Flow Processes in Faults And Shear Zones (Geological Society Special Publication No. 224)

Flow Processes in Faults and Shear Zones Geological Society Special Publications Society Book Editors R. J. PANKHURST...

Author: G. I. Alsop | R. E. Holdsworth | K. J. W. McCaffrey | M. Hand

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Flow Processes in Faults and Shear Zones

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 I 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: ALSOP, G. L, HOLDSWORTH, R. E., MCCAFFREY, K. J. W. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224. CASCIELLO, E., CESARANO, M. & COSGROVE, J. 2004. Shear deformation of pelitic rocks in a largescale natural fault. In: ALSOP, G. L, HOLDSWORTH, R. E., MCCAFFREY, K. J. W. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224,113-125.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 224

Flow Processes in Faults and Shear Zones

EDITED BY

G. I. ALSOP

University of St Andrews, UK

R. E. HOLDSWORTH University of Durham, UK

K. j. w. MCCAFFREY University of Durham, UK and

M. HAND University of Adelaide, Australia

2004 Published by The Geological Society London

THE GEOLOGICAL SOCIETY

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Contents ALSOP, G. I. & HOLDSWORTH, R. E. Shear zones - an introduction and overview

1

DIJKSTRA, A. H., DRURY, M. R., VISSERS, R. L. M., NEWMAN, J. & VAN ROERMUND, H. L. M. Shear zones in the upper mantle: evidence from alpine- and ophiolite-type peridotite massifs

11

WHITE, J. C. Instability and localization of deformation in lower crust granulites, Minas fault zone, Nova Scotia, Canada

25

LLOYD, G. E. Microstructural evolution in a mylonitic quartz simple shear zone: the significant roles of dauphine twinning and misorientation

39

PIAZOLO, S., ALSOP, G. L, M0LLER NIELSEN, B. & VAN GOOL, J. A. M. The application of GIS to unravel patterns of deformation in high grade terrains: a case study of indentor tectonics from west Greenland

63

VIGNERESSE, J. L. Rheology of a two-phase material with applications to partially molten rocks, plastic deformation and saturated soils

79

COLLETTINI, C. & BARCHI, M. R. A comparison of structural data and seismic images for low-angle normal faults in the Northern Apennines (Central Italy): constraints on activity

95

CASCIELLO, E., CESARANO, M. & COSGROVE, J. Shear deformation of pelitic rocks in a large-scale natural fault

113

MALTMAN, A. & VANNUCCHI, P. Insights from the Ocean Drilling Program on shear and fluid-flow at the mega-faults between actively converging plates

127

JANSSEN, C, LUDERS, V. & HOFFMANN-ROTHE, A. Contrasting styles of fluid-rock 141 interaction within the West Fissure Zone in northern Chile GUPTA, S. & BICKLE, M. J. Ductile shearing, hydrous fluid channelling and high-pressure metamorphism along the basement-cover contact on Sikinos, Cyclades, Greece

161

ALSOP, G. I. & HOLDSWORTH, R. E. Shear zone folds: records of flow perturbation or structural inheritance?

177

MAC!NNES, E. A. & WHITE, J. C. Geometric and kinematic analysis of a transpression terrane boundary: Minas fault system, Nova Scotia, Canada

201

PEREIRA, M. F. & SILVA, J. B. Development of local orthorhombic fabrics within a simple-shear dominated sinistral transpression zone: the Arronches sheared gneisses (Iberian Massif, Portugal)

215

OCCHIPINTI, S. A. & REDDY, S. M. Deformation in a complex crustal-scale shear zone: Errabiddy Shear Zone, Western Australia

229

BAILEY, C. M., FRANCIS, B. E. & FAHRNEY, E. E. Strain and vorticity analysis of transpressional high-strain zones from the Virginia Piedmont, USA

249

GIORGIS, S. & TIKOFF, B. Constraints on kinematics and strain from feldspar porphyroclast populations

265

GUMIAUX, C., BRUN, J. P. & GAPAIS, D. Strain removal within the Hercynian Shear Belt of Central Brittany (western France): methodology and tectonic implications

287

vi

CONTENTS

CARRERAS, J., DRUGUET, E., GRIERA, A. & SOLDEVILA, J. Strain and deformation history in a syntectonic pluton. The case of the Roses granodiorite (Cap de Creus, Eastern Pyrenees)

307

MOLLI, G. & TRIBUZIO, R. Shear zones and metamorphic signature of subducted continental crust as tracers of the evolution of the Corsica/Northern Apennine orogenic system

321

CHEW, D. M., DALY, J. S., FLOWERDEW, M. I, KENNEDY, M. J. & PAGE, L. M. Crenulation-slip development in a Caledonian shear zone in NW Ireland: evidence for a multi-stage movement history

337

MAZZOLI, S., INVERNIZZI, G, MARCHEGIANI, L., MATTIONI, L. & CELLO, G. Brittle-ductile shear zone evolution and fault initiation in limestones, Monte Cugnone (Lucania), southern Apennines, Italy

353

Index

375

Referees The editors are very grateful to the following people for their help in refereeing papers for this volume. C. M. Bailey, College of William and Mary, USA W. Bailey, CSIRO Petroleum, Australia M. Barchi, Universita di Perugia, Italy M. Bestmann, Liverpool University, UK M. Bickle, Cambridge University, UK M. Brown, University of Maryland, USA J. Carreras, Universitat Autonoma de Barcelona, Spain C. Ehlers, Abo Akademi University, Finland D. Gapais, Universite de Rennes 1, France J. Imber, University of Durham, UK R. Jamieson, Dalhousie University, Canada D. Jiang, University of Maryland, USA S. Johnson, University of Maine, USA R. R. Jones, CognIT, Norway M. Krabbendam, British Geological Survey, UK R. Law, Virginia Tech, USA S. Lin, University of Waterloo, Canada G. E. Lloyd, University of Leeds, UK A. Maltman, Aberystwyth University, UK A. McCaig, University of Leeds, UK B. Miller, San Jose State University, USA T. Needham, University of Leeds, UK G. J. H. Oliver, University of St Andrews, UK C. W. Passchier, Johannes Gutenberg Universitat, Germany

N. Petford, Kingston University, UK G. Potts, Liverpool University, UK G. Roberts, University College London, UK A. Robertson, Edinburgh University, UK C. Rosenberg, Freie Universitat, Germany M. Sandiford, University of Melbourne, Australia S. Sherlock, The Open University, UK R. Sibson, University of Otago, New Zealand C. Simpson, Boston University, USA C. J. Spiers, Utrecht University, The Netherlands R. Strachan, Portsmouth University, UK E. Tavarnelli, Universita di Siena, Italy B. Tikoff, University of Wisconsin-Madison, USA S. Treagus, University of Manchester, UK J. Turner, University of Birmingham, UK A. Vauchez, Universite de Montpellier II, France J.-L. Vigneresse, CREGU, France R. Weinberg, University of Western Australia, Australia M. Wells, University of Las Vegas, USA B. Yardley, University of Leeds, UK

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Shear zones - an introduction and overview G. I. ALSOP1 & R. E. HOLDSWORTH2 Crustal Geodynamics Group, School of Geography &. Geosciences, University of St Andrews, St Andrews, Fife, Scotland, KY16 9AL, UK (e-mail [email protected]) ^Reactivation Research Group, Department of'Earth Sciences, University of Durham, Durham, DH1 3LE, UK l

Faults and their deeper-level equivalents, shear zones are localized regions of higher strain which effectively accommodate differential movement in the Earth's crust and mantle during deformation of the lithosphere. Shear zones may be more precisely defined as approximately tabular regions of concentrated deformation and flow across which adjacent relatively undeformed rock units are offset. They are recognized at all sizes from micro to plate boundary scale (Ramsay 1980; Sornette et al. 1990) (Figs 1 & 2). Faults and shear zones are therefore important examples of the heterogeneous nature of deformation in natural rocks, and profoundly influence the location, architecture and evolution of a broad range of geological phenomena (e.g. Rutter et al. 2001). The topography and bathymetry of the Earth's surface is marked by mountain belts and sedimentary basins which are controlled by faults and shear zones. In addition, faults and shear zones control fluid migration and transport, including hydrothermal fluids and hydrocarbons of economic significance (e.g. McCaig 1997). Magma transport, emplacement and eruption are also frequently controlled by faults and shear zones, as are earthquakes. Once faults and shear zones are established, they are often long-lived features prone to multiple reactivation over very large time-scales (e.g. Holdsworth et al. 1997). Faults and shear zones are typically arranged into complex interlinked networks that permit 3D strain in response to plate tectonic forces (Dewey et al. 1986). However, analysis of ductile shear zones is complicated as they are only directly accessible to geoscientists after exhumation to the Earth's surface. In such cases, the relationships between the observed finite deformation patterns, the preserved microstructures at any given location, and the deformation path and strain rate history are potentially difficult to resolve (e.g. Knipe 1989). It is therefore necessary to first consider the bulk deformation behaviour of the lithosphere and the nature and strength of deformed rocks at depth within shear zones.

Fig. 1. Banded orthogneiss with darker amphibolite layers displaying dextral offset across minor shear zones in West Greenland. Note the attenuation of layering, variable displacement and minor melt component along the shears. Pocket knife for scale.

Strength, strain-rate histories and fault rocks at depth Laboratory-based rock deformation experiments, together with geophysical studies and field observations of natural examples suggest that fault and shear zone deformation products, processes and rheology change with depth

From: ALSOP, G. L, HOLDSWORTH, R. E., MCCAFFREY, K. J. W. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224,1-9. 0305-8719/$15.00 © The Geological Society of London 2004.

Fig. 2. Aeromagnetic map of part of the 180 km long Nordre Str0mfjord shear zone in the Nagssugtoqidian belt of West Greenland. Magnetic patterns clearly define the ENE-WSW-trending shear zones which are marked by pronounced swings in the trend of the regional magnetic signature (see Rasmussen & Van Gool 2000 and references therein).

INTRODUCTION TO SHEAR ZONES

Fig. 3. General graphs (partly modified from Knipe 1989) and schematic sketches illustrating variations in shear resistance, deformation regime, fault rocks and strain rate with depth in the lithosphere. The lower diagram represents a schematic profile through crustal faults and shear zones illustrating the frictional to viscous regime together with the typical fault rocks and products of deformation.

(Fig. 3; Sibson 1977, 1983). For a crustal-scale structure, most models suggest that an upper crustal network of brittle faults and cataclastic deformation products will connect directly at depth with a generally broader system of anastomosing shear zones where deformation products are mylonitic in character (Fig. 3). In the upper crustal region, deformation mechanisms involve brittle failure and frictional sliding, with fault

3

strength generally rising with increasing depth due to increasing effective pressure (Fig. 3) (Byerlee 1978; Paterson 1978; Sibson 1983). This frictional deformation regime is often seismogenic. In the shear zones at greater depths, the regime changes to one of viscous flow where deformation is generally considered to be aseismic with dominant operative mechanisms exemplified by crystal plasticity and diffusional creep (Fig. 3) (Sibson 1977; Tullis & Yund 1977; Schmid & Handy 1991). Here, the main controls on strength are temperature, strain rate and grain size. Due to increasing temperature, strength is believed to decrease with increasing depth. The transition between frictional and viscous regimes is typically thought to lie at depths between 10 to 15 km in continental regions and to coincide with the strength maxima of the crust and - in some cases - perhaps the lithosphere (Figs 3 & 4) (Sibson 1983). The character and deformation behaviour within this transition zone is complicated by the pre-existing compositional heterogeneity of rocks on different scales (e.g. Handy 1990) and by changes in composition, deformation mechanisms and rheology that occur during fault rock deformation and associated, often retrograde, metamorphism (e.g. Schmid & Handy 1991; Wintsch et al 1995; Holds worth et al 2001). A number of recent experimental and field-based studies (e.g. Bos & Spiers 2001,2002; Stewart et al 2000; Imber et al 1997) have highlighted the importance of fluidrock interactions and how these may lead to marked fault zone weakening in the region of the frictional viscous transition along crustalscale faults.

4

G. I. ALSOP & R. E. HOLDSWORTH

plete understanding of fault and shear zone processes and their controls, it is necessary to consider the relationships and interactions of processes across all scales.

Lithospheric-scale controls

Fig. 4. Schematic strength profile through the crust and upper mantle illustrating the principal loadbearing regions of the lithosphere. The strength of the crust and mantle are based on quartz-feldspar and olivine rheologies respectively.

Based on what we know regarding strain-rate evolution under different deformation regimes, it seems likely that strain-rate histories will also show fundamental variations with increasing depth (Fig. 3) (Knipe 1989). Although the shallow frictional regime is dominated by cataclasites associated with punctuated high strainrate episodes, the frictional-viscous transition is marked by a greater component of aseismic flow between such short-lived high strain-rate events. The deeper viscous regime is dominated by ductile mylonites in which deformation is accommodated by more continuous and relatively steady flow (Fig. 3). Once again, the strainrate behaviour in the main load-bearing frictional-viscous transition zone will be of considerable importance, with the possibility that periods of unstable viscous flow in mylonites could trigger seismic episodes in the overlying brittle crust (Figs 3 & 4) (e.g. Hobbs et al 1986).

Lithospheric-scale controls on faults and shear zones include the general tectonic regime (shortening, extension and strike-slip) and plate boundary conditions (e.g. Sibson 1983; Teyssier & Tikoff 1998). The thickness of the quartzofeldspathic crust and age of the (mantle) lithosphere, which reflects its thermal history, also provides large-scale controls on lithosphere deformation and therefore fault and shear zone development (e.g. see Tommasi et al. 1995). Experimental data suggest that the primary load-bearing region of the lithosphere lies in the upper mantle, whereas a secondary load-bearing region is developed at the frictional-viscous transition in the crust (Fig. 4). The general view then is that large-scale deformation response is determined by development of shear zones in the upper mantle (e.g. Molnar 1992; RegenauerLieb & Yuen 2004). If the crust and mantle are coupled, then flow within the mantle will ultimately control deformation in the crust (e.g. Tikoff etal. 2002).

Grain-scale controls Grain-scale controls on faults and shear zones include lithological controls such as the composition, mineralogy and grain size of the host, together with the presence of pre-existing fabrics and introduction of a fluid or melt. Local environmental conditions of pressure, temperature and strain rate will also affect faulting and shear zone processes. The influx of hydrous fluids into active deformation zones at depth appears to be particularly important and will lead to profound changes in the chemistry and mineralogy of the host, grain size and dominant deformation mechanisms and ultimately therefore the rheology and strength of the fault/shear zone rock (e.g. Bos & Spiers 2002; Holdsworth et al. 2001).

Controls on fault and shear zone development

Network geometry-scale processes

Three sets of controls on fault and shear zone development exist: lithospheric-scale, network geometry-scale, and grain-scale. Each scale range tends to be the focus of different research communities, e.g. geodynamicists, structural geologists and microstructural geologists, respectively. However, in order to gain a com-

Geometric controls are central to the transfer of displacement between mantle shear zones driving lithosphere deformation and overlying systems of crustal faults and shear zones (e.g. see Tommasi et al. 1995; Tikoff et al. 2002). More fundamentally, they form the key mechanical link between grain-scale and lithospheric-scale


5

Fig. 5. Schematic 3D sketch illustrating shear zones anastomosing around low-strain augen. The relatively small amount of mechanically weak but interconnected shear zone rock forms a network which effectively controls the bulk strength of the overall rock volume.

