 301 Okinawa-Kyushu junction area evolution model accomodation of extension through reactivation of thrust faults 307 cross-section model 307-310, 309 perpendicular extention 308 transtension in northern region 307 Laramide shortening, vertical coupling controlled by crustal heterogeneity 117-121, 137 lattice prefered orientation (LPO), olivine 3, 15 deformation types 55 mantle fabric observations 42-43, 44 Lewis and Clark lineament (LCL) 779, 184, 186-187 Lindenow Fjord 233, 237, 239 lithosphere, idealized deformation diagrams 54 Los Angeles 68 faults and seismicity 75 Los Angeles Basin 74 Malaga 250 Manapouri, Lake 799 mantle-driven deformation of orogenic zones 41-42, 60 crustal deformation and mantle fabric ancient orogens 48 interpretation of data 49-51 neotectonic orogens 49 strain history 51 lithosphere/asthenosphere connections continental settings 47-48 cratons 46

oceanic settings 46-47, 47 lithospheric deformation 42 mantle fabric laboratory experiments 43 numerical experiments 43-45 olivine LPO and shear-wave splitting 42-43, 44 mantle viscosity and seismic attenuation 45-46 Marbella 250 Maria 255-256, 260 Marlborough fault system, South Island, New Zealand 94-95, 96 Masuda City 148 Median Tectonic Line (MIL) 142,143,148, 282, 285, 299 Mexico, Caribbean-North American transform boundary 117-120, 137 analogue modelling of Late Cretaceous to Early Tertiary deformation model construction 124-126, 725, 726 model results 129-132, 130, 131, 133, 134 model rheological structure and analogue materials 126-127, 727 scaling of models 127-129, 128 geological and tectonic setting crustal structure 120, 727 Early Tertiary deformation 121-123, 722, 723 Laramide deformation 120-121 Tertiary deformation 124 lithological units 779 qualitative comparison of model results with geology 134-136, 135 model limitations 132-134 vertical coupling and decoupling 136 terrane boundaries 118 Midternass 233, 235 Milford 200 Milford Sound 799, 200, 209 P-T data 201, 202 Missoula 184 Mixteco Terrane 722 Mixteco-Oaxaca-Juarez block (MOJB) 120, 727 Early Tertiary deformation 121-123, 122 Laramide deformation 120-121 Tertiary deformation 124 Mogens Heinesen Fjord 233, 235 Montana disturbed belt (MDB) 779 Monterey fault zone 67 Monterey Microplate 67 Morro fault zone 67 mountain belts 1 Murcia 250 Nagssugtoqidian orogenic belt (Nag) 233 Nankai Trough 143, 299 Nanortalik 233 Napasorsuaq Fjord 233, 235, 237 New Zealand

INDEX absolute plate motion (APM) and relative plate motion (RPM) 50 seismic anisotropy 9-13, 31-32, 32-34 applicability of models 36-37 model results 36 modelling results 16-31, 17, 18-19, 20-21, 22

shear wave splitting measurements 10 study methods 13-16, 13 shear-wave splitting 50 Newport-Inglewood fault zone 68, 74-76 Niaqornaarsuk 233 Nobeoka Tectonic Line (NIL) 299 N0rrearm 233, 239 North American Cordillera 167-168, 168 crustal architecture 178-183 cross-sections 181-182 crustal thickening and deep flow 173 -174 crustal thickening and unroofing 168-169 Central Cordillera 169-170 Northern Cordillera 169, 170 Southern Cordillera 170-171 displacement transfer in decollement systems 177-178, 185-187, 191 generalized tectonic map 179 Late Cretaceous-Tertiary structural elements 186-187 oblique ramp system in Idaho-Montana basement 183-185, 184 coupling v. decoupling 190-191 influence on strike-slip systems 188 influence on Tertiary extensional systems 188-189 kinematic model for linked decollement system 189-190,759 regional expression 187-188 plateaux 171-172 unroofing mechanisms 172-173 Nunnarsuit 233 Oaxaca 118, 119 Oboke nappe 281, 283, 285, 287 Oita 282 Oita-Kumamoto Tectonic Line (OKTL) 252 Okanagan extensional complex (OC) 779 Okinawa trough back-arc basin 297, 310 diffuse extension across south Kyushu 301, 305-307 Beppu region 305 Hitoyoshi-Ichifusa region 305, 306 Osumi region 301-305, 302, 303, 304, 305 geodynamical and geological outline cross-section 300 Kyushi geological structure 301, 302, 303, 304, 305 present-day plate configuration and recent evolution 297-298, 299 structure 298-301