Fig. 6. Diffuse zone of shearing between quartzofeldspathic orthogneiss (right-hand side of photograph) and amphibolite-rich gneisses (left-hand side of photograph) in West Greenland. Note how variable rheology results in boudinage and folding of amphibolites. Hammer for scale (top centre).

processes. Such controls include the size and interconnectivity of adjacent shear zones, the orientation and dip of shear zones in relation to far-field stress, and displacement compatibility. In general, larger and more interconnected shear zones will require smaller volumes of mechanically weak fault rocks in order to form an interconnected weak network that controls the overall strength of the system on crustal or lithospheric scales (Fig. 5). Relatively low strain units may be entirely surrounded by high strain (and weak) interconnected shear zones which effectively mechanically isolate these regions of relatively strong rock as augen (Fig. 5). Displacement compatibility also needs to be considered in relation to how stick-slip faulting at shallow depths relates to steady-state flow at depth. In order to understand faults and shear zones better, we must consider the interaction of all of these factors on lithosphere, grain and networkscale controls. We shall now consider some of the main points arising from this volume.

timescales (e.g. see Tommasi etal. 1995; Vauchez & Tommasi 2003). The present volume opens with seven papers which discuss these topics from deeper to shallower levels using a variety of techniques and scales of observation. Dijkstra et al. describe localization mechanisms of deformation in alpine- and ophiolite-type mantle massifs. They suggest that softening processes responsible for localization are associated with melt-related weakening, together with a change in the dominant deformation mechanism from dislocation to grain size-sensitive creep. Such reactions occur over a broad range of P-T conditions in the upper mantle, suggesting that mantle shear zones are widespread and may significantly reduce the bulk strength of the lithosphere. White describes localization of deformation in lower crust granulites in Nova Scotia. Localization may occur at several different length scales resulting in finer-grained material associated with dislocation creep microstructures. The finest-grained mylonites are marked by the introduction of partial melts and pseudotachylites associated with transient frictional events. Microstructures which formed during an event which initiated an instability can be obliterated by subsequent ductile flow. The role of microstructure in the evolution of an amphibolite facies shear zone is described by Lloyd. This contribution explores the role twinning plays in both the microstructural and petrofabric evolution of shear zones by assisting in the initial grain size comminution processes. Twinning may thus help to accommodate high shear strains whilst maintaining a stable microstructure and constant (single crystal) petrofabric

Lithosphere deformation and rheology of shear zones As noted above, shear zones processes and products change with depth and the type of material being deformed (Fig. 6). The nature of these processes and how they may be linked kinematically and mechanically, particularly across major geological and geophysical interfaces such as the Moho, will profoundly affect the overall rheological behaviour of the lithosphere during deformation on geological

6


which may be detected by seismic investigation of deep shear zones. The testing of remote geophysical techniques via the application of Geographical Information Systems (GIS) to granulite and amphibolite facies shear zones in West Greenland is presented by Piazolo et al. Multidisciplinary geophysical and geological datasets are compared in terms of subject, space and scale and suggest that Theologically weak amphibolite facies rocks are moulded and sheared around a cooled granulite facies block which effectively behaves as a rigid indentor on a regional scale. The rheology of deforming material is also addressed by Vigneresse in terms of two phase materials such as crystallizing magma. The general behaviour of viscosity is taken as a function of the strain rate and the amount of solid phase present. At high strain rates, the viscosity contrast between the two phases at its lowest, whilst at low strain rates, the viscosity contrast between the phases is at its highest. These relationships will clearly control the structures that are developed. The role of large-scale extensional faults in crustal deformation in the Northern Appenines of Central Italy is studied by Colletini & Barchi. Recorded microseismicity suggests that faults are presently active with a vertical o^ and that fluid movement along the gently-dipping fault planes has resulted in the faults being mechanically weak. Casciello et al. also investigate the role of fluids in upper crustal faults via analysis of the hydraulic characteristics and hence frictional properties of sheared clays. Shear strain can induce mineralogical changes in smectite resulting in their replacement with anhydrous illite minerals. Localization of the illitization process along the shear zone may generate water, leading to fluid overpressure and hydraulic circulation. Partitioning processes in shear zones The central section of the book contains a series of eight papers on partitioning and localization processes, which are an increasingly recognized phenomena in shear zones across a range of scales from microstructural to outcrop to regional (Fig. 7). Partitioning may be analysed in terms of the role of fluid flow, together with the localization and focusing of fluids along faults and shear zones. Mega-faults between actively converging plates are described by Maltman & Vannucchi. Lithological influence on localization of fault propagation seems absent with faults responding to continuing deformation by intensifying and focusing strain inwards rather than propagating outward splays. Faults penetrated by the Ocean Drilling Programme may be

Fig. 7. Outcrop-scale dextral shear zone developed in pelite at Cap de Creus, Spain (see Carreras 2001 for further details). Regional deformation is focused into the retrogressive shear zone, which further partitions and localizes deformation. Note the sheared pegmatite and pronounced strike-swing of regional foliation into the shear zone.

just tens of metres in thickness and yet can efficiently channel fluids for distances of tens of kilometres whilst at the same time inhibiting fluid flow across them. Styles of fluid rock interaction within the West Fissure Zone in Chile are examined by Janssen et al. Their results demonstrate considerable variation in the degree of fluid interaction with limited ascending hydrothermal fluids in some traverses, whereas in other cases fluid-enhanced weakening mechanisms are dominant, reflecting the heterogeneity and partitioning of fluid flow in largescale continental fault zones. The partitioning and localization of fluid flow in the Cyclades of Greece is described by Gupta & Bickle. Their work shows an association of higher strain with increased hydration in the footwall to a basement-cover shear zone along which fluids have been focused. The restricted availability of water outside of this zone allows the preservation of earlier metamorphic assemblages which are otherwise destroyed. Partitioning processes in shear zones can also be investigated in terms of deformation


localization and partitioning, which results in the multiple generation of a range of structures displaying overprinting relationships at a variety of scales. Alsop & Holdsworth describe deformation within shear zones in which folding displays predictable geometric patterns analysed on fabric topology plots. Multiple generations of folds may be explained in terms of a fold evolution model, where early sheath folds represent a more highly deformed and evolved variety of synshearing folds originally generated during perturbations in ductile flow. Alternatively, the fold inheritance model suggests that the structural architecture generated during sheath folding may subsequently control the geometry and govern the orientation of later synshearing folds. Machines & White describe the largescale partitioning of deformation into distinct external (contractional structures) and internal (shear) zones along the Minas Fault System of Nova Scotia. Ductile structures are composite features derived from the multiple transposition of pre-existing layers, with strain accommodated within progressively narrower volumes of rock marked by contrasting deformation micromechanisms. Pereira & Silva investigate the partitioning of deformation in a transpressive sinistral shear zone in the Iberian Massif of Portugal. Although the shear zone is dominated by widespread monoclinic fabrics, mechanical conditions locally exist in which deformation has been partitioned into orthorhombic fabrics highlighting the importance of scale when investigating deformation processes. Occhipinti & Reddy undertake a detailed investigation of the crustal-scale Errabiddy Shear Zone of Western Australia which shows a temporal evolution from generation of ductile fabrics, to folding to a brittle overprint. Deformation is locally partitioned into simple shear dominated displacement zones that separate regions of flattening and pure shear. Analyses of small segments of shear zones may thus not give the complete history of an evolving shear zone because of strain localization and partitioning over time. Bailey et al. analyse transpressional high strain zones from the Virginia Piedmont of the southern Appalachians. Estimates of strain based on quartz grain shapes, porphyroclasts and folded/boudinaged pegmatites suggest significant strike-slip displacement is associated with orogen-parallel material elongation together with 40-70% contraction normal to this zone.

Unravelling shear zone histories Deformation associated with faults and shear zones may be both spatially and temporally variable resulting in a complex interplay of

7

Fig. 8. Minor aplite intrusion rotated and attenuated by sinistral shears formed within the Roses Granodiorite (Cap de Creus, Spain). Minor intrusions may provide effective markers in the unravelling of shear zone histories (see paper by Carreras et al.).

structures developed across a range of scales. This final section of the book comprises six papers concerned with unravelling shear zone histories through a variety of techniques and approaches. Estimates of strain, metamorphic history, isotopic age dating, fluid influx and minor igneous intrusions acting as 'timemarkers' may all be particularly useful in interpreting and unravelling such histories (Fig. 8). Giorgis & Tikoff constrain the kinematics and finite strain of deformation in shear zones based on 3D numerical models of the rotation populations of rigid clasts. This numerical model suggests that there is no consistent relationship between the asymmetrical orientation of a population of rigid markers and the simple shear component of deformation. This suggests that the asymmetrical alignment of a population of porphyroclasts is not a reliable shear sense indicator. They apply the technique to the Western Idaho Shear Zone and suggest significant oblique convergence during transpressional deformation. Gumiaux et al. undertake a statistical analysis of cleavage orientation data in

8


order to restore major deformation associated with the Hercynian shear belt of central Brittany. Strain patterns reflect domains on different scales and the occurrence of local superimposed deformations highlighting spatial and temporal shear zone evolution on a variety of scales. Carreras et al. investigate the strain and deformation history of a syntectonic pluton in the Eastern Pyrenees. This pluton displays a continuous deformation history from magmatic structures developed during cooling to latestage (solid state) mylonitic fabrics developed along shear zones when the pluton had crystallized. Quantitative strain analysis of magmatic and solid state fabrics suggest that both record bulk finite strains of similar magnitudes, but with marked differences in the distribution of strain which becomes increasingly localized during solid state deformation. Molli & Tribuzio examine the metamorphic and structural history of major shear zones in the Tenda massif of Corsica. Contractional shear zones display a HP/LT metamorphic signature which is overprinted by retrogressive (greenschist facies) exhumation-related extensional structures indicating a protracted metamorphic and structural history. Chew et al. also describe a protracted structural evolution within Caledonian shear zones of NW Ireland. Early recumbent fold nappes are overprinted by dextral strike-slip crenulations which are followed by a subsequent phase of sinistral deformation. Major crustal scale shear zones may thus be reactivated over significant time scales. Mazzoli et al. investigate the structural evolution of brittle-ductile shear zones developed in low grade settings of the Southern Appenines of Italy. Analysis of shear zone related vein arrays suggests that they are controlled by displacement accumulation and shear strain developed at approximately constant temperature. Detailed analysis of such vein fills may therefore provide important constraints on the factors which control fluids and fault development during brittle-ductile deformation. Summary Lithospheric-scale controls on the development of faults and shear zones are clearly important with current work concentrating on crust-mantle linkages, together with the attributes and processes in mantle shear zones. Improved remote sensing of deep shear zones will aid in the interpretation of large-scale deformation localization and magma transport. Grain-scale controls relate to deformation mechanisms and recognition of diffusive creep

in shear zones associated with studies of microrheology. Further work may concentrate on chemical, strain and energy/heat partitioning processes in polymineralic rocks which will undoubtedly be aided by modelling flow and grain-scale deformation mechanisms. Geometric controls on faults and shear zones are reflected in their kinematic and flow characteristics, which may be marked by patterns of vorticity and spin together with strain partitioning and localization. Future work may in general attempt to quantify and model fault and shear zones processes across this range of controls. This may involve the development of new analytical and geochronological techniques, together with the application and integration of analogue/numerical modelling to studies of natural faults and shear zones. Thus, although significant advances have undoubtedly been made in the understanding of fault and shear zone processes, new work may attempt to integrate studies further encompassing a range of techniques varying from fieldbased to remote geophysical, and a range of scales from lithospheric to microscopic. Such a mixed, multidisciplinary approach will provide a more holistic view of fault and shear zone development and the controls and evolution of lithospheric deformation. Most of the papers in this volume were presented at a joint international meeting of the Tectonic Studies Group of the Geological Society, London, the Structural Geology and Tectonics Division of the Geological Society of America and the Geological Society of Australia on the 2-3 September 2002. We would like to thank all contributors, and participants at the Transport and Flow Processes in Shear Zones meeting and field trip. We would also like to thank Lorna Stewart for her help with this Special Publication.

References Bos, B. & SPIERS, C.J. 2001. Experimental investigation into the microstructural and mechanical evolution of phyllosilicate-bearing fault rock under conditions favouring pressure solution. Journal of Structural Geology, 23,1187-1202. Bos, B. & SPIERS, CJ. 2002. Frictional-viscous flow of phyllosilicate-bearing fault rock: Microphysical model and implications for crustal strength profiles. Journal of Geophysical Research, 107, (B2). BYERLEE, J.D. 1978. Friction of rocks. Pure and Applied Geophysics, 116, 615-626. CARRERAS, J. 2001. Zooming on Northern Cap de Creus shear zones. Journal of Structural Geology, 23,1457-1486. DEWEY, J.F., HEMPTON, M.R., KIDD, W.S.F., SAROGLU, F. & SENGOR,A.M.C. 1986. Shortening of continental lithosphere: the neotectonics of Eastern Anatolia- a young collision zone. In: COWARD, M.P &

INTRODUCTION TO SHEAR ZONES RIES, A.C (eds) Collision tectonics. Geological Society, London, Special Publications, 19, 3-36. HANDY, M.R. 1990. The solid-state flow of polymineralic rocks. Journal of Geophysical Research, 95, 8647-8661. HOBBS, B.E., ORD, A. & TEYSSIER, C. 1986. Earthquakes in the ductile regime. Pure and Applied Geophysics, 124, 310-336. HOLDSWORTH, R.E., BUTLER, CA. & ROBERTS, A.M.

1997. The recognition of reactivation during continental deformation. Journal of the Geological Society, London, 154, 73-78. HOLDSWORTH, R.E., STEWART, M., IMBER, J. & STRACHAN, R.A. 2001. The structure and rheological evolution of reactivated continental fault zones: a review and case study. In: MILLER, J.A., HOLDSWORTH, R.E., BUICK, I.S. & HAND, M. (eds) Continental reactivation and reworking. Geological Society, London, Special Publications, 184, 115-137. IMBER, I, HOLDSWORTH, R.E., BUTLER, C.A. & LLOYD, G.E. 1997. Fault-zone weakening processes along the reactivated Outer Hebrides Fault Zone, Scotland. Journal of the Geological Society, London, 154,105-110. KNIFE, R.J. 1989. Deformation mechanisms - recognition from natural tectonites. Journal of Structural Geology, 11,127-146. McCAiG, A.M. 1997. The geochemistry of volatile fluid flow in shear zones. In: HOLNESS, M. (ed.) Deformation enhanced melt segregation and metamorphic fluid transport. Chapman and Hall, 227-260. 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. PATERSON, M.S. 1978. Experimental rock deformation - the brittlefield.Springer-Verlag, Berlin. RAMSAY, J.G. 1980. Shear zone geometry: a review. Journal of Structural Geology, 2, 83-99. RASMUSSEN, T.M. & VAN GOOL, J.A.M. 2000. Aeromagnetic survey in southern West Greenland: project Aeromag 1999. Geology of Greenland Survey Bulletin, 186, 73-77. REGENAUER-LIEB, K. & YUEN, D.A. 2004. Modeling shear zones in geological and planetary sciences: solid- andfluid-thermal- mechanical approaches. Earth Science Reviews, 63, 295-349. RUTTER, E.H., HOLDSWORTH, R.E. & KNIPE, R.J. 2001. The nature and tectonic significance of fault zone weakening: an introduction. In: HOLDSWORTH, R.E., STRACHAN, R.A., MAGLOUGHLIN, J.F. & KNIPE, R.J. (eds) The nature and tectonic significance of fault zone weakening. Geological Society, London, Special Publications, 186,1-11.

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SCHMID, S.M. & HANDY, M.R. 1991. Towards a genetic classification of fault rocks: geological usage and tectonophysical implications. In: MULLER, D.W., MCKENZIE, J.A. & WEISSERT, H. (eds) Controversies in modern geology: evolution of geological theories in sedimentology, earth history and tectonics. Academic, San Diego, California, USA, 339-361. SIBSON, R.H. 1977. Fault rocks and fault mechanisms. Journal of Geological Society, London, 133, 191-213. SIBSON, R.H. 1983. Continental fault structure and the shallow earthquake source. Journal of Geological Society of London, 140, 741-767. SORNETTE, D., DAVY, P. & SORNETTE, A. 1990. Structuration of the lithosphere in plate-tectonics as a self-organized critical phenomenon. Journal of Geophysical Research, 95,17353-17361. STEWART, M., HOLDSWORTH, R.E. & STRACHAN, R.A. 2000. Deformation processes and weakening mechanisms within the frictional-viscous transition zone of major crustal-scale faults: insights from the Great Glen Fault Zone, Scotland. Journal of Structural Geology, 22, 543-560. TEYSSIER, C. & TIKOFF, B. 1998. Strike-slip partitioned transpression of the San Andreas Fault system: a lithospheric scale approach. In: HOLDSWORTH, R.E., STRACHAN, R.A. & DEWEY, J.F. (eds) Continental transpressional and transtensional tectonics. Geological Society, London, Special Publications, 135,143-158. TIKOFF, B., TEYSSIER, C. & WATERS, C. 2002. Clutch tectonics and the partial attachment of lithospheric layers. In: BERTOTTI, G, SCHULMANN, K. & CLOETINGH, S.A.P.L. (eds) Continental Collision and the Tectono-sedimentary Evolution of Forelands. EUG Stephan Mueller Special Publication Series, 1, 57-73. TOMMASI, A., VAUCHEZ, A. & DAUDRE, B. 1995. Initiation and propogation of shear zones in a heterogeneous continental lithosphere. Journal of Geophysical Research, 100, 22083-22101. TULLIS, J. & YUND, R.A. 1977. The brittle-ductile transition in feldspathic rocks. Transactions American Geophysical Union, 68,1464. VAUCHEZ, A. & TOMMASI, A. 2003. Wrench faults down to the asphenosphere: geological and geophysical evidence and thermo-mechanical effects. In: STORTI, F, HOLDSWORTH, R.E. & SALVINI,F. (eds) Intraplate strike-slip deformation belts. Geological Society, London, Special Publications, 210,15-34. WINTSCH, R.P, CHRISTOFFERSON, R. & KRONENBERG, A.K. 1995. Fluid-rock reaction weakening of fault zones. Journal of Geophysical Research, 100, 13021-13032.

Shear zones in the upper mantle: evidence from alpine- and ophiolite-type peridotite massifs ARJAN H. DIJKSTRA1, MARTYN R. DRURY2, REINOUD L. M. VISSERS2, JULIE NEWMAN3 & HERMAN L. M. VAN ROERMUND2 1 Department of Applied Geology, Curtin University of Technology, Perth, Western Australia Postal address: GPO Box U1987, Perth, WA 6845, Australia (e-mail: a. dijkstra@curtin. edu. au) 2 Vening Meinesz School of Geodynamics, Utrecht University, Utrecht, The Netherlands 3 Department of Geology and Geophysics, Texas A &M University, College Station, Texas,

USA Abstract: There is abundant field and microstructural evidence for localization of deformation in alpine- and ophiolite-type mantle massifs. On the basis of field relationships and microstructures we recognize two types of tectonite shear zones (medium- to coarse- and fine-grained), as well as two types of mylonitic shear zones (anhydrous and hydrous peridotite mylonites). In tectonite shear zones, softening processes responsible for localization are probably melt-related weakening in the medium to coarse tectonites and a change in limiting slip system in the fine-grained tectonites. In peridotite mylonites, the most likely cause for softening and localization is a change in dominant deformation mechanism from dislocation to grain size sensitive creep. Microstructural and petrological study of mylonite rocks reveals that reactions, either continuous net-transfer reactions (anhydrous and hydrous) or melt-rock reactions, play a key role in the formation of fine-grained material that promotes grain size sensitive creep. These reactions occur over a broad range of pressure-temperature conditions encompassing a large part of the lithospheric upper mantle. We conclude that mantle shear zones are widespread and that they reduce the (bulk) strength of the lithosphere significantly.

Attempts to understand the dynamics of tectonic processes depend critically on quantitative knowledge concerning the mechanical properties of the lithosphere and the relative strength of the crust and upper mantle (e.g. Brun 2002). For many years, it has been accepted that the upper part of the mantle lithosphere is strong compared to the crust (e.g. Kirby 1983; Kuznir & Park 1986). The notion of a strong uppermost mantle is based on extrapolation of laboratory flow laws (Kohlstedt et al 1995) and the depth distribution of seismicity in the lithosphere (Chen & Molnar 1983). The inferred high strength of the shallow mantle implicitly refers to homogeneous flow. However, at large strains, deformation of the mantle lithosphere may become localized within weak shear zones (Fig. 1) so the application of low-strain flow laws may seriously overestimate the strength of the mantle (Rutter & Brodie 1988; Handy 1989; Vissers et al. 1995). In a recent study on the seismicity and elastic thickness of continental lithosphere, Magi et al. (2000) have suggested that the crust is in fact the strongest layer. Thus, it is

emerging that the strength of lithospheric upper mantle may vary considerably. Over the last decade we have studied a number of peridotite massifs, both of orogenic ('alpine-type') and ophiolitic origin, in order to establish under which conditions, and to what extent, deformation in mantle rocks is localized. In this contribution we present an overview of shear localization features in these peridotite massifs. From our data as well as published data from other workers it emerges that, under lithospheric conditions, localization of deformation in mantle rocks is a common phenomenon (Fig. 1; Table I). In addition, shear localization seems to occur over a wide range of pressure-temperature conditions, producing a number of distinct shear zone types. The range of shear zone types is comparable to types of deformation structures found in crustal rocks (e.g. Carreras et al. 1980). Microstructural work has highlighted the importance of mineral reactions as a softening mechanism in the development of mantle mylonite shear zones (Handy 1989; Drury et al.

From: ALSOP, G. L, HOLDSWORTH, R. E., MCCAFFREY, K. IW. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224,11-24. 0305-8719/$15.00 © The Geological Society of London 2004.

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A. H. DIJKSTRA ET AL.

Fig. 1. Shear zone features in selected mantle massifs (see references in figure). Tectonites in Voltri (Italy), Othris (Greece) and Ronda (Spain) are fine-grained tectonites. Tectonites in Oman are mostly coarse-grained, 'asthenospheric' in character, except for some relatively fine-grained tectonites in the 'sub-Moho shear zone' in the Hilti Massif (see text and Fig. 3).

1990; Newman et al. 1999; Furusho & Kanagawa 1999; Handy & Stunitz 2002). Because of the key role of these reactions, we predict that the mechanical strength of the mantle lithosphere may vary in different tectonic settings, and that the greatest reduction of the bulk strength may be expected in systems in which mantle rocks move quickly through pressure-temperature space, such as in lithosphere-scale extensional detachment and/or subduction zone systems. We also conclude that reaction-derived mantle mylonite zones are relatively stable features, that may be easily reactivated, and that mechanical models of the mantle lithosphere should take the possible presence of weak shear zones into account.

Structures and microstructures in mantle rocks Based on mapping and microstructural analysis of different peridotite massifs, we recognize three main types of deformation structures and microstructures in mantle rocks: medium- to coarse-grained tectonites, fine-grained tectonites and peridotite mylonites. We use the term 'tectonite' to describe strongly deformed rocks with foliation and lineation defined by the shape of pyroxene, spinel and olivine grains. The recrystallized grain size varies from coarse (>4 mm), medium (1-4 mm) to fine (1-0.1 mm). The term mylonite is used to describe a cohesive fault (or shear zone) rock with strong foliation

Table 1. Peridotite massifs with evidence for localized deformation Massif

Type

Type of shear zone

References f

Anita Bay (New Zealand) Balmuccia (Italy)

a* a

Beni Bousera (Morocco) Hidaka (Japan) Horoman (Japan) Lanzo (Italy) Lizard (England) Oman

a t t t t o

Othris (Greece) Ronda (Spain) Table Mountain (Canada) Turon de Tecouere (France)

o a o a

Peridotite mylonites (tp) Peridotite mylonites (e) Hydrous peridotite mylonites (c) Peridotite mylonites (e)

Voltri (Italy)

t

Vourinos (Greece)

o

Raudhaugene (Norway) Zabargad (Red Sea)

a t

Fine grained tectonites (e) Hydrous peridotite mylonites grading into serpentinite mylonites (e) Peridotite gauge zones Hydrous peridotite mylonites grading into serpentinite mylonites (c) Peridotite mylonites (?) Hydrous peridotite mylonites (t)

Hydrous peridotite mylonites (e ) Fine grained tectonites (e) Pseudotachylytes (e?) Mylonites (e) Peridotite mylonites (c) Peridotite mylonites (c) Peridotite mylonites (e) Hydrous peridotite mylonites (e) Coarse grained tectonites (c?) Peridotite mylonites (c)

Wood (1972); Hill (1995) Boudier et al (1984); Brodie & Rutter (1987); Skrotzki et al. (1990) Jin et al. (1998) Reuberetal. (1982) Furusho & Kanagawa (1999) Sawaguchi (2002) Boudier (1978) Cook etal (2000) Ceuleneer et al. (1988); Ceuleneer & Rabinowicz (1992); Dijkstra et al. (20026) Boudier et al. (1988) Rassios & Konstantopoulou (1993); Dijkstra et al. (20020) Van der Wai & Vissers (1993,1996) Suhr (1993) Vissers etal. (1997); Fabrics et al. (1998); Newman et al. (1999) Drury et al. (1990); Vissers et al. (1991); Hoogerduijn Strating et al. (1993); Vissers et al. (1995,1998) Rassios et al. (1994); Rassios & Smith (2002); Personal observation (A.H.D.) Van Roermund & Drury (1998); Unpublished results (H.v. R.) Nicolas et al. (1987); Piccardo et al. (1993)

*)a = alpine-type (subcontinental mantle); o = ophiolitic (oceanic mantle); t = transitional (generally rifted margin) )e = extensional (including diapiric); c = contractional; tp = transpressional; t = transcurrent

f

14


SHEAR ZONES IN MANTLE ROCKS

and fine scale compositional banding resulting from tectonic grain size reduction. Usually, peridotite mylonites have a very fine (0.1-0.01 mm) to ultra-fine (1 mm) olivine crystals showing extensive evidence for recovery and dynamic recrystallization by subgrain rotation and fast grain boundary migration. There is generally a strong lattice preferred orientation present, with the dominant slip system [a] (010) (where [a] is the slip direction and (010) the slip plane). Medium- to coarse-grained tectonites often contain evidence for the presence of small amounts of interstitial melt during deformation. It is generally assumed that these tectonites form by high temperature (>1200 °C, i.e. near- to supersolidus temperatures), low stress (of the order of a few MPa) deformation, at conditions similar to those in the asthenosphere. These medium- to coarse-grained tectonites are mainly restricted to the mantle sections of ophiolite massifs and are usually non-localized at the scale

15

of the massifs. However, there is evidence for localization of high-temperature, 'asthenospheric' deformation in a 500-800 m wide meltrich, high-strain tectonite horizon immediately below the crust-mantle boundary in the Oman Ophiolite (Fig. 3; Dijkstra et al. 20020). The tectonites in this 'sub-Moho-shear zone' are medium- to fine-grained, with an average recrystallized grain size of about 0.5 mm, but otherwise strongly resemble highly recrystallized, 'asthenospheric' tectonites (Fig. 3).

Fine-grained tectonites Fine-grained tectonites (Fig. 2a), equivalent to the fine-porphyroclastic 'lithospheric' microstructures of Nicolas (1986), Ceuleneer et al. (1988) and Ildefonse et al. (19986), have olivine crystals generally smaller than 1 mm, often with a strongly developed dislocation substructure (undulatory extinction and well defined, closely spaced subgrain walls). Lattice preferred orientations are often consistent with [a] {Okl} slip. These are produced by relatively high stress (generally >10MPa) deformation, at temperatures at which recovery is not very efficient (950-1200 °C). Recrystallization is dominantly by subgrain rotation with moderate grain boundary migration. Fine-grained tectonites have been found in a number of different massifs, including the Voltri (Italy), Othris (Greece), and Ronda (Spain) massifs (Fig. 1). In Voltri, there is good evidence that fine-grained tectonites are essentially a form of localized deformation; in the Tugello and Tobbio areas, fine-grained tectonites occur in a kilometre-scale deformation zone within or adjacent to a kilometre-scale domain of granular peridotites (Fig. 1). Often, however, the limited dimensions of mantle massifs do not allow determination as to whether the deformation in the fine-grained tectonites was homogeneous, or localized in kilometre-scale shear zones.

Fig. 2. Photomicrographs of typical mantle deformation microstructures. Photo (e) in plane polarized light, all others in cross polarized light, (a) Fine-grained tectonite from Othris, Greece. Crystals in field of view are all olivine; (b) Mylonite from Turon de Tecouere, France, showing olivine porphyroclast in a fine-grained matrix consisting of olivine, pyroxene, spinel, plagioclase and minor amphibole; (c) Mylonite from Othris, showing olivine lenses and stretched orthopyroxene porphyroclast in fine-grained matrix of olivine and orthopyroxene; (d) Mylonitic band which formed in a peridotite which shows extensive development of fine-grained, reactionderived olivine and orthopyroxene. Also note the clast-like behaviour of plagioclase in the mylonitic band, suggesting that plagioclase was stronger than the fine-grained matrix during deformation. Sample from Othris; (e) Spinel mylonite from Voltri, Italy, with spinel porphyroclast in a fine-grained polyphase matrix consisting of bands of mostly olivine (light bands) and olivine, pyroxene, spinel and amphibole (dark bands); (f) Coarse amphibole (hornblende) fibres in the pressure shadows of a stretched orthopyroxene porphyroclast. The finegrained material mantling the pyroxene is also mainly amphibole. Sample from Voltri; (g) Fine-grained tectonite from Othris, showing abundant fine-grained olivine and orthopyroxene produced by the breakdown of orthopyroxene porphyroclasts, probably by a reaction involving a melt. The same reaction produced the fine-grained matrix of the Othris mylonites; (h) Peridotite gouge and possible mylonite precursor in dunite sample from Vourinos, Greece, consisting of olivine and minor (opaque) chromite.

16

A. H. DIJKSTRA ETAL.

Peridotite mylonites Peridotite mylonites (Figs 1, 2b-e) generally consist of a very fine grained matrix with olivine grain sizes as small as a few microns. The matrix often contains porphyroclasts of olivine (Fig. 2b) as well as other phases (e.g. spinel, pyroxene, plagioclase, garnet, amphibole; Fig. 2c-f), or lenticular domains of relatively coarse-grained olivine (Fig. 2c). Pyroxene porphyroclasts are often strongly stretched within the foliation plane (Fig. 2c); it is not uncommon to find pyroxene porphyroclasts with aspect ratios more than 10. Peridotite mylonite shear zones occur over a wide range of scales, from millimetrewide bands to kilometre-scale shear zones (e.g. in Othris; Fig. 1). Peridotite mylonites can be anhydrous (Othris; Ronda; Raudhaugene, Norway; and Oman), or weakly hydrous with the growth of small amounts of amphibole (Turon de Tecouere, N. Pyrenees, France). Strongly hydrous peridotites with the development of abundant amphibole (Fig. 2f) are found in Voltri (Hoogerduijn Strating et al. 1993) and in the mantle section of the Lizard ophiolite, England (Cook et al. 2000). The hydrous mylonites of the Voltri massif also contain chlorite, as well as various amounts of antigorite serpentine, and some of these mylonites grade into serpentinite mylonites. The formation of hydrous minerals requires fluid infiltration into the shear zones and the fluids may have been derived from adjacent crustal units during exhumation and after crustal emplacement of the peridotites. Relative softening mechanisms In the absence of large temperature, stress or fluid pressure gradients, shear localization in rocks deforming by crystal plastic or diffusion creep mechanisms requires a softening mechanism that weakens the material inside the shear zone with respect to the material in the wall rock (Bowden 1970; Poirier 1980; White et al. 1980; Hobbs etal. 1990; Drury etal 1991; Rutter 1999). Without a relative softening mechanism, the zone of localized deformation would widen rapidly resulting in relatively homogeneous deformation on the scale of the peridotite body. In many natural shear zones a history of progressive localization can be reconstructed (e.g. Vissers et al. 1991; Drury et al 1991; Jin et al. 1998). Softening mechanisms in general have been discussed in White & Knipe (1978), Poirier (1980) and White et al. (1980). Softening mechanisms in mantle rocks have been discussed in Drury et al. (1991) and Jin et al. (1998).

Below, we summarize the relevant mechanisms responsible for localization of deformation in tectonites and peridotite mylonites.

Relative softening in medium- to coarsegrained peridotite tectonites High temperature mantle flow producing medium- to coarse-grained tectonites appears to be homogeneous on the scale of individual peridotite bodies. Localization of high temperature, 'asthenospheric' mantle flow has so far only been observed in the mantle section of the Oman Ophiolite, in a melt-rich horizon immediately below the crust-mantle boundary (Ceuleneer etal. 1988; Ceuleneer & Rabinowicz 1992). Based on our microstructural study of such a melt-rich shear zone in the Hilti Massif (Fig. 3), we concluded that melt weakening was responsible for softening of the peridotites in the shallowmost 500-800 m of the mantle section (Dijkstra et al. 20020). Weakening of olivine aggregates deforming by dislocation creep containing an interstitial melt fraction is observed in laboratory deformation experiments (Hirth & Kohlstedt 1995; Bai et al. 1997). This weakening effect, which is proportional to the melt content but which only becomes significant when more than 4% melt is present, is probably related to an enhancement of grain boundary sliding and/or diffusion along melt-wetted grain boundaries (Hirth & Kohlstedt 1995). As this type of melt weakening requires high melt contents, higher than normally present in the upper mantle, shear localization of a result of melt weakening is only feasible in regions of melt accumulation (Dijkstra et al. 20020). Holtzmann et al. (2003) have recently shown that the compaction length of the melt flow system controls the spacing of localized shear zones produced by melt weakening.

Relative softening in fine-grained tectonites Drury et al. (1991) argued that shear heating could probably be ruled out as a softening mechanism in fine-grained tectonites. Recent experimental studies suggest a possible softening mechanism related to a change in the ratecontrolling slip system in olivine (Fig. 4). In olivine aggregates deforming by dislocation creep a transition may occur from creep controlled by the hard [c](010) system, to creep controlled by the weak [a]-slip system accommodated by grain boundary sliding and diffusion (see Hirth & Kohlstedt 1995 for details). Dislocation creep in fine-grained


17

Fig. 3. Cross-section across the Hilti mantle section, Oman, showing the presence of a melt-rich, tectonite shear zone immediately below the crust-mantle boundary. Also shown are tracings of representative microstructures in dunites from shear zone and wall rock (modified from Dijkstra et al. 20020).

Fig. 4. Olivine deformation mechanism map for the Othris fine grained tectonites, showing the fields for [c]- and [a]-slip limited creep. Deformation conditions for the Othris tectonites are shown by grey field. Weakening in fine-grained tectonites may be the result of switching off the hard [c]-slip system. For details on flow laws used see Dijkstra et al. (20026).

Fig. 5. Scanning electron microscope (SEM) backscatter image of grain boundary alignments in Turon de Tecouere mylonites.

controlled creep to [a]-slip controlled creep (Jin et al. 1998).

Relative softening in peridotite mylonites olivine limited by the weak [a]-slip system can be up to two orders of magnitude faster than creep limited by hard [c]-slip (Mackwell et al. 1990; Drury & Fitz Gerald 1998). Concentration of deformation in a tectonite shear zone, causing faster grain size reduction by dynamic recrystallization within the shear zone compared to the wall rock, may thus lead to strain softening of the shear zone tectonites by a change from [c]-slip

Peridotite mylonite zones are characterized by the small grain size of the matrix. There is often evidence that grain size sensitive (GSS) creep probably a combination of diffusion creep plus grain boundary sliding - has been the dominant deformation mechanism particularly in the polyphase layers (Drury et al. 1991; Newman et al. 1999; Jaroslow et al. 1996; Dijkstra et al. 20025). For instance, in the Turon de Tecouere

18


Fig. 6. SEM Backscatter images showing evidence for production of fine-grained material by reactions, (a) Clinopyroxene porphyroclast reacting to fine-grained plagioclase+spinel-lherzolite assemblage in Turon de Tecouere mylonites; (b) Spinel porphyroclast from coarse spinel-lherzolite assemblage reacting to finegrained plagioclase+spinel-lherzolite assemblage in Turon de Tecouere mylonites; (c) Garnet porphyroclast in garnet-lherzolite assemblage reacting to new garnet+spinel-lherzolite assemblage in Raudhaugene mylonites.

mylonites (Fig. 1), there is only a weak lattice preferred orientation in the matrix olivine and the matrix grains are virtually dislocation free (Newman et al. 1999). In addition, grain boundaries in the matrix are often aligned subparallel to the foliation (Fig. 5), over several grain lengths, and these alignments are interpreted as sliding surfaces (Newman et al. 1999). The Othris mylonites (Fig. 1) are characterized by the absence of a lattice preferred orientation in the matrix grains and by numerous grain boundary alignments (Dijkstra et al. 20026). This evidence suggests that softening and localization in mylonites are the result of a change of deformation mechanism, from dislocation to grain size sensitive creep, brought about by extreme grain size reduction. Detailed microstructural and petrological work in peridotite mylonites in Turon de Tecouere, Othris, and Raudhaugene has further shown that reactions play a key role in produc-

ing the very fine grain size of the matrix. Newman et al. (1999) demonstrated that in Turon de Tecouere small, micron-sized grains were produced by a continuous net transfer reaction associated with the transformation of spinel-lherzolite assemblage into a plagioclaseIherzolite assemblage (Fig. 6a,b). This reaction took place over a broad pressure range of 5-11 kbar (Fig. 7). The same reaction causing grain size reduction was also responsible for the development of the Hidaka mylonites (Furusho & Kanagawa 1999). Similarly, grain size reduction associated with the formation of mylonites in the Raudhaugene peridotite body in western Norway (Van Roermund & Drury 1998; Van Roermund et al. 2001) involved a continuous reaction in which a garnet-lherzolite assemblage was transformed into a spinel-lherzolite assemblage (Fig. 6c). In the Othris peridotites, a reaction between harzburgites and a percolating melt led to the formation of a mixture of