341

Okinawa-Kyushu junction area evolution model accomodation of extension through reactivation of thrust faults 307 cross-section model 307-310, 309 perpendicular extention 308 transtension in northern region 307 Olinala 722 olivine lattice preffered orientation (LPO) see lattice preffered orientation (LPO), olivine Omineca batholith belt 168 Omineca-Sevier batholith belt 168 Orizaba 118, 119 Orofino shear zone (OSZ) 779, 184, 186-187 orogenic float 1 erogenic zones, mantle-driven deformation 41-42, 60 crustal deformation and mantle fabric ancient orogens 48 interpretation of data 49-51 neotectonic orogens 49 strain history 51 lithosphere/asthenosphere connections continental settings 47-48 cratons 46 oceanic settings 46-47, 47 litho spheric deformation 42 mantle fabric laboratory experiments 43 numerical experiments 43-45 olivine LPO and shear-wave splitting 42-43, 44 mantle viscosity and seismic attenuation 45-46 Osburn fault 184 Oshima 282 Osumi Peninsula 299, 301-305 ages of pseudotachylite veins 303, 304, 305 cross-section of Osumi pluton 302 geological map 302 Otte Rud 0er 235 Oxnard 77 Oztotitlan Basin 722 Paatusoq 233, 237 Palos Verdes fault zone 68, 76-77 Papalutla 775 Papalutla Thrust 722 Patton Escarpment 67 Patton Ridge 67 Pembroke Valley (P) 200 P-T data 201, 202 Peninsular Ranges batholith (PRB) batholith belt 169, 772

Pinchi fault (PI) 186-187 Poison Bay 200 P-T data 201, 202 Priest River extensional complex (PC) 779, 184 Prins Christian Sund 233, 237 Puebla 778, 779 Puente Hills 75

342

INDEX

Puerto Angel 118, 119 Puisortoq 233, 235 Puisortoq Fjord 235 Qaqortoq 233 Qernertoq 233, 237 Qomoq augen granite 235 Qoornoq 235 Rambla Seca Basin 255, 262 cross-section 263 Raymond fault 68, 75 relative plate motion (RPM), compared to absolute plate motion (APM) 58-60, 59 rheology of partially molten rocks, strain-rate dependency 327-328, 334 basic rheological laws 328-329, 329 bulk response to low strain rates 332 bulk response to tectonic stress 331-332, 332 melt segregation at outcrop scale 333, 333 melt segregation within a vein 333-334, 333 pseudo-fluids 329-330, 330 two-phase materials 330-331 viscosity determination 334 Ryoke metamorphic belt 257, 282 Ryukyu Arc 299 Sakamoto antiform 281 Salina Cruz 779 Salton Trough 67 Sambagawa, flow patterns during exhumation of metamorphic rocks 279, 293-294 extreme ductile layer, normal thinning and arc-parallel stretching 288-290 exhumation scenario 292-293, 293 normal fault development 289 normal fault stereographs 290 spacial distribution of recrystallized quartz grain sizes 297 strike-slip displacements 290 uniform shear sense and reversal by late-stage folding and faulting 290-291 variable strain geometry of exhuming rocks 292 five possible exhumation mechanisms 280 geological outline 280-282, 287 geological map 282 study results 3D strain geometries 284-288, 284-285, 286 brittle deformation 288 mesoscopic structures 282-284, 283 shear sense distribution 287 San Andreas fault (SAF) 67 absolute plate motion (APM) and relative plate motion (RPM) 59 seismic anisotropy 9-13 Absolute Plate Motion (APM) model 29-30, 29, 30 modelling results 16-31, 77, 78-79, 20-21, 22

Pacific plate viscosity 29 shear wave splitting measurements 10 study methods 13-16, 73 symmetric weak fault model 23-25, 24-26, 24 San Benito fault 67 San Clemente fault 67, 70-73 aeromagnetic anomaly map 77 faults and seismicity in the Santa Barbara Channel area 73 seismic reflection profile 72 San Clemente Island 77 San Diego 68 San Fernando fault zone 75 San Gabriel fault 67, 68, 75 San Gregorio fault 67 San Isidro fault 67 San Jacinto fault 67 San Nicolas Island 77 San Pedro Basin 74 San Pedro Basin fault zone 73-74, 74 San Pedro Bay 74 San Quentin Basin 67 Santa Barbara 68 Santa Barbara Basin 77 Santa Barbara Channel 72-73 faults and seismicity 73 Santa Barbara fault 77 Santa Barbara Island 77 Santa Catalina Island 77 Santa Cruz Basin 67 Santa Cruz Island 77, 73 Santa Cruz-Catalina Ridge 77, 73 Santa Lucia fault 67 Santa Monica Basin 77, 74 Santa Monica Bay 74 Santa Monica-Hollywood fault zone 68 Sardlog shear zone 233 Saruta, River 287, 283, 285, 287 Satukujoq granite 235 seismic anisotropy New Zealand and California 9-13, 31-37, 32-34 applicability of models 36-37 comparison between northern and southern California 35-36 compression 34-35 modelling results 16-31, 77, 78-79, 20-21, 22, 36 study methods 13-16, 73 two-layer anisotropy 12 seismic attenuation 45-46 Sermilagaarsuk 233 Serreta de Guadalupe 255-256, 262 Seto Inland Sea 144, 148, 149, 281 Sevier batholith belt 168 Sevilla 250 shear-wave splitting 3 as a function of wave polarization 23

INDEX California 10, 50 changing depth to isotropy/anisotropy boundary 14 crustal block rotation by mantle flow 90-91 evolution of parameters 22 mantle fabric observations 42-43, 44 New Zealand 10, 50 oceanic material 46-47, 47 San Andreas fault (SAF) 10 Tibet 50 Trinidad 50 Venezuela 50 Shikoku Island 143, 281, 299 Shikoku-western Honshu region, Japan 144 Shimanto metamorphic belt 281, 282 Shimanto Terrane 306 Shimbara Peninsula 299 Shimokawa, River 281, 283, 285, 287 Shiraga, Mount 281 Shuswap extensional complex (SC) 779, 186-187 Sierra Nevada 10 Sierra del Gigante 255-256 Sierra del Maimon 255-256 Sierra del Pericay 255-256, 262 Sierra Larga 255-256 Sierra Madre fault zone 68, 75 Sierra Nevada (SN) batholith belt 169, 777 Sierra San Pedro Martir (SSPM) 772 Sikhote-Alin fault system 143 SKS phases 9-10 Snake River plain (SRP) 779 Snow Peak 184 Solana Formation 260 S0ndre Igaliku 233 S0ndre Sermilik 233, 243 Sorte Nunatak 233 Southern Japan Sea fault zone (SJSFZ) 143 strain gradients 101, 112-114 attachment tectonics 101-103, 102 modelling 103-104, 103, 104 transpression and transtension attachments 107-108, 109-112, 770, 773 transpression attachments 108-109, 777 transtension attachments 109, 772 wrench attachments 104-105, 105-107, 706, 707, 708, 709, 770 foliation and lineation patterns 105 strain modelling 9-13 California applicability of models 36-37 comparison between north and south California 35-36 model results 36 northern California 35 southern California 32-34 New Zealand 31 -32, 32-34 applicability of models 36-37 model results 36 results 16

343

effect of viscosity structure 22-30 other flow models 30-31 relative plate motion with isoviscous model 16-22,77,78-79,20-21,22 study methods 13-14 changing depth to isotropy/anisotropy boundary 14 model parameters 13 relation between deformation parameters and anisotropy 15-16 strike-slip fault systems 1 applicability of models 36-37 block rotation and intracrustal vertical decoupling 141-142, 158, 160 California comparison between north and south California 35-36 northern California 35 southern California 32-34 coupling at boundaries 9-13 modelling results 16-31, 77, 78-79, 20-21, 22 study methods 13-16, 73 model results 36 New Zealand 31-32, 32-34 partitioning 58, 58 termination at convergence zones 65-66, 79-80 bending fault termination 76-78, 76 Pacific-North America tranform plate boundary 66-68, 67, 68 straight fault termination 69-76, 70 termination styles 66, 68-78 vertical coupling and decoupling along WTR boundary 78-79 Tanakura Tectonic Line (TTL) 143 Tanegashima Island 299 Tasermiut 233 tectonic processes, implications of clutch tectonics 53 convergence 54, 51 divergence 54, 56-57 strike-slip partitioning and homogeneous mantle deformation 58, 58 transcurrent boundaries 53-56, 54 Tehuacan 779 Tehuantepec 779 Tehuantepec, Gulf of 118 Tibet absolute plate motion (APM) and relative plate motion (RPM) 50 lithospheric deformation 51 shear-wave splitting 50 Tintina fault (TI) 786-787 Tixtla 722 Tokara Line 299 Tokara Ridge 299, 300 Tokara Sub-basin 299, 300 Toluca 778