Fig. 7. Pressure-temperature grid with exhumation paths for Pyrenean ('Py'; e.g. Turon de Tecouere) and Norwegian ('N'; e.g. Raudhaugene) peridotites, showing the possible reaction lines that may be responsible for reaction-softening. Note that, as phases involved are solid solutions, reactions occur continuously over a range of pressures and temperatures (shaded fields).

fine-grained material which later formed the matrix for the peridotite mylonites (Fig. 2d,g; Dijkstra et al. 20026). A similar sort of reaction is observed in Lizard (Cook et al. 2000). Finally, in Voltri, amphibole-, chlorite-, and antigoriteproducing reactions played a role in the development of the hydrous mylonites (Fig. 2e,f; Hoogerduijn Strating et al. 1993). Clearly, these examples show how that reaction-enhanced softening, by both subsolidus and near-solidus reactions, as well as by hydrous and anhydrous reactions, is an important mechanism leading to grain size reduction and localization in mantle rocks during metamorphism. Metamorphic reactions may also induce softening owing to transformational induced plasticity (White & Knipe 1978). Reactions do not always produce grain size reduction. In fact, grain size reduction should occur during retrograde reactions, whereas, prograde reactions may result in coarsening and strain hardening. Metamorphic reactions are often considered to occur over a limited range of P-T conditions. Many reactions in peridotites, however, are continuous, involving solid solution phases, so the reactions occur over a broad region of

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pressure-temperature space spanning a large part of the lithospheric mantle (Fig. 7). Moreover, because of the polyphase nature of finegrained reaction products, subsequent grain growth after the reaction is inhibited (Olgaard 1990). This means that once formed, reactionderived mylonites are relatively stable features, which can be reactivated during subsequent deformation events. Another consequence of the key role of reactions in the development of peridotite mylonites is that fine grained rocks, such asfine-grainedtectonites - if polyphase might be better candidates for future mylonites than coarse-grained, undeformed rocks because reaction kinetics will tend to be faster in finergrained rocks. The growth of hydrous minerals in mylonites requires fluid-infiltration and probably a component of brittle-dilatant deformation. It is therefore possible that hydrous peridotite mylonites form by localization of the brittle component of deformation (Hobbs et al. 1990; Drury et al. 1991; Jin et al 1998; Handy and Stiinitz 2002). Cataclasis may also contribute to grain size reduction (Jaroslow et al. 1996; Handy & Stiinitz 2002). In the mantle section of the Vourinos Ophiolite, Greece, there is a close spatial and temporal association of brittle gouge zones, hydrous peridotite mylonites and serpentinite mylonites, which are often the locus of high-temperature mineralizing fluid flow (Rassios et al. 1994). Some of the peridotite mylonites in Vourinos may have started off as fine grained brittle gouge zones such as shown in Fig. 2h. Gouge zones and serpentinite mylonites may be formed after the peridotite bodies are emplaced in the upper crust, but low-temperature fault gouges and mylonites like this should also occur in the shallow mantle beneath oceanic crust and thinned continental margins. Discussion There is now a wealth of field and detailed microstructural observations suggesting that localization of deformation in tectonite and mylonite shear zones may be the rule rather than the exception in mantle rocks. Evidence for localized deformation is found in alpine, ophiolite and transitional type mantle massifs. Similar evidence for tectonite and mylonite mantle structures is also found in peridotite xenoliths from alkali basalts (e.g. Mercier & Nicolas 1975; Hunter et al. 1984; Cabanes & Brique 1986; Downes 1987,1990; Cabanes & Mercier 1988; Xu et al. 1993). There remains, however, the question as to how representative such structures are for true mantle deformation. The structures

20


Fig. 8. Possible tectonite and mylonite shear zones in extensional and transcurrent systems under hydrous conditions. In transcurrent settings, the formation of reaction-derived mylonites is much less likely than in extensional (or contractional) systems, as mantle rocks do not move rapidly through pressure-temperature space. Lithospheric weakening can be caused by the formation of tectonite shear zones in such settings. The formation of hydrous mylonites as the result of ingress of fluids can lead to the formation of mylonites in the roots of transcurrent systems.

present in exposed upper mantle rocks may have been formed when the peridotites were part of the crust, or in-situ within the upper mantle. Insitu temperature (T) and pressure (P) conditions vary widely within the upper mantle, ranging from very low T and P conditions beneath oceanic crust through to ultra-high P and Tin the convecting upper mantle. In the ophiolitic massifs discussed, relations between deformation and magmatism (Othris, Oman) and/or high-temperature fluid infiltration (Vourinos) demonstrate that deformation structures are related to on-ridge or early, intra-oceanic, deformation events taking place well before the rocks became emplaced onto continental margins. There is little doubt that such structures were produced within the shallow oceanic mantle. In continental settings, there is considerable overlap between P—T conditions within the shallow upper mantle and crust, depending on crustal type and thickness, making it difficult to identify mantle or crustal structures in alpinetype peridotite bodies solely on the basis of P-T conditions of deformation. Even so, we argue that there is abundant evidence for localization of deformation in rocks with mantle compositions under conditions that are at least equivalent to those in the mantle lithosphere. Experimental studies show that quartzo-feldspathic crustal rocks are significantly weaker than dry olivine-pyroxene mantle rocks (Kohlstedt et al. 1995). As a result, dry mantle rocks will behave as strong bodies within the crust after incorporation into that crust. The common discordance of structures in peridotite bodies compared to crustal country rock structure is consistent with limited deformation of the peridotites after crustal emplacement. A weak serpentinite rim and internal serpentinite shear zones usually accommodate any crustal defor-

mation in peridotite bodies. Some examples of intense coherent deformation of crust and mantle rocks occur in hydrated peridotite bodies included in high-pressure crustal terrains (Wood 1972; Cordellier et al. 1981). The extensive deformation of these mantle rock bodies within the crust can be explained by water weakening of olivine (Chopra & Paterson 1984). Shear zone structures in ophiolitic massifs are generally contractional, related to the ophiolite emplacement, whereas shear zones in alpinetype peridotite massifs are often related to extensional exhumation of the massifs. This is essentially a sampling bias because we tend to observe the structures that have brought the mantle massifs to the earth surface. In order to extrapolate our observations to the unexposed part of the mantle it is crucial to understand the processes that are responsible for shear zone development. As discussed above, we have found that during exhumation of the mantle footwall in an extensional detachment system (Fig. 8), drastic weakening occurs through the development of peridotite mylonites by deformation and reactions in mantle mineral assemblages under changing pressure and temperature conditions. The same processes may produce similar mylonites during burial in a thickening lithospheric mountain root, or in a subducting slab. In contrast, we expect that peridotite tectonites rather than mylonites are the dominant mode of localized deformation in the root zones of lithospheric scale transcurrent shear zones. In such systems driving forces for reactions are absent, as mantle rocks are essentially stationary in pressure-temperature space. Only as soon as fluid infiltration and/or brittle processes set in may very fine-grained mylonites be produced, leading to significant weakening of the lithospheric mantle (Fig. 8).


Implications for the strength of the mantle The upper part of the mantle is potentially the strongest part of the subcontinental lithosphere (e.g. Kirby 1983; Kuznir & Park 1986), however, the development of mantle shear zones may reduce the strength of the mantle lithosphere significantly (Rutter & Brodie 1988). Under some conditions, mantle lithosphere containing mylonite shear zones can be shown to be weaker than dry lower crustal, feldspar-dominated rocks (Fig. 2d; Dijkstra et al 20026). Mantle shear zones will, when present, play an important role during mountain building or during continental break-up (Vissers et al 1995; Vissers et al 1997; Handy & Stunitz 2002). The notion of a relatively weak continental mantle lithosphere is supported by recent assessments of lithosphere strength based on the depth distribution of seismicity (Magi et al 2000). Mylonite shear zones may be relatively stable structures that are not easily annealed (probably in contrast to tectonite shear zones) because of the polyphase fine-grained matrix in reactionderived peridotite mylonites (which suppresses grain growth). Mylonitic shear zones may therefore persist, and be easily reactivated during subsequent deformation events (e.g. Bailey et al 2000). Models of the strength of the mantle lithosphere need to take the presence of such shear zones, which act as pre-existing weaknesses, into account. In Turon de Tecouere it was found that strain rates in ultramylonitic and protomylonitic peridotites were equal, but that stresses were two times lower within the ultramylonites (Newman & Drury 2000). In Othris, estimated strain rates in peridotite mylonites were up to four orders of magnitude faster than strain rates in the wall rock (Dijkstra et al 20025). Conversely, lithospheric mantle containing only one volume percent of Othris-like, finegrained mylonitic material, for instance in an anastomosing network of shear zones, has a bulk strength that is a hundred times lower than mantle lithosphere without any mylonites. Conclusions Shear localization is a common process in mantle peridotite massifs. Reactions play a key role in the development of many peridotite mylonite zones, and these reactions occur over a large pressure and temperature range, both under hydrous/anhydrous and subsolidus/nearsolidus conditions. Shear zones are probably omnipresent in the lithospheric mantle and need to be taken into account in modelling studies. They may reduce the bulk strength of the mantle by up to two orders of magnitude.

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This contribution is partly based on the work of many students and postgraduate students at Utrecht University. In particular, Eilard Hoogerduijn Strating and Dirk Van der Wai have made crucial contributions. Geoff Lloyd and Mike Bickle are thanked for numerous constructive comments and suggestions. This work has been supported by NWO Pioneer subsidy No. 030-75-346.

References BAI, Q., JIN, Z-M. & GREEN II, H. 1997. Experimental investigation of the rheology of partially molten peridotite at upper mantle pressures and temperatures. In: HOLNESS, M.B. (ed.) Deformationenhanced fluid transport in the Earth's crust and mantle. Chapman & Hall, London, 40-61. BAILEY, W.R., HOLDSWORTH, R.E. & SWARBRICK, R.E. 2000. Kinematic history of a reactivated oceanic suture: the Mamonia Complex Suture Zone. Journal of the Geological Society, London, 157, 1107-1126. BOUDIER, F. 1978. Structure and petrology of the Lanzo peridotite massif (Piedmont Alps). Geological Society of America Bulletin, 89,1574-1591. BOUDIER, E, JACKSON, M. & NICOLAS, A. 1984. Structural study of the Balmuccia Massif (Western Alps): a transition from mantle to lower crust. Geologic en Mijnbouw, 63,179-188. BOUDIER, E, CEULENEER, G. & NICOLAS, A. 1988. Shear zones, thrusts and related magmatism in the Oman ophiolite: initiation of thrusting on an oceanic ridge. Tectonophysics, 151, 275-296. BOWDEN, PB. 1970. A criterion for inhomogeneous plastic deformation. Philosophical Magazine, 22, 455-462. BRODIE, K. & RUTTER, E. 1987. Deep crustal extensional faulting in the Ivrea Zone of northern Italy. Tectonophysics, 140,193-212. BRUN, J.-P 2002. Deformation of the continental lithosphere: insights from brittle-ductile models. In: DE MEER, S., DRURY, M.R., DE BRESSER, J.H.P & PENNOCK, G.M. (eds) Deformation mechanisms, rheology and tectonics: current status and future perspectives. Geological Society, London, Special Publications, 200, 355-370. CABANES, N. & BRIQUE, L. 1986. Hydration of an active shear zone: Interactions between deformation, metasomatism and magmatism: the spinel Iherzolites from the Montferrier (Southern France) Oligocene basalts. Earth and Planetary Science Letters, 81, 233-244. CABANES, N. & MERCIER, J.-C.C. 1988. Insights into the upper mantle beneath an active extensional zone: the spinel-peridotite xenoliths from San Quintin (Baja California, Mexico). Contributions to Mineralogy and Petrology, 100, 374-382. CARRERAS, J., COBBOLD, PR., RAMSAY, J.G. & WHITE, S.H. 1980. Shear zones in rocks. Journal of Structural Geology, 2,1-287. CEULENEER, G. & RABINOWICZ, M. 1992. Mantle flow and melt migration beneath oceanic ridges: models derived from observations in ophiolites. In: MORGAN, J.P, BLACKMAN, D.K. & SINTON, J.M.

22


(eds) Mantle flow and melt generation at midoceanic ridges, Geophysical Monograph 71, American Geophysical Union, 123-154. CEULENEER, G, NICOLAS, A. & BOUDIER, F. 1988. Mantle flow patterns at an oceanic spreading center: the Oman peridotites record. Tectonophysics, 151,1-26. CHEN, W.-P. & MOLNAR, P. 1983. Focal depths of intracontinental and intraplate earthquakes and their implications for the thermal and mechanical properties of the lithosphere. Journal of Geophysical Research, 88,4183-4214. CHOPRA, P.N. & PATERSON, M.S. 1984. The role of water in the deformation of dunite. Journal of Geophysical Research, 89, 7861-7876. CORDELLIER, F., BOUDIER, F. & NlCOLAS, A.

1981.

Structural study of the Almklovdalen peridotite massif (Southern Norway). Tectonophysics, 77, 257-281. COOK, C.A., HOLDSWORTH, R.E., STYLES, M.T. & PEARCE, J.A. 2000. Pre-emplacement structural history recorded by mantle peridotites: an example from the Lizard Complex, SW England, Journal of the Geological Society, London, 157, 1049-1064. DIJKSTRA, A.H., DRURY, M.R. & FRIJHOFF, R.M. 20020. Microstructures and lattice fabrics in the Hilti mantle section (Oman Ophiolite): evidence for shear localization and melt weakening in the crust-mantle transition zone? Journal of Geophysical Research, 107, Bll, 2270, doi: 10.1029/2001JB000458. DIJKSTRA, A.H., DRURY, M.R., VISSERS, R.L.M. & NEWMAN, J. 2002£. On the role of melt-rock reaction in mantle shear zone formation in the Othris Peridotite Massif (Greece). Journal of Structural Geology, 24,1431-1450. DOWNES, H. 1987. Relationship between geochemistry and textural type in spinel Iherzolites, Massif Central and Languedoc, France. In: NIXON, PH. (ed.) Mantle xenoliths. John Wiley & Sons Ltd, New York, 125-133. DOWNES, H. 1990. Shear zones in the upper mantle: relation between geochemical enrichment and deformation in mantle peridotites. Geology, 18, 374-377. DRURY, M.R. & FITZ GERALD, J.D. 1998. Mantle rheology: insights from laboratory studies of deformation and phase transition. In: JACKSON, I.N.S. (ed.) The Earth's mantle - composition, structure and evolution. Cambridge University Press, 503-559. DRURY, M.R., HOOGERDUIJN STRATING, E.H. & VISSERS, R.L.M. 1990. Shear zone structures and microstructures in mantle peridotites from the Voltri Massif, Ligurian Alps, NW Italy. Geologie en Mijnbouw, 69, 3-17. DRURY, M.R., VISSERS, R.L.M., VAN DER WAL, D. & HOOGERDUIJN STRATING, E.H. 1991. Shear localization in upper mantle peridotites. Pure and Applied Geophysics, 137, 439-460. FABRIES, J., LORAND, J.-P & BODINIER, J.-L. 1998. Petrogenetic evolution of erogenic Iherzolite massifs in the central and western Pyrenees. Tectonophysics, 292,145-167.

FURUSHO, M. & KANAGAWA, K. 1999. Transformationinduced strain localization in a Iherzolite mylonite from the Hidaka metamorphic belt of central Hokkaido, Japan. Tectonophysics, 313, 411-432. HANDY, M.R. 1989. Deformation regimes and the rheological evolution of fault zones in the lithosphere: the effects of pressure, temperature, grain-size and time. Tectonophysics, 163,119-152. HANDY, M.R. & STUNITZ, H. 2002. Strain localization and reaction weakening: a mechanism for initiating exhumation of subcontinental mantle beneath rifted margins. In: DE MEER, S., DRURY, M.R., DE BRESSER, J.H.P. & PENNOCK, G.M. (eds) Deformation mechanisms, rheology and tectonics: current status and future perspectives. Geological Society, London, Special Publications, 200, 387-407. HILL, EJ. 1995. A deep crustal shear zone exposed in western Fiordland, New Zealand. Tectonics, 14, 1172-1181. HIRTH, G. & KOHLSTEDT,D.L. 1995. Experimental constraints on the dynamics of the partially molten upper mantle 2. Deformation in the dislocation creep regime. Journal of Geophysical Research, 100,15441-15449. HOBBS, B.E., MUHLHAUS, H.-B. & ORD, A. 1990. Instability, softening and localization of deformation. In: KNIPE, R.J. & RUTTER, E.H. (eds) Deformation mechanisms, rheology and tectonics. Geological Society, London, Special Publications, 54, 143-165. HOLTZMAN, B.K., GROEBNER, N.J., ZIMMERMAN, M.E., GINSBERG, S.B. & KOHLSTEDT, D.L. 2003. Deformation-driven melt segregation in partially molten rocks. Geochemistry, Geophysics and Geosystems, 4, doi: 10.1029/2001GC000258. HOOGERDUIJN STRATING, E.H., RAMPONE, E., PicCARDO, G.B., DRURY, M.R., VISSERS, R.L.M. 1993. Subsolidus emplacement of mantle peridotites during incipient oceanic rifting and opening of the Mesozoic Tethys (Voltri Massif, NW Italy). Journal of Petrology, 23, 901-927. HUNTER, R.H., UPTON, B.C. & ASPEN, P. 1984. Metaigneous granulite and ultramafic xenoliths from basalts of the Midland Valley of Scotland: petrology and mineralogy of the lower crust and upper mantle. Transactions of the Royal Society of Edinburgh: Earth Sciences, 75, 75-84. ILDEFONSE, B., BOUDIER, F. & NICOLAS, A. 19980. Asthenospheric deformation in ophiolitic peridotite. In: SNOKE, A.W., TULLIS, J. & TODD, V.R. (eds) Fault-related rocks - a photographic atlas. Princeton University Press, Princeton, New Jersey, 588-589. ILDEFONSE, B., BOUDIER, F. & NICOLAS, A. 19986. Lithospheric deformation in ophiolitic peridotite. In: SNOKE, A.W., TULLIS, J. & TODD, V.R. (eds) Fault-related rocks - a photographic atlas. Princeton University Press, Princeton, New Jersey, 590-591. JAROSLOW, G.E., HIRTH, G. & DICK, H.J.B. 1996. Abyssal peridotite mylonites: implications for grain-size sensitive flow and strain localization in the oceanic lithosphere. Tectonophysics, 256, 17-37.

SHEAR ZONES IN MANTLE ROCKS JIN, D., KARATO, S.-I. & OBATA, M. 1998. Mechanisms of shear localization in the continental lithosphere: inference from the deformation microstructures of peridotites from the Ivrea zone, northwestern Italy. Journal of Structural Geology, 20,195-209. KiRBY,S.H. 1983. Rheology of the lithosphere. Reviews of Geophysics and Space Physics, 21,1458-1487. KOHLSTEDT, D.L., EVANS, B. & MACKWELL, S.M.

1995.

Strength of the lithosphere: constraints imposed by laboratory experiments. Journal of Geophysical Research, 100,17587-17603. KUZNIR, N.J. & PARK, R.G. 1986. Continental lithosphere strength: the critical role of lower crustal deformation. In: DAWSON, IB., CARSWELL, D.A., HALL, J. & WEDEPOHL, K.H. (eds) The nature of the lower continental crust. Geological Society, London, Special Publications, 24, pp. 79-93. MACKWELL, S.J., BAI, Q. & KOHLSTEDT, D.L. 1990. Rheology of olivine and the strength of the lithosphere. Geophysics Research Letters, 17, 9-12. MAGI, A., JACKSON, J.A., MCKENZIE, D. & PRIESTLY, K. 2000. Earthquake focal depths, effective elastic thickness, and strength of the continental lithosphere. Geology, 28, 495-498. MERCIER, J.-CC. & NICOLAS, A. 1975. Textures and fabrics of upper-mantle peridotites as illustrated by xenoliths from basalts. Journal of Petrology, 16, 454-487. NEWMAN,! & DRURY,M.R. 2000. Deformation history along a peridotite shear zone with implications for dynamic processes within the upper mantle. Geological Society of America Abstracts with Programs, 32. NEWMAN, J., LAMB, W.M., DRURY, M.R. & VISSERS, R.L.M. 1999. Deformation processes in a peridotite shear zone: reaction-softening by an H2Odeficient, continuous net transfer reaction. Tectonophysics, 303,193-222. NICOLAS, A. 1986. Structure and petrology of peridotites: clues to their geodynamic environment. Reviews of Geophysics, 24, 875-895. NICOLAS, A., BOUDIER,F. & MONTIGNY, R. 1987. Structure of Zabargad Island and early rifting of the Red Sea. Journal of Geophysical Research, 92, 461-474. OLGAARD, D.L. 1990. The role of second phase in localizing deformation. In: KNIPE, R.J. & RUTTER, E.H. (eds) Deformation mechanisms, rheology and tectonics. Geological Society, London, Special Publications, 54,175-181. PICCARDO, G.B., RAMPONE, E., VANUCCI, R., SHIMIZU, N., OTTOLINI, L. & BOTTAZZI, P. 1993. Mantle processes in the sub-continental lithosphere: the case study of the rifted spirel-lherzolites from Zabargad (Red Sea). European Journal of Mineralogy, 5,1039-1056. POIRIER, J.P. 1980. Shear localization and shear instability in materials in the ductile field. Journal of Structural Geology, 2,135-142. RASSIOS, A. & KONSTANTOPOULOU, G. 1993. Emplacement tectonism and the position of chrome ores in the Mega Isoma peridotites, SW Othris, Greece. Bulletin of the Geological Society Greece, XXVIII/2, 463-474.

23

RASSIOS, A. & SMITH, A.G. 2000. Constraints on the formation and emplacement age of western Greek ophiolites (Vourinos, Pindos, and Othris) inferred from deformation structures in peridotites. In: DILEK, Y, MOORES, E.M., ELTHON, D. & NICOLAS, A. (eds) Ophiolites and oceanic crust: new insights from field studies and the ocean drilling program. Geological Society of America Special Paper 349, Boulder, Colorado, 473-483. RASSIOS, A., GRIVAS, E., KONSTANTOPOULOU, G. & VACONDIOS, I. 1994. The geometry of structures forming around the brittle-ductile transition in the Vourinos-Pindos-Othris oceanic slab. Bulletin of the Geological Society Greece, 30,109-121. REUBER, I., MICHARD, A., CHALOUAN, A. & JUTEAU,T. 1982. Structure and emplacement of the alpinetype peridotites from Beni Bousera, Rif, Morroco: a polyphase tectonic interpretation. Tectonophysics, 82, 231-251. VAN ROERMUND, H.L.M. & DRURY, M.R. 1998. Ultrahigh pressure (P>6GPa) garnet peridotites in Western Norway: exhumation of mantle rocks from more than 185 km. Terra Nova, 10,295-301. VAN ROERMUND, H.L.M., DRURY, M.R., BARNHOORN, A. & DE RONDE, A. 2001. Relict majoritic garnet microstructures from ultra-deep erogenic peridotites in Western Norway. Journal of Petrology, 42,117-130. RUTTER, E.H. 1999. On the relationship between the formation of shear zones and the form of the flow law for rocks undergoing dynamic recrystallization. Tectonophysics, 303,147-158. RUTTER, E.H. & BRODIE, K. 1988. The role of tectonic grain size reduction in the rheological stratification of the lithosphere. Geologische Rundschau, 77, 295-308. SAWAGUCHI, T. 2002. Structual petrology of the Horoman peridotite: deformation history and exhumation process of the Alpine-type peridotite massif. PhD Thesis, Waseda University, Japan. SKROTZKI, W, WEDEL, A., WEBER, K. & MULLER, WE 1990. Microstructure and texture in Iherzolites of the Balmuccia massif and their significance regarding the thermomechanical history. Tectonophysics, 179, 227-251. SUHR, G. 1993. Evaluation of upper mantle microstructures in the Table Mountain massif (Bay of Islands ophiolite). Journal of Structural Geology, 151,1273-1292. VAN DER WAL, D. & VISSERS, R.L.M. 1993. Uplift and emplacement of upper mantle rocks in the western Mediterranean. Geology, 21,1119-1122. VAN DER WAL, D. & VISSERS, R.L.M. 1996. Structural petrology of the Ronda Peridotite, SW Spain: deformation history. Journal of Petrology, 37, 23-43. VISSERS, R.L.M., DRURY, M.R., HOOGERDUIJN STRATiNG,E.H.& VANDERWAL,D. 1991. Shear zones in the upper mantle: a case study in an Alpine Iherzolite massif. Geology, 19, 990-993. VISSERS, R.L.M., DRURY, M.R., HOOGERDUIJN STRATING, E.H., SPIERS, C.J. & VAN DER WAL, D. 1995. Mantle shear zones and their effect on lithosphere strength during continental break-up. Tectonophysics, 249,155-171.

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VISSERS, R.L.M., DRURY, M.R., NEWMAN, J. & FLIERVOET, T.F. 1997. Mylonitic deformation in upper mantle peridotites of the North Pyrenean Zone (France): implications for strength and strain localization in the lithosphere. Tectonophysics, 279, 303-325. VISSERS, R.L.M., HOOGERDUIJN STRATING, E.H., DRURY, M.R. & VAN DER WAL, D. 1998. Microstructures formed during shear localization in the upper mantle. In: SNOKE, A.W.,TuLLis, J. & TODD, V.R. (eds) Fault-related rocks - a photographic atlas, Princeton University Press, Princeton, New Jersey, 582-587. WHITE, S.H. & KNIPE, RJ. 1978. Transformation- and reaction-enhanced ductility in rocks. Journal of the Geological Society London, 135, 513-516.

WHITE, S.H., BURROWS, S.E., CARRERAS, J., SHAW, N.D. & HUMPHREYS, FJ. 1980. On mylonites in ductile shear zones. Journal of Structural Geology, 2, 175-187. WOOD, B.L. 1972. Metamorphosed ultramafites and associated formations near Milford Sound, New Zealand. New Zealand Journal of Geology and Geophysics, 15, 88-127. Xu, Y.-G., Ross, J.V & MERCIER, J.-C.C. 1993. The upper mantle beneath the continental rift of Tanlu, Eastern China: Evidence for the intralithospheric shear zones. Tectonophysics, 225, 337-360.

Instability and localization of deformation in lower crust granulites, Minas fault zone, Nova Scotia, Canada JOSEPH C. WHITE Department of Geology, University of New Brunswick, Fredericton, NB, Canada, E3B 5A3 (email: [email protected]) Abstract: Blocks of granulite from within the megabreccia at Clarke Head, Nova Scotia, Canada contain extremely well preserved mylonitic and ultramylonitic textures developed in mineral assemblages for which thermobarometic calculations have indicated temperatures and pressures between 700-860 °C and 750-950 MPa. Deformation within these rocks is characterized by localization at several discrete length scales associated with the development of new microstructures comprising finer-grained material. Mylonitized granulite exhibits dislocation creep microstructures, with development of intense S-C fabrics and shear bands during the transition to ultramylonite. Dynamically recrystallization of plagioclase can be followed through progressive grain size reduction to about 5 um, but there remain extensive zones with grains less than 1 um in diameter. Localization of the these finest-grained ultramylonites occurs by transient frictional events associated with the introduction of partial igneous melts and formation of pseudotachylyte which produces abrupt decreases in grain size that cannot arise during dislocation mediated grain size reduction. The heterogeneous response of these rocks demonstrates the importance of considering characteristic length scales when assigning evidence from the rock record (e.g. palaeopiezometry) to bulk behaviour of the lithosphere. Associated with the localization of strain and subsequent strain softening is the observation that microstructures formed during the event that initiated the instability can be an obliterated by ductile flow. In instances where critical components of the microstructural evolution are known to have been largely overprinted, it becomes possible to reconcile contradictions in the rock record, such as production of ultra-fine-grained superplastic aggregates in what otherwise appears to be a dominantly dislocation creep regime.

The Minas fault system (MFS) in northwestern Nova Scotia, Canada (Fig. 1) demarcates a longlived, crustal-scale Appalachian tectonic boundary that juxtaposes Avalon and Meguma tectono-stratigraphic terranes (Keppie 1982; Webster et al 1998). The temporal and spatial complexities of this fault system that reflect transpressional deformation (Waldron et al 1989) provide opportunities to examine deformation processes at a range of crustal conditions. Of specific interest to this study are occurrences of strongly deformed granulites at Clarke Head (Fig. 1) that, as samples of the deep crust, are unique within the MFS. The granulites are central to the geological history along the terrane boundary as they reflect the crustal conditions most remote from the current upper crustal, extensional basin setting. Exposed granulites (Gibbons & Murphy 1995) take the form of isolated mylonitized blocks within a carbonate-evaporite-clay matrix (Fig. 2a) and are an integral component of the Clarke Head 'megabreccia', itself an anomalous feature of the MFS. A multidisciplinary study of the granulites by Gibbons et al (1996) has reported

Fig. 1. Location map of Clarke Head megabreccia containing mylonitized granulite blocks as part of the Minas fault system (MFS).

geothermobarometric, isotopic, TEM and tectono-stratigraphic data used to establish the P-T-t trajectory of the Clarke Head megabreccia over approximately 55 million years of convergence between Laurentia and Gondwana. The latter study has demonstrated the occurrence of syntectonic metamorphic mineral

From: ALSOP, G. I., HOLDSWORTH, R. E., MCCAFFREY, K. J. W. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224, 25-37. 0305-8719/$15.00 © The Geological Society of London 2004.

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Fig. 2. Field relationships of the granulites. (a) Megabreccia near Clarke Head, Nova Scotia. The large dark clasts are Carboniferous sedimentary rock. Granulite blocks are at the top of the cliff and eroded out on to beach, (b) Compositionally banded mylonitic granulite block with cross-cutting plagioclase-rich leucosome introduced as a melt, (c) Leucosome in mylonitic granulite in a stage of partial transposition, (d) The contact between granulite mylonite and ultramylonite comprising a change in fabric and bulk composition, although the mineral assemblage remains the same.

assemblages and microstructures indicative of lower crust conditions (700-860 °C and 750-950 MPa). Some retrogression of granulite blocks to amphibolite assemblages is observed, both by exhumation-related hydration and during introduction of late Cl-rich amphibole veins. Occurrence of the 'megabreccia' in its current configuration is distinct and unrelated to deformation of the granulites. This contribution pursues the nature of deformation microstructures in the granulites through detailed analytical electron microscopy with the aim of elucidating micromechanical records and their implication to the rheology of deep crustal rocks.

Clarke Head granulites: field observations The granulite blocks exhibit contrasting domainal layering (Fig. 2) defined by variations in mineralogy, grain size and deformation textures. Three domains are recognized within the blocks: host mylonite, ultramylonite and 'cherty' ultramylonite.

The host mylonite comprises the bulk of the granulites and although strongly deformed, it represents the earliest deformation and background strain. The mineralogical assemblage consists of clino- and orthopyroxene, plagioclase, K-feldspar and scattered garnet; ilmenite and magnetite are significant oxide phases. The principal foliation is a transposed, composite layering (Fig. 2b-d) defined by modal variations of feldspars and pyroxenes. Because the granulites occur as isolated blocks within the Clarke Head megabreccia, any reference of their internal structure (e.g. foliations) to an external geometric or kinematic reference frame is precluded. In this study, the compositional layering is taken as the primary reference frame (Cfoliation), in conjunction with a weak lineation within the compositional layering. A second foliation (S-foliation) oblique to, but within the compositional layers is defined by elongate plagioclase and pyroxene. Porphyroclasts of pyroxene and garnet within the compositional layering are typically 200-1000 um across, whereas feldspars dominate the finer-grained matrix. Plagioclase-rich

INSTABILITY AND DEFORMATION LOCALIZATION

layers (leucosomes) are observed both crosscutting this main compositional layering, and forming intrafolial isoclinal folds (Fig. 2b, c). The latter relationships record the introduction of anatectic melt (cross-cutting layers) and progressive high-strain transposition associated with formation of the compositional layering. The feldspathic layers commonly localize deformation as evidenced by their large attenuation and grain size reduction relative to pyroxenerich layers. Ultramylonite is recognized as zones of more intense deformation up to 40 cm thick that disrupt the predominant compositional layering of the mylonite at low angles and are recognized by a transition in colour, fabric and grain size (Fig. 2d). Feldspar-rich layers (leucosome selvages) are common at the mylonite-ultramylonite contacts. Shear foliation planes (C-surfaces) in the ultramylonite overprint the compositional layering (pre-exisiting C-surface) at a small angle. The compositional layering in the ultramylonite is much thinner, usually less than 1 mm thick. Ultramylonite S-foliation defined by elongate to curviplanar feldspar and pyroxene has the same orientation as that of the host mylonite and is pervasively developed. Shear bands, which do not occur in the host mylonite, are characteristic of the ultramylonite, commonly observed linking C-planes. Cherty ultramylonite comprises bands of cherty- or glassy-looking material up to 5 mm thick that occur sparsely throughout the granulite blocks. Although grossly parallel to the compositional layering and ultramylonite foliation, cherty mylonite bands have extremely sharp contacts that cut both of the latter domains. The distinctive appearance and crosscutting relationships in the field serve as the bases on which this domain is identified. Experimental procedures Thin sections were cut parallel and perpendicular to the lineation and compositional layering, respectively, observed in the host mylonite. All thin sections are doubly polished and mounted in Crystalbond™ adhesive in order to remove areas for transmission electron microscopy. Material was examined at progressively finer scales using light microscopy (LM), scanning electron microscopy (SEM), electron microprobe (EMP) and transmission electron microscopy (TEM). The SEM is JEOL 6400 operated at 15 kV and 2.5 nA. The analytical system combines a JEOL wavelength-dispsersive spectrometer for light elements with a LINK Pentafet™ energy-dispersive detector and analyser. The

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EMP is a JEOL 733 operated at 15 kV and 10 nA. The TEM is a Philips EM400T operated at 120 kV, with a LINK Pentafet™ light element energy-dispersive detector and analyser. Specimens for TEM were mounted on copper grids directly from the thin sections, and ionmilled with argon. For compositional analyses by TEM, samples of coarser-grained granulites were first analysed in SEM and electron microprobe (EMP) and then used as empirical internal elemental standards for EDS analysis by TEM. All electron microscopy and the bulk of compositional analyses referred to herein were done in the UNB Electron Microscopy Unit in order to maintain strict spatial control of the analysis areas; samples analysed by EMP at the University of Wales and University of Calgary gave consistent results (Gibbons et al. 1996). Deformation microstructures

Host mylonite Microstructures (Fig. 3a) within the host (i.e. least deformed) granulite comprise orthopyroxeneclinopyroxene-garnet-feldspar assemblages, with both clinopyroxene and orthopyroxene forming equant, to slightly elongate porphyroclasts (100-1000 urn grains; aspect ratio typically < 2:1) in a finer-grained, dominantly feldspar matrix (25-150 urn grains). Pyroxene grains both contain and are mantled by ilmenite-magnetite aggregates. The plagioclase (An42-An48) grain size range is quasi-bimodal, with smaller grains being the recrystallization products (rotation recrystallization) of the larger porphyroclasts. A few primary grains with incomplete dynamic recrystallization are observed (Fig. 3b). Feldspar grain boundaries are irregular. Dislocations and dislocation networks indicative of intracrystalline creep occur in both pyroxene and plagioclase (Fig. 3c,d). Additionally, pyroxene-oxide aggregates would suggest concurrent diffusive mass transfer (Fig. 3a).

Ultramylonite As was observed in field exposures, the boundary (Fig. 4a) between mylonite and ultramylonite is commonly associated with feldspar-rich layers (leucosome) that reflect a broader compositional change. The plagioclase in these layers is compositionally distinct from that (An27-An39) in the mylonite and where pyroxenes are present, they occur in lower modal proportions than in the host mafic mylonite. The

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Fig. 3. Mylonite microstructures. (a) Typical pyroxene grains in plagioclase matrix. Pyroxene grains are mantled by ilmenite-magnetite (Fe-Ti oxide) aggregates. LM Crossed-polarizers with gypsum plate, (b) Relict primary plagioclase in leucosome exhibiting partial dynamic recrystallization (rotation recrystallization) to mylonitic texture. LM Crossed-polarizers. (c) Dislocation substructure in clinopyroxene. TEM Bright field, (d) Dislocations and twins in plagioclase. Subgrain boundaries indicated by arrows. TEM Bright field.

most common textural transition is an abrupt Zone H: plagioclase aggregates (5-25 jam) decrease in grain size and/or distortion of elonformed by dynamic recrystallization of grains from Zone I; these plagioclase grains form gate pyroxenes and feldspars into a new, finerelongate aggregates parallel to a new compograined foliation, concomitant with formation of sitional layering in the ultramylonite. a strong S-C fabric and shear bands Fig. 4a). Pyroxenes are associated with dramatic elongaZone III: discrete, optically isotropic, 100 urn thick seams with grain sizes routinely less tion of the mantling Ti-Fe oxide aggregates parallel to both S- and C-planes. In detail, the than 1 (im and rarely larger than 5 jim. ultramylonite has a very heterogeneous microstructure that enables a tripartite division The mineral assemblage remains effectively by grain size into three distinct textural zones unchanged throughout all three zones and no (Fig. 4b): systematic variation in mineral chemistries was observed among them. The most significant minZone I: plagioclase and pyroxene are com- eralogical change is the new occurrence of quartz parable in size to those in the host mylonite and a modal increase in K-feldspar, particularly (15-50 jam), but are more elongate; these in Zone III. The latter mineralogical changes, in elongate grains occur within enveloping conjunction with the cross-cutting field relationsurfaces that define zones parallel to the com- ships support the origin of the leucosomes as positionallayering in the host mylonite, but lower crustal anatectic melts. The black colour of are characterized by a more prominent Zone III arises primarily from the extremely S-foliation. small grain size. Semi-quantitative image


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Fig. 4. Ultramylonite microstructures. (a) Contact between mylonite and ultramylonite bands separated by plagioclase selvage. Compositional layering in the mylonite and C-planes (dark seams) in ultramylonite are subparallel. S-foliations in the mylonite become more strongly developed in the selvage, with a strong S-C fabric, plus shear bands in the ultramylonite. LM Crossed-polarizers with gypsum plate, (b) Details of ultramylonite bands. Zone I and Zone II both exhibit dynamic recrystallization, and are effectively a continuous deformation. Zone III are thin optically opaque seams parallel to C-planes. LM Crossed-polarizers with gypsum plate, (c) Zone III ultramylonite band parallel to C-plane demonstrating the very fined grained, polymineralic nature of the optically opaque seams. It is contained within Zone II ultramylonite. SEM Backscattered Image, (d) Dislocation-free K-feldspar grain typical of Zone III ultramylonite. Hole in the upper left is X-ray analysis spot induced during collection of mineral composition data. TEM Bright field.

processing of oxide phases showed there to be no significant variation in oxide proportions at length scales on the order of the thickness of the associated zonal foliation; for example, compositional layering vs. Zone I vs. Zone II. Grains in Zones I and II exhibit dislocation substructures comparable to those in the host mylonite (Fig. 3d). Zone III grains are distinctive by virtue of the absence of dislocations, irrespective of the mineral phase (Fig. 4d), except for occasional, larger feldspar grains. The most commonly observed transition from Zone II to Zone III textures is the apparent continuously ductile reorientation of S-foliations toward the C-plane (Fig. 5a) along which elongate Fe-Ti oxide aggregates and Zone III textures develop.

Close examination shows that Zone III seams (Fig. 5b) initiate parallel to both C-surfaces and shear band orientations as discrete brittle displacement and cataclastic zones (Fig. 5c,d). Zone III ultramylonite bands initially that form parallel to incipient shear bands (Fig. 5c) rotate into the C-plane orientation with increasing strain. The fractures are accompanied by finegrained aggregates (veins) of ilmenite/magnetite, plagioclase, K-feldspar, pyroxene and quartz that have substantively the same assemblage as seen throughout the mylonite and ultramylonite. The association of fracturing and veining leads to interpretation of these textures as fluid-assisted fracture events, where the fluid is related to anatectic melts.

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Fig. 5. Microstructures associated with the transition from Zone II to Zone III ultramylonite, from most to least common. The evolutionary sequence actually progresses in reverse as (d)-(c)-(b)-(a). (a) Typical, highly evolved contact between Zone I and II ultramylonites with well-defined S-fabric delineated by deformed orthopyroxene and truncated by Zone III ultramylonite parallel to C-plane. Backscattered electron contrast for the labelled minerals is the same in (b)-(d). SEM Backscattered Image, (b) Discrete Zone III seam (veinlet) with pyroxene, plagioclase, ilmenite and magnetite with no clear evidence of a transition. SEM Backscattered Image, (c) Elongated pyroxene and Fe-Ti oxides within a fracture cross-cutting the S-foliation along an incipient shear band. SEM Backscattered Image, (d) Fracturing and cataclasis along the shear band orientation associated with introduction of a pyroxene-plagioclase-oxide vein. The fracture and vein cross-cut orthopyroxene that has been previously deformed into the S-foliation. The rarity of these microstructures is interpreted to reflect their rapid transformation to textures such as in (a). SEM Backscattered Image.

Cherty ultramylonite microstructures The cherty ultramylonite bands are as distinctive in thin section as they are in the field. The sharpness of the macroscopic contacts arises from an abrupt and homogeneous reduction in grain size that is more extreme than the mylonite to ultramylonite transition (compare Fig. 6a with Fig. 4b). The contacts, although abrupt, are distinctly curviplanar, commonly have an ultra-fine-grained pyroxene selvage at contacts (Fig. 6a) with the host and truncate the host compositional layering, including individual grains. Flow folds (Fig. 6b) are common and restricted to these bands. Pyroxene porphyroclasts within the bands have two features not

seen in elsewhere: a shard-like morphology, particularly near the contacts and/or irregular grain boundaries suggestive of fluid corrosion (Fig. 6c-e). The absence of evidence for progressive grain size reduction would be consistent with, but not solely demonstrable of, catastrophic grain size reduction (Fig. 6d). Similarities with Zone III ultramylonite include matrix grain size of the order of 1 um (Fig. 6d) and the near complete absence of dislocations (see Fig. 4d), but the 'cherty' bands can be considerably thicker (> 1 mm). Garnet porphyroclasts likewise have distinctive rounded shapes not typical of typical of pseuodotachylyte (Fig. 6f and Lin 1999).


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Fig. 6. Cherty ultramylonite (deformed pseudotachylyte). (a) Cherty mylonite - mylonite contact. Unlike the common ultramylonite, there is no progressive transition from the host mylonite; the contact is both abrupt and eroded. A dark, pyroxene-rich selvage at the contact may be a chill margin. LM Crossed-polarizers with gypsum plate, (b) Typical flow folds that are common in these zones, and totally unobserved elsewhere. LM Plane-polarized light, (c) Highly elongated and distorted pyroxene at the contact with ultra-fine-grained cherty ultramylonite matrix. The corroded pyroxene grain boundary is indicative of interaction with a disequilibrium fluid. SEM Backscattered Image, (d) Pyroxene-matrix contact demonstrating attrition and grain size reduction by brittle processes. SEM Backscattered Image, (e) Overview of a cherty ultramylonite with deformed pyroxene and a foliation that are distinctive from those in the dominant ultramylonite. LM Crossed-polarizers with gypsum plate, (f) Elliptical garnet with extremely smooth contact with the matrix indicative of interaction with a melt. LM Plane-polarized light.

Interpretation of microstructures and deformation mechanisms Deformation environment The syndeformational granulite metamorphic minerals in each microstructural regime have effectively constant mineral chemistries that

correspond to those used in the thermobarometric calculations of Gibbons et al. (1996). The associated temperature range of 700-860 °C and 750-950 MPa is applicable to all the deformation states, with the possible exception of the cherty mylonite. Hence, the microstructures clearly evolved in the deep crust. Preservation of these microstructures can be explained by the

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small grain size and anhydrous mineral phases which will inhibit fluid ingress and also kinetically limit retrogression during exhumation of the rocks. Where there is imperfect preservation, the ability to capture a range of behaviour is potentially more useful than simply assigning some single set of parameters to the rock body (Means 1995).

ing and veining localized on pre-existing localized ductile structures provide an additionally mechanism for grain size reduction. The association of grain fragmentation, introduction of oxide/silicate veins and cataclasis (Fig. 5c,d) argues for operation of a pressuredependent brittle process induced by elevated fluid pressure. Based on the observed mineralogy, the fluid from which veins form is akin to that of the high temperature anatectic leucoDeformation mechanisms some melts in the host granulite. The fractures The bulk of the microstructures observed in and veins are concentrated at zones of ductile both mylonite and ultramylonite reflect disloca- strain localization such as shear bands, with subtion-mediated deformation-recovery-recrystal- sequent solid-state deformation of the crystallization cycles that reduce the grain size. lized melt obliterating any brittle textures. The However, the grain size reduction by demon- formation of elongate Fe-Ti oxide aggregates strable dynamic recrystallization contrasts with during ultramylonite deformation plays a the grain size reduction observed at the tran- central role in the brittle event. Fe-Ti oxides are sition from Zone II to Zone III ultramylonite anticipated to be involved with the vein as remowhere the latter relates to fracturing and vein bilized, extremely ductile phases with low visintroduction. Two aspects of the microstructural cosity behaviour. Similar, though not identical evolution are central to any interpretation of influences of oxides on fracturing has been these observations: (i) the increase in defor- reported for lower ocean crust shear zones mation intensity, for which decreasing grain size (Agar & Lloyd 1994). is a proxy, correlates with concomitant localizThe importance of fracturing, vein introducation of the deformation into narrower zones; tion and cataclasis is the reduction in grain size and (ii) microstructures that correlate to the beyond that obtainable in the dislocation creep initial stages of the localization process can be field by dynamic recrystallization (White 1982; masked by subsequent textural equilibration De Bresser 2001); this in turn allows a transition during ductile flow. to grain-size sensitive (GSS) flow (Figs 7 and 8) The host mylonitic texture is interpreted as and phenomenological superplasticity (Boullier having developed during deep-crustal flow of & Gueguen 1975; Edington etal 1976). Diffusive the granulite, wherein intracrystalline, disloca- control of the GSS flow in Zone III ultramylonite tion-mediated processes dominate, concomitant is inferred from the submicrometre grain size and with development of the macroscopic composi- dislocation-free grains that contrast with dislocational layering. These textures represent the tion dominated microstructures throughout the closest approach in the granulites to a bulk bulk of the granulite. The polymineralic nature quasi-steady-state deformation. Nevertheless, of Zone III ultramylonite favours mutual pinning heterogeneous flow occurs even under these of grain boundaries and suppression of grain conditions in response to the rheological con- growth (White 1982; Panda et al 1985; Chen & trast of pyroxene- and plagioclase-rich layers as Xue 1990; Olgaard 1990); hence, overprinting of shown by both more intense recrystallization the latter of the fine-grained material is not within the plagioclase layers and the associated anticipated even if readjustment to ambient localized development of S-C fabrics. stresses comparable to Zone II ultramylonite Mylonitic microstructures evolve progres- should occur. Thus it is possible to achieve consively to Zone I, then Zone II ultramylonite. trasting mixed-mode deformation in adjacent Incontrovertible evidence for dynamic recrystal- volumes of rocks while requiring only transient lization of plagioclase to grain sizes less than variations in stress and strain rate throughout the about 5 (im is not observed (Fig. 4b) even rock mass. though significantly smaller grain sizes are Cherty ultramylonite is interpreted to be an common in Zone III ultramylonite. Notably, extreme example of localization, whereby a fricboth the reduction in grain size and the ensuing tional melt (disequilibrium melt) develops (Fig. change in intracrystalline defect microstructure 7) with subsequent super- and sub-solidus flow. from Zone II to Zone III ultramylonites are This interpretation rests largely on the gross texabrupt (e.g. Fig. 5a), and there is no evidence for tural differences and obliquity of these zones to the progressive reduction in grain size observed the overall rock fabric, in conjunction with their between Zone I to Zone II ultramylonite. Tran- similarity to demonstrable pseudotachylytes. sient microstructures (Fig. 5) formed by fractur- Unlike the other inferred crystallized melts (e.g.


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Fig. 7. Schematic chart of microstructure evolution during deformation. The diagram illustrates the interplay between strain or strain-rate hardening of the system, the corresponding localization feature formed in response to hardening, and the subsequent strain or strain-rate softening processes within the localized zone that produces the stable microstructure observed in the rocks. Deformation mechanism regimes are demarcated in terms of how localization is controlled during the evolution of the ultramylonites.

leucosome, Zone III ultramylonite), cherty ultramylonite layers entrain pyroxene porphyroclasts that have scalloped contacts suggestive of corrosion by a fluid, and have irregular contacts with the host rock, as is observed in disequilibrium frictional melts (Magloughlin 1992; Spray 1992; White 1996). The flow folding that, within the granulite blocks, is unique to these zones is typical of ductile flow within pseudotachylyte (Sibson 1980; Passchier 1982; Koch & Masch 1992; White 1993, 1996). Upon crystallization of the frictional melt, the extremely fine-grained, polymineralic material moves directly into the grain-size sensitive flow regime, in much the same way as Zone III ultramylonite. The intensity of deformation immediately adjacent to these zones (e.g. very high aspect ratio pyroxene grains indicative of glide with insufficient time for recovery), in combination with the inherent strength of granulites in a pore fluid pressure deficient environment (White 1996; Kiister & Stockhert 1999) is consistent with melting initiated as a plastic instability; this in turn would explain the contrast in features between the 'cherty' and Zone III ultramylonites which initiated as transient high fluid pressure events. Similar textures have been observed in peridotite massifs (Vissers et al

1997) and are interpreted similarly as 'dry' localization events, in contrast to 'wet' localization, where water or melts play a role.

Mechanism transitions The processes leading to microstructural and/or deformation mechanism transitions in the Clarke Head granulites can be subtle. Inferring the absence of dislocations in small equant grains within high-strain zones to be an indicator of diffusion-controlled grain size sensitive flow, a transition in the dominant deformation regime occurs at about 5 um grain size (Fig. 8); that is, Zone II and Zone III ultramylonites are distinct. Additionally, grain size reduction by dislocation processes (dynamically recrystallization) is observed not to provide a path from grain sizes characteristic of Zone II to those of Zone III ultramylonite. Instead, the rarely observed and rarely preserved cases of fracturing, cataclasis and vein injection are responsible for strain localization and grain size reduction. Brittle deformation as an overprinted precursor to ductile deformation is commonly observed in experiments (e.g. Tullis & Yund 1987) and nature (Pennacchioni & Cesare 1997; Guermani & Pennacchioni 1998). The mechanism regime path then becomes:

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Fig. 8. A plot of the width of localized deformation zones (black bars) and the corresponding grain size of plagioclase within specific types of zones (shaded area). The localization zones are arrayed sequentially in terms of the degree to which they would exhibit a strain rate faster than the bulk rate; i.e. perfectly homogeneous deformation would give a ratio of 1.0 (see text). The figure illustrates that as the characteristic length scale of the localization decreases, so does the grain size; that is, the more discrete the localization, the higher the stress, and as interpreted here, the higher the strain rate during initiation of the localization. Based on microstructural observations, approximate boundaries are drawn between dislocation- and brittle-mediated localization, as well as the limit of dynamic grain size observed. It is suggested by these constraints that those ultramylonites observed in these rocks with characteristics of GSS flow required a brittle precursor. This is not to infer that this is an attribute for all examples of GSS flow.

dislocation-mediated creep —> fracture —> diffusion-mediated GSS creep. Deformation is sustained within the GSS creep field by the suppression of grain growth due by mutual pinning of the different. If there was some grain growth, it would be limited by the 'stable' recrystallized grain size appropriate for the dislocation creep regime (White 1982). These observations emphasize the possibility of transitional microstructures, for which little evidence is preserved, acting as necessary intermediary states between outwardly stable, but contrasting microstructures.

Deformation partitioning and localization Within the Clarke Head granulites, the localization of strain corresponds to progressive reductions in grain size. Development of discrete C-plane fabrics and shear bands reflects mechanical partitioning of the strain in response to inherent system hardening and/or a change in

boundary conditions. The contributory effects to the final microstructures of systemic hardening and localized softening are illustrated in Fig. 7. Strain hardening of an overall deforming system - a result of changes in kinematics, P-T fluid environment or fabric, where no one factor is independent of the other - will lead to increasing stress levels if imposed tectonic displacement rates continue. For continued accommodation of such displacements, there must eventually develop some mode of strain or strain-rate softening. In effect, system hardening leads to instability and localization at more extreme (e.g. higher stress) conditions than the pre-existing deformation state, while subsequent softening within the localized zones masks the processes of localization; softening can not be isolated from induced instability and the localization of deformation (as reviewed for experimental systems byHobbsetal. 1990). Microstructural steady state or stability during deformation must meet the criteria of


Means (1981) that both weakening and softening be equally effective. The latter will be reflected as microstructures associated with a controlling micromechanism(s) (e.g. dislocation glide + climb) and which describe the deformation over a characteristic spatial and time scale. At any stage of a dislocation-mediated microstructural evolution, maintenance of an apparent steady-state microstructure is tied to grain size reduction (localization) events that enable sufficient softening to establish a degree of microstructural stability. Trigger events (or bifrucations) arising from system hardening provide the impetus for localization, and associated softening, i.e. in the absence of perfect spatial and temporal homogeneity, heterogeneous behaviour will be the rule. This is clearly demonstrated by the requirement for gradients in strain energy at the grain scale in order to induce dynamic recrystallization, even in nominally steady-state fabrics. The rarity with which these trigger events are preserved reflects the ubiquity of subsequent ductile overprints. The question of localization is a function of the length scale of interest and, in turn, a function of the prevailing defect structure. Central to this is the question of what parameters can be extracted from the rock record (e.g. palaeopiezometers), and to what volume of rock these are applicable. Within the Clarke Head granulites, each type of localization structure has its own characteristic length scale (thickness), plus a length scale (grain size) specific to its specific microstructure. Figure 8 relates the thickness of localized deformation zones to an estimate of the associated deviation of localized strain rate from the system or bulk strain rate. Ductile shear strain rate is partitioned as a function of zone thickness for a set ambient strain rate (White & Mawer 1992); for example, perfect homogeneity and steady-state at a bulk strain rate of 10~14 s"1 is taken as a ratio of 1.0; mylonites ~ 102; Type II ultramylonite « 104, etc. Typical frictional displacement rates are used to give a comparable strain rate estimate in the brittle regime. Each level of localization corresponds to a decrease in length scale (thickness of localized zone), an increase in the number of localized sites (e.g. foliation => foliation with multiple shear bands) and an increase in deformation intensity within that zone (smaller grain size). The variability of the thickness of the cherty ultramylonite distinguishes it from the other localized deformation zones, consistent with it forming by a unique process (pseudotachylyte). In cases of localized (heterogeneous) defor-

35

mation, there may not be a single answer to questions of strain rate and stress as determined from microstructures. During a phase of hardening, localization of more-or-less bulk mylonitization into discrete C-planes, suggests that the weak zones (C-planes) have the highest strength (smallest recrystallized grain size). If instead we couch questions in terms of how displacements are best accommodated by the rock mass, then the weak zones do carry evidence of a higher stress, but also higher strain rates distributed over smaller volumes of rock; that is, the timeaveraged rate of strain energy dissipation may stay more of less constant, whereas the path of displacement accommodation varies dramatically. Rutter (1999) has noted the need for sufficient strain rates in the localized zones to balance the bulk displacements. At a plate length scale, deformation rates of 10~14 s"1 could be appropriate, and the related microstructures could be non-descript. At the same time localization due to system hardening will record only stresses and strain rates greater than background. What are ultimately interpreted as the weakest zones may well be those with the initially fastest rate of strain hardening i.e. rock volumes within which stresses concentrated most quickly lead to a trigger that transforms microstructures to that characteristic of the new deformation regime.

Conclusions Sequences of mylonites and ultramylonites have granulite grade mineral assemblages and associated deformation microstructures for which thermobarometry indicates deformation at temperatures in excess of 700 °C and 750 MPa. In addition to high-temperature creep and grain size reduction by dynamic recrystallization, there is localization of deformation into discrete, but always narrower zones that are tied to introduction of a fluid phase (leucosome, veins, pseudotachylyte). Elevated fluid pressure transients (anatectic melts) that develop in concert with system hardening introduce material that is more amenable to deformation (strain softening) than the existing constitutive state. In extreme cases, frictional melting produces pseudotachylyte that is subsequently deformed ductily by grain size sensitive flow upon crystallization. The oxide-related localizations, as well as the more obvious pseudotachylyte events, may be coseismic events. The array of strain localization structures and concomitant variations in microstructures demonstrate the inherent heterogeneity of deformation, wherein characteristic length scales must be considered

36

J. C. WHITE

before assigning evidence from the rock record, such as palaeopiezometric measurements, to bulk behaviour of the lithosphere. The overprinting of microstructures associated with the initial localization appears to be common. Knowledge that such microstructural evidence has been lost may allow reconciliation of contradictions in the rock record, such as the production of ultra-fine-grained superplastic aggregates in what otherwise appears to be a dominantly coarser grained dislocation creep regime. Instrumentation and analytical facilities were supported through the UNB Electron Microscopy Unit. Financial support by the Natural Sciences and Engineering Research Council is gratefully acknowledged. Collaboration with Terry Gordon, Brendan Murphy and particularly Wes Gibbons, who provided the impetus for this work, has contributed significantly to these ideas, errors in which are solely the author's responsibility. Comments by R. Holdsworth and an anonymous reviewer contributed to clarification of the presentation.

References AGAR, S.M. & LLOYD, G.E. 1994. Rheology of shear zones in the lower ocean crust: the role of oxide deformation. EOS Transactions of the American Geophysical Union, 75/44, 650. BOULLIER, A.M. & GUEGUEN, Y. 1975. SP-mylonites: origin of some mylonites by superplastic flow. Contributions to Mineralogy and Petrology, 50, 93-104. CHEN, I.W. & XUE, L.A. 1990. Development of superplastic structural ceramics. Journal of the American Ceramic Society, 73, 2585-2609. DE BRESSER, J.H.P., TER HEEGE, J.H. & SPIERS, CJ. 2001. Grain size reduction by dynamic recrystallization: can it result in major rheological weakening? International Journal of Earth Sciences, 90, 28-45. EDINGTON, J.W., MELTON, K.N. & CUTLER, C.P. 1976. Superplasticity. Progress in Materials Science, 21, 63-170. GIBBONS, W. & MURPHY, J.B. 1995. Mylonitic mafic granulite in fault megabreccia at Clarke Head, Nova Scotia: a sample of Avalonian lower crust? Geological Magazine, 132, 81-90. GIBBONS, W., DOIG, R., GORDON, T., MURPHY, B., REYNOLDS, P. & WHITE, J.C. 1996. Mylonite to megabreccia: tracking fault events within a transcurrent terrane boundary in Nova Scotia, Canada. Geology, 24,411-414. GUERMANI, A. & PENNACCHIONI, G. 1998. Brittle precursors of plastic deformation in a granite; an example from the Mont Blanc Massif (Helvetic, Western Alps). In: RUTTER, E.H., BORIANI, A., BRODIE, K.H., BURLINI, L. & TREAGUS, S.H. (eds) Structures and properties of high strain zones in rocks. Journal of Structural Geology, 20,135-148.

HOBBS,B.E.,MULHAUS,H.B.& ORD,A. 1990. Instability, softening and localization of deformation. In: KNIPE, R.J. & RUTTER, E.H. (eds) Deformation mechanisms, rheology and tectonics. Geological Society, London, Special Publications, 54, 143-165. KEPPIE, ID. 1982. The Minas Geofracture. In: ST. JULIEN, P.& BELAND, J. (eds) Major structural zones and faults of the Northern Appalachians. Geological Association of Canada, Special Paper, 24, 263-280. KOCH, N. & MASCH, L. 1992. Formation of Alpine mylonites and pseudotachylytes at the base of the Silvretta nappe, Eastern Alps. Tectonophysics, 204, 289-306. KUSTER, M. & STOCKHERT, B. 1999. High differential stress and sublithostatic pore fluid pressure in the ductile regime - microstructural evidence for short-term post-seismic creep in the Sesia Zone, Western Alps. Tectonophysics, 303,263-277. LIN, A. 1999. Roundness of clasts in pseudotachylytes and cataclastic rocks as an indicator of frictional melting. Journal of Structural Geology, 21, 473-478. MAGLOUGHLIN, J.F. 1992. Microstructural and chemical changes associated with cataclasis and frictional melting at shallow crustal levels: the cataclasite B pseudotachylyte connection. Tectonophysics, 204,243-260. MEANS, W.D. 1981. The concept of steady-state foliation. Tectonophysics, 78,179-199. MEANS, W.D. 1995. Shear zones and rock history. Tectonophysics, 247,157-160. OLGAARD, D.L. 1990. The role of second phase in localizing deformation. In: KNIPE, R.J. & RUTTER, E.H. (eds) Deformation mechanisms, rheology and tectonics. Geological Society. London, Special Publications, 54,175-181. PANDA, PC, RAJ, R. & MORGAN, P.E.D. 1985. Superplastic deformation in fine-grained MgO.2Al2O3 spinel. Journal of the American Ceramic Society, 68, 522-529. PASSCHIER, C.W. 1982. Pseudotachylyte and the development of ultramylonite bands in the SaintBarthelemy massif, French Pyrenees. Journal of Structural Geology, 4, 69-79. PENNACCHIONI, G. & CESARE, B. 1997. Ductile-brittle transition in pre-Alpine amphibolite facies mylonites during evolution from water-present to water-deficient conditions (Mont Mary Nappe, Italian Western Alps). Journal of Metamorphic Geology, 15, 777-791. RUTTER, E.H. 1999. On the relationship between the formation of shear zones and the form of the flow law for rocks undergoing dynamic recrystallization. Tectonophysics, 303,147-158. SIBSON, R.H. 1980. Transient discontinuities in ductile shear zones. Journal of Structural Geology, 2, 165-171. SPRAY, J.G. 1992. A frictional basis for the frictional melting of some rock-forming minerals. Tectonophysics, 204, 205-221. TULLIS, J. & YUND, R.A. 1987. Transition from cataclastic flow to dislocation creep of feldspar:

INSTABILITY AND DEFORMATION LOCALIZATION mechanisms and microstructures. Geology, 15, 606-609. VISSERS, R.L.M, DRURY, M.R., NEWMAN. J. & FLIERVOET, T.F. 1997. Mylonitic deformation in upper mantle peridotites of the North Pyrenean Zone (France): implications for strength and strain localization in the lithosphere. Tectonophysics, 279, 303-325. WALDRON, J.W.F., PIPER, DJ.W. & PE-PIPER, G. 1989. Deformation of the Cape Chignecto pluton, Cobequid highlands, Nova Scotia: thrusting at the Meguma-Avalon boundary. Atlantic Geology, 25, 51-62. WEBSTER, T.L., MURPHY, IB. & BARR, S.M. 1998. Anatomy of a terrane boundary: an integrated structural, geographic information system, and remote sensing study of the late Paleozoic Avalon-Meguma boundary. Canadian Journal of Earth Sciences, 35, 787-801.

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WHITE, J.C. 1982. Quartz deformation and the recognition of recrystallization regimes in the Flinton Group conglomerates, Ontario. Canadian Journal of Earth Sciences, 19, 81-93. WHITE, J.C. 1993. Melt instabilities and superplasticity during ductile shear of silicates. Proceedings of the Microscopical Society of Canada, 20,118-119. WHITE, J.C. 1996. Transient discontinuities revisited: pseudotachylyte, plastic instability and the influence of low pore fluid pressures on deformation mechanisms in the mid-crust. Journal of Structural Geology, 18,1471-1486. WHITE, J.C. & MAWER, C.K. 1992. Deep-crustal deformation textures along megathrusts from Newfoundland and Ontario: implications for microstructural preservation, strain rates and strength of the lithosphere. Canadian Journal of Earth Sciences, 29,129-141.


Microstructural evolution in a mylonitic quartz simple shear zone: the significant roles of dauphine twinning and misorientation GEOFFREY E. LLOYD School of Earth Sciences, The University, Leeds LS2 9JT, UK Abstract: SEM/EBSD-based orientation and misorientation analyses are described for a lower amphibolite facies simple shear zone (Torridon, NW Scotland). It is shown that as well as conventional crystal-slip processes (i.e. basal-a, prism-a, rhomb-a and negative second order rhomb-a slip), dauphine twinning also plays a role in both microstructural and petrofabric evolution. Twinning assists in the initial grain size comminution processes, including dynamic recrystallization, from originally coarse wall rock grains to a typical mylonitic microstructure in the centre of the shear zone. Subsequently, twinning helps to accommodate high shear strains in the mylonite whilst maintaining a stable microstructure and constant 'single crystal' petrofabric. The role of dauphine twinning appears to be to allow efficient switching between relatively 'soft' and relatively 'hard' slip directions that possibly exploit a distinction between negative and positive crystal forms. Misorientation analysis emphasizes the relationships between crystal-slip systems and grain boundary network, including dauphine twin planes, and suggests that the mylonitic microstructure contains preferred orientations of both tilt and twist boundaries that help to explain shear zone microstructural evolution and stability.

Shear zones occur in many deformed rocks on scales ranging from the microscopic to the macroscopic. They act to localize deformation and accommodate much of the ductile displacement that occurs during orogenesis. Consequently, the physical, mechanical and chemical processes involved in the formation and evolution of shear zones have attracted considerable and wide-ranging attention (e.g. Ramsay 1980; Rutter et al. 1998). In particular, many shear zones develop strong crystal lattice preferred orientation (LPO) during formation. However, the processes by which LPO develops during shear zone evolution remains controversial, but it must be possible to derive this knowledge from study of shear zone microstructure (e.g. Law 1987). Until recently, the only techniques widely available for studying microstructural and LPO evolution have been universal stage optical microscopy and X-ray texture goniometry. Although both techniques remain useful, they have their limitations (e.g. the former is restricted to minerals for which unique crystal directions can be identified optically, whereas the latter is restricted to a specific grain size range and derives only bulk data lacking microstructural control). Fortunately, recent developments in scanning electron microscopy (SEM), specifically electron-backscattered diffraction (EBSD), have provided solutions to these problems. SEM/EBSD (Venables & Harland 1973) permits the recognition of all crystal orientations for most minerals in one-

to-one relationship with microstructure, and automation provides statistically significant datasets. This contribution assesses the question of shear zone microstructural evolution and LPO development from the point of view of an almost perfectly simple shear zone developed within a quartz vein, using SEM/EBSD to derive the crucial observations and data.

Specimen and analytical details Sample details The sample used in this study (Fig. 1) was collected (Wheeler 1984) from a 30cm wide deformed quartz vein within Proterozoic gneisses at the head of Upper Loch Torridon, NW Scotland (UK GR NG 840530). The shear zone developed under lower amphibolite facies conditions typical of the mid-lower crust. Law et al. (1990) have previously described this sample in terms of its finite microstructure and LPO, as follows (see also Lloyd et al. 1992). The vein was deformed by crystal plastic processes to form a dextral shear zone with an intense mylonitic foliation and lineation (e.g. Ramsay & Graham 1970). In XZ section (i.e. where X>Y>Z), the vein microstructure is that of a classic Type II S-C mylonite (Lister & Snoke 1984) and consists of two planar domains (A and B) aligned parallel to the macroscopic mylonitic foliation (SA). A domains consist of


40

G. E. LLOYD

MICROSTRUCTURAL EVOLUTION

equant, less than 5 um quartz grains and subsidiary feldspar, whereas (the volumetrically more important) B-domains consist of dynamically recrystallized, elongate, less than 100 um quartz grains with long axis (SB) oblique to SA. The obliquity between SB and SA confirms a dextral shear sense. Shear zone microstructures (see Law et al. 1990, fig. 2) are indicative of constant volume, strongly non-coaxial, essentially plane strain deformation that closely approximates to simple shear. The development of a typically mylonitie microstructure is achieved via crystal slip systems that exploit the ease of slip parallel to (e.g. basal-a, rhomb-a and/or {±7i}), in accordance with a simple shear kinematic framework that aligns with lineation (X) and (c) normal to foliation (XY). The LPO (e.g. Fig. 2) therefore can be explained in terms of the simple shear kinematic framework indicated by shear zone geometry, in agreement with the fabric evolution model for bulk simple shear proposed by Etchecopar (1977). Furthermore, it appears that slip is favoured on negative rather than positive forms, which might suggest a lower resistance to slip on the former compared to the latter. Alternatively, the non-equivalence of positive and negative rhombs and the r point maximum in XZ at 30-50° to XY, subparallel to the NW-SE trending maximum principal stress direction inferred from the simple shear kinematic framework, is consistent with dauphine twinning (e.g. Tullis & Tullis 1972). The significance of dauphine twinning during quartz deformation has received relatively little attention (e.g. Tullis 1970; Baker & Wenk 1972; Baker & Riekels 1977; Olesen & Schmidt 1990; Lloyd et al. 1991; Barber & Wenk 1991; Mainprice et al. 1993; Neumann 2000; Lloyd 2000). The role of dauphine twinning during the microstructural and LPO evolution of the Torridon quartz simple shear zone is considered in detail in this contribution.

SEM/EBSD This contribution makes use of several complementary aspects of SEM crystallographic analysis, namely EBSD (e.g. Prior et al. 1999)

41

and orientation contrast imaging of specimen microstructure (e.g. Lloyd 1987; Adams et al. 1993; Trimby & Prior 1999). Although EBSD is used primarily to index crystal orientations, it can also provide several simulated images of samples (e.g. Jensen & Schmidt 1991; Adams et al. 1993; Field 1997; Lloyd 2000). Here, use is made of the grey-scale coded band contrast (BC) image (e.g. Fig. Ib), which reflects variations in the quality of the electron backscattered diffraction patterns (EBSP). EBSP are indexed via computer pattern recognition programs (e.g. Schmidt & Olesen 1989), which provide conventional petrofabric diagram representations and their misorientation equivalents. Several general or detailed automated and manual SEM/EBSD analyses were performed (see Fig. la for locations and Table 1 for details). In general, automated EBSD indexing success rates were poor and a dependence on both grid step size and sample grain size was observed. As the former increased and/or latter decreased, indexing success rate also decreased due to the increased probability of sampling 'artefacts' (i.e. grain boundaries, 'overlapping' grains, fractures, etc.). Particular care was taken to avoid the incorrect indexing of certain EBSP due to 'pseudosymmetry' effects (e.g. Prior et al. 1999; Lloyd 2000). Automated SEM/EBSD analysis of quartz is especially prone to this problem because the dauphine twinning operation, a 180° (crystallographically equivalent to 60°) rotation about the c-axis, results in very similar pattern configurations that are readily confused. This problem typically leads to an apparent increase in frequency of dauphine twinning in the SEM/EBSD LPO. However, as both Law et al. (1991) and Lloyd et al. (1992) recognized a real contribution from dauphine twinning in the mylonitie LPO in their studies using different techniques, it was important to minimise any mis-indexing due to pseudo-symmetry. The approach used here involved reducing the overall pattern contrast, such that the background diffraction bands that contain the significant differences in configuration are emphasized relative to the common main (brightest)

Fig. 1. SEM images of Torridon shear zone (XZ section plane), (a) Electron channelling orientation contrast image of the whole sample: SZWR: shear zone wall rock (analysis G 030600, see Table 1 for details); SZM: shear zone margin (G040800); MM: mature mylonite (G270700); and MD: mylonite detail (G050800). Also indicated are the positions of the detailed analyses (see Figs 5 & 6). (b) The same image based on EBSP band contrast. The grey scale reflects the quality of the EBSP at each analytical point (dark, poor; bright, good). This image clearly reveals significant detail of the microstructure (e.g. quartz grain size reduction and foliation development and deflection into the shear zone; break-up of feldspar; twinning and subgrains in wall rock feldspar; etc.).

Fig. 2. Equal area, upper hemisphere quartz pole figures for different parts of the Torridon shear zone constructed via auto-EBSD analysis (see Table 1) using the program Pf2k (D. Mainprice; see also Mainprice 1990, Mainprice & Humbert 1994). (a) Shear zone wall rock (SZWR). (b) Shear zone margin (SZM). (c) Mature mylonite (MM), (d) Mylonite detail (MD).


43

Table 1. Summary ofTorridon shear zone auto-EBSD experiments (see Fig. la for locations; BC: band contrast; BS: band slope; BN: band number; n/a, not available). In all but G030100, plagioclase was ignored. JOB (format: username-day-month-year)

DETAILS G030600 Specific comments Area analysed

* blown filament shear zone wall rock (Fig. la) Dimensions (mm) 10.0 X 13.0 No. data pts. 136640 Grid step (urn) 50 Time (hr:min:sec.) 41:47:25 Index rate (sec/ebsp) 1.1 Minerals indexed quartz & plagioclase (Anl6) % Low BC *46.6 % Low BS 0.0 % Low BN 0.1 % Not indexed 27.0 % Indexed 26.2 No. indexed 35800

G040800

G270700

G050800

G120702

user break shear zone margin (Fig. la) 17.0 X 6.0 120537 25 29:37:52 0.88 quartz

_ mature mylonite (Fig. la) 13.5 X 7.25 228825 20 41:18:56 0.65 quartz

_ mylonite detail (Fig. la) 1.5 x 1.0 150001 1 22:11:31 0.53 quartz

*blown filament linear traverse (Fig. 6) 0.1 X 4.0 334087 1 69:06:00 0.36 quartz

2.2 0.0 13.0 68.5 16.2 19527

5.1 0.0 0.2 87.2 7.5 17162

1.3 0.0 2.1 74.8 21.8 32700

n/a n/a n/a 27.49 72.51 290040

diffraction bands (Lloyd 2000). In addition, a relatively tight maximum 'mean angular deviation' (i.e. the difference between observed and predicted pattern configurations) of 0.75 was used (e.g. Krieger-Lassen 1996). Unfortunately, this meant that some 'good' patterns were rejected, further reducing successful indexing rates (see Table 1). In spite of these considerations, indexing rates were sufficient to simulate microstructural detail accurately (e.g. compare Fig. la, b) and to derive representative LPO (see Figs 2 & 6).

Crystallographic misorientation analysis The advent of SEM/EBSD analysis has focused attention on the crystallographic misorientation relationships that exist between adjacent regions (i.e. grains, subgrains, twins, etc.) and therefore upon the formation and crystallographic orientation of the separating boundaries (e.g. Lloyd et al 1997; Trimby et al 1998, 2000; Paul & Fitz Gerald 1999; Fliervoet et al. 1999; Prior 1999; Neumann 2000; Wheeler et al. 2001). The misorientation petrofabric or LPO of deformed rocks may prove significant as the (crystallographic) properties and impact of grain boundaries, particularly where they differ from the grain properties, are largely unknown (see Lloyd & Kendall, in press). Although misorientation data can be dis-

played simply as histograms of misorientation angle frequency, the misorientation between two adjacent crystal lattices is more accurately represented by a single axis about which one of the lattices needs to be rotated by a specific angle to bring it into complete coincidence with the other lattice. This leads to the concept of the misorientation axis/angle pair (e.g. Pospiech et al. 1986; Randle & Ralph 1986; Randle 1992, 1993; Mainprice et al. 1993; Lloyd et al. 1997; Kruse et al. 2001). Any misorientation can be defined by a number of symmetrically equivalent axis/angle pair combinations but current convention is to adopt the pair with the minimum misorientation angle (e.g. Mainprice et al. 1993; Lloyd et al. 1997). Misorientation axis/angle pair data can be represented by either pole figures or inverse pole figures. The former plots the data in 'sample space', whereas the latter plots the data in 'crystal space'. Both provide useful but different information on the nature and origin of (sub)grain boundary misorientations (Lloyd et al. 1997, Wheeler et al. 2001). Misorientation pole and inverse pole figure diagrams are typically presented as series in which each individual figure covers a narrow range of misorientation angles. However, due to problems in determining misorientation axes accurately as misorientation angles approach zero, plots of misorientations less than 5° are usually excluded (Prior 1999).

44

G. E. LLOYD

3b), referred to as the 'forbidden region' (Wheeler et al. 2001). The maximum misorientation angle possible is 104.5° about {2021}. Quartz exhibits also enantiomorphism and its left- and right-handed forms cannot be distinguished via electron diffraction techniques due to Friedel's law, which imposes an arbitrary centre of symmetry. Current convention therefore decrees that a right-handed pseudo-hexagonal definition is used for misorientation analyses of polycrystalline quartz studies (e.g. Neumann 2000). However, for individual quartz grains, EBSD does distinguish positive and negative forms (e.g. , {r} and {z}, [n] and {^'}, etc.), such that an unambiguous crystallographic indexing of quartz to its trigonal symmetry is possible. The possibility exists therefore that a trigonal construction may be more appropriate for quartz misorientation analysis (see below and also Mainprice et al. 1993).

Grain boundary (mis)orientation analysis

Fig. 3. Quartz crystallography, (a) Principal directions in quartz (after Linker et al. 1984) and comparison of the 60° hexagonal (darker shading) and 120° trigonal (lighter shading) inverse pole figure constructions used in misorientation analysis, (b) 60° misorientation inverse pole figure construction and the definition of 'forbidden regions' that develop progressively for misorientation axes above 60° (indicated by lower and upper bound angles). This construction fails to distinguish between or .

The range of possible misorientation angles is determined by the symmetry of the mineral concerned. Quartz has trigonal D3 - 32 symmetry but for convenience often has assumed to be pseudo-hexagonal (Fig. 3a). The former requires that the basal plane subtends a 120° angle with the c-axis, whereas the latter requires only a 60° unit triangle. The hexagonal definition imposes a constraint that misorientation angles above 60° lie progressively further from the c-axis (compare upper and lower bounds in Fig.

A full definition of grain boundary orientation involves the dimensional orientation of the boundary plane as well as the crystallographic misorientation across the boundary (e.g. McLaren 1986). It is difficult to measure accurately and rapidly the dimensional orientations of boundaries (e.g. Kruhl & Peternell 2002), particularly in the large quantities now being produced regularly for misorientation data via EBSD analysis. Fortunately, Randle (1992) has shown that the misorientation axis/angle pair definition alone often provides a meaningful approximation. In addition, the operation of crystal-slip (and twin) systems can be associated with the formation of different types of boundary in specific orientations (Lloyd et al. 1997; Neumann 2000; Kruse et al. 2001). The procedure for predicting boundary orientation from misorientation axis/angle pair data and slip/twin systems is as follows. Each crystal-slip system has a unique rotation axis (R) about which individual crystal directions migrate or disperse on small/great circles during slip (Lloyd & Freeman 1991, 1994). In contrast, different types of boundary have specific misorientation axes (i). The misorientation axis for tilt boundaries (-^TIB) is parallel both to the boundary plane (Fig. 4a) and also R (Fig. 4b). Such boundaries therefore form parallel to jR/^TiB and the slip plane normal (SPN) but normal to the slip direction (SD). The relationships between crystal-slip systems recognized for the Torridon shear zone (Law et al. 1990) and the predicted orientations of misorientation axes and tilt boundaries are given in Table 2.

45


requires at least two slip systems with different Burgers vectors (e.g. Hull & Bacon 1984), these ideal relationships cannot be sustained. In his study of recrystallization in quartzites, Neumann (2000, fig. 2b) included cross-slip systems (e.g. [m} and {z}) in his application of the edge dislocation model and predicted misorientation axes by combining the orientations appropriately (i.e. £{1122} and ^'{1132} respectively; see Fig. 3). His approach suggests a potential modification of the tilt boundary-edge dislocation to twist boundaries. Twist boundaries may form as a compromise to the combinations of slip systems responsible. Thus, -6TWB is most likely to form parallel to the Fig. 4. Relationships between crystal slip systems resultant of the slip plane normal directions (i.e. and boundary orientations: SD: slip direction; SPN: SPN), whereas the boundary plane is most likely slip plane normal; R: slip system rotation axis; £TlB to form_parallel to the resultants of the rotation and -^TwB? tilt and twist boundary misorientation axes, axes (R) and slip directions (SD) and hence (a) Tilt boundaries (TiB). (b) General relationship normal to SPN/£TwB. For example, if (c) and between TiB and crystal slip systems, (c) Twist (c)<m> slip contribute equally to twist boundary boundaries (TwB). (d) General relationship between formation, then SD = + <m> = {2130}, R = TwB and crystal slip systems. The actual situation is <m> + = {2130} and SPN/£TwE = (c) + (c) = more complex - see text for discussion. (0001). The relationships between crystal slip systems recognized in the Torridon shear zone (Law et al. 1990) and the orientations of twist The situation for twist boundaries is more boundaries predicted using this approach are complex because the relationship envisaged given in Table 3. It may therefore be possible to reconcile slip between slip systems, misorientation axes and boundary formation strictly applies only to slip systems and rotation axes with formation of due to edge dislocations. By definition, the mis- both tilt and twist boundaries via a combination orientation axis for twist boundaries (-^TWB) *s of EBSD-based LPO and misorientation analynormal to the boundary plane (Fig. 4c), which ses. This is the approach taken here, which also should be parallel to the slip plane and hence SD highlights the role of dauphine twinning in and R, whereas -€TWB *s parallel to SPN (Fig. 4d). boundary formation during shear zone However, as the formation of twist boundaries microstructural evolution.

Table 2. Relationships between specific quartz crystal slip systems (SPN: slip plane normal; SD: slip direction; R: unique rotation axis) recognized in the Torridon shear zone in both symbol and Miller-Bravais forms (after Linker et al. 1984), tilt boundary misorientation axes (-C-Tis) and potential crystallographic orientation of tilt boundaries. Orientations of boundaries that may be due to dauphine twinning are indicated by shading. Crystal-slip systems Symbol

Miller-Bravais {hkil} [UVTW]

SPN

SD

SPN

SD

(c) (c) {m} W [A M

{m}

(0001) (0001) {10-10} {1-102} {10-11} {01-11}

[11-20] [1-100] [1-210] [11-20] [1-210] [2-1-10]

Tilt boundary orientations R/^TiB

M

W

M M M

SPN

SD

(c) (c)

(m} M M M

{»}

Table 3. Relationships between quartz, crystal-slip systems (R: unique rotation axis; SD: slip direction; SPN: slip plane normal) recognized in the Torridon shear zone (see Table 2) and twist boundary mis orientation axes (&TWB)> assuming slip systems operate equally in combinations (see also Neumann 2000). Also shown (first row) is fcj + (c){m} combination, although latter has not been recognized in the shear zone. The top line in each row gives individual orientations of R, SD and SPN/£TwB for each combination of slip systems read horizontally and vertically. Other lines in each row give resultant crystallographic orientations ofR, SD and SPN/£TwB. Slip system Slip system

(c) [±a]

R

W

SPN

R

SD

[m\ + [±a] + <m> {2130} (2130)

SPN

R

SD

-P-TwB

£-TwB

(c) [m]

{^} [±a]

(ml [±«]

M [±«]

SPN

SD

•^TwB

(c)

-

-

-

-

SPN -P-TwB

_

-

(m][±a]

{m} + (c) r{1011} z{0111}

[1120]

(c) +{m] r{1011) z{0111)

-

-

-

-

M[±«]

W +W {2021}

[1120]

(c) + M (0113)

(c) +(/•) ^(1012)

[1120]

{m} + {*} z{0111}

-

M[±«]

{m} + {*} z{0111)

[1120]

(c) +1/1 7r(1012[

(c) +(*} {1013}

[1120]

{m} + {r} {2021}

w{1123} +w

[1120]

{n'} + {r} {1123}

fz)[±a]

{jiijr{1011) +w

[1120]

7f|0112)

(c) +{^| (1013)

[1120]

{m} + {z} {0221}

W + {^1 {2023}

[1120]

{n'} + {z} {0123}

(c) +M

R

_

-

_

-

-

{n} + {n'} {1124}

[1120]

{r} + {z} ^{1122}


Results and interpretations Microstructure and LPO Optical (Law et al. 1990, figs 1 & 2), orientation contrast (Fig. la; see also Lloyd et al, 1992, figs 2-6) and EBSD-simulated microstructural images (Fig. Ib) show a rapid development of shear zone microstructure, including foliation development, grain size reduction and mylonitization, with fractured clasts of feldspar floating in a quartz mylonite. The shear zone strain gradient and dynamic recrystallization rates can be expected to have been concomitantly steep and/or rapid. The mylonite consists of highly flattened regions that appear to represent dynamically recrystallized original grains (Lloyd et al. 1992) and it is likely that these will exhibit similar crystal orientations or orientation relationships due to a parental grain orientation control during recrystallization (e.g. Lloyd & Freeman 1991,1994). Comparison of LPO from different regions within the shear zone (Fig. 2) indicates a relatively rapid migration of wall rock crystal pole directions towards those exhibited by the mature mylonite LPO, in agreement with microstructural observations. The mature mylonitic LPO exhibits an maximum parallel to X and a (c) maximum normal to the basal/XY foliation planes and approximates to that of a single crystal orientation. This observation, frequently observed in mylonites, poses the question as to how do these LPO and associated microstructures develop and persist. SEM/EBSD analysis reveals considerably more information on LPO and hence microstructural evolution than do conventional techniques. Such information is crucial to the understanding of this and similar shear zones. Close examination of the region where the shear zone margin transforms rapidly into the mature mylonite (e.g. Fig. 1) reveals a distinctive microstructure of narrow intragranular 'bands', comprising 'grains', that are being deflected progressively into the mylonitic foliation (see also Lloyd et al. 1992, figs 2-4). Manual EBSD analysis of 'grains' from one 'band' (e.g. Fig. 5a) identifies the crystal orientation relationships expected for dauphine twinning (Fig. 5b, d), such as six rather than three pairs of rhomb orientations. Furthermore, examination of the orientation relationships between adjacent bands (e.g. Fig. 5a, points 5/24, 6/23,10/21,15/19, etc.) also indicates dauphine twinning, although with a different Euler-3 angle relationship between 'parent' and 'twin' grains (Fig. 5d). Thus, the overall microstructure consists of not only alternating 'grains' within individual 'bands' but also

47

alternating 'bands', all of which have dauphine twin relationships to each other. The overall effect is a 60° 'corrugation' of the crystallographic microstructure in two effectively orthogonal directions to form long, narrow 'bands' and small, elongate grains (Fig. 5a). Although the mature mylonite LPO approximates to that of single crystal quartz, the rhomb poles display more than three orientation clusters (Fig. 2c, d), as would be expected for a single crystal orientation. Such a configuration is typical of dauphine twinning. This behaviour has been confirmed by a detailed SEM/EBSD linear traverse at 1 um step size (i.e. smaller than the grain sizes traversed) across part of the mylonite (see Fig. la for location). The pole figures for the traverse (Fig. 6a) are similar to those for the general mylonite petrofabric (Fig. 2c): an a-axis maximum subparallel to X; a great circle distribution of a-axes and m-poles subparallel to the XY foliation plane, normal to a strong c-axis maximum subparallel to Z; and superposition of the r and z poles. However, the dispersion patterns shown by the pole figures are stronger than for the general mylonite petrofabric and are compatible with slip on a prism system such as {m} (e.g. Lloyd & Freeman 1994), although this system was not recognized by Law et al. (1990) or Lloyd et al. (1992). The SEM/EBSD analyses show that dauphine twinning plays a role in both grain size reduction and mylonite development during shear zone evolution. Initial (i.e. relatively low shear strain) microstructural and petrofabric evolution is in part accommodated by dauphine twinning of original quartz grains, which helps to reduce grain size from several millimetres to a few 100 um. The mylonite LPO approximates a dauphine twinned quartz single crystal orientation onto which specific crystal slip systems have operated. However, a more complete appreciation of the role of dauphine twinning during shear zone microstructural and LPO evolution, and its relationship to crystal-slip systems that operate concurrently, can be achieved by assessing the misorientation relationships between adjacent (sub)grains, including the formation and orientation of (sub)grain boundaries.

Misorientation analysis The only EBSD data that can be used for 'correlated' (i.e. adjacent (sub)grains sharing a common boundary) misorientation analysis are from the manual analysis of the dauphine twin microstructure (Fig. 5a) and the linear traverse across the mylonite (Fig. la). The former

Fig. 5. SEM/EBSD analysis of localized dauphine twin microstructures (see Fig. la for location). Parts (c)-(f) were constructed using the program Channel5®. (a) Orientation contrast image, showing numbered points (1-18, along the band; 19-28 adjacent to the band on either side) from where EBSP were obtained, (b) Equal area, upper hemisphere pole figures illustrating basic crystallographic relationships for points 1-28 (dotted line indicates orientation