344

INDEX

transpressional zones 2 transpressional/transtensional attachment zones, strain gradients 101, 107-108, 109-112, 110, 112-114,113 attachment tectonics 101 -103,702 modelling 103-104, 103, 104 transpression attachments 108 -109, 777 transtension attachments 109, 772 wrench attachments 104-105, 105-107, 106, 107,

108, 109, 110 foliation and lineation patterns 105 transtensional deformation 297, 310 Okinawa-Kyushu junction area evolution model accomodation of extension through reactivation of thrust faults 307 cross-section model 307-310, 309 northern region 307 perpendicular extention 308 Trinidad absolute plate motion (APM) and relative plate motion (RPM) 50, 59 shear-wave splitting 50 Tsuneyama synform 257, 283 Tsushima fault system (TFS) 143 Tuliman 722 Tunua 235 Tuzantian Basin 722 Velez Blanco 255-256, 260 Velez Rubio 255-256 Venezuela absolute plate motion (APM) and relative plate motion (RPM) 50 shear-wave splitting 50 Verdugo fault zone 75 vertical axis rotations 3-4, 83-85, 84, 97-98 assumptions and applicability of model 97 Betic Cordillera 249-250, 273-275 evolution of structures and implications for coupling and decoupling 273, 274 geological map 250, 255-256, 260 geological setting 250-253, 252 pattern of vertical-axis rotations 272-273 regional constraints 271-272 rock successions 253-259, 254 structures 257-255, 259-271, 267, 262, 264, 265, 270 rigid rotations application to attachment/detachment zones 92 boundary conditions and background 85-87, 55, 56,57 experimental apparatus and design 87-89, 55, 59 experimental results 89, 90, 91 mantle deformation in obliquely convergent environments 91-92 natural systems 89-92 shear-wave splitting 90-91

side-driven v. bottom-driven systems 95-97 upper crustal rotation coinciding with mantle deformation 92, 94, 95 Eastern Transverse Ranges, California 92-94, 93, 94 Marlborough fault system, South Island, New Zealand 94-95, 96 Western Transverse Ranges (WTR), California 69 vertical coupling in the lithosphere channel flow 5 geophysical constraints shear wave splitting 3 vertical axis rotations 3-4 island arcs and marginal basins 4-5 orogenic belts and exposed attachment zones 4 viscosity and strain modelling effect of viscosity structure 22-23 Absolute Plate Motion (APM) model 29-30, 29,

30

asymmetric viscosity models 25-28, 27-25 effect of increasing compression component 28-29, 29 symmetric weak fault model 23-25, 24-26, 24 isoviscous model 16-22, 77, 75-79, 20-21, 22 Vizcaino Peninsula 67 western Idaho shear zone (WISZ) 779, 186-187 Western Metamorphic Belt (WMB) 777 western Nevada shear zone (WNSZ) 186-187 Western Transverse Ranges (WTR), California 66-69, 67, 79-80 clockwise vertical-axis rotation mechanisms 69 strike-slip fault termination styles 68-69 bending fault termination 76-78, 76 straight fault termination 69-76, 70 vertical coupling and decoupling along boundary 78 basal shear-driven block rotation 79 edge-driven block rotation 78-79 Whittier-Elsinore fault zone 77-78 Wind River Mountains (WR) 779 wrench attachments 104-105 foliation and lineation patterns 105 strain gradients 105-107, 106, 107, 108, 109, 110 Xochipala 722 Yakushi antiform 257, 253 Yakushima Island 299 Yamaguchi City 745, 749, 757 Yangsan fault system (YFS) 743 Yanhuitlan 775 Yoshiwa 746 Zarcilla de Ramos 255-256 Zarcilla de Ramos Basin 255 Zihuatanejo 775 Zitlala 722 Zumpango 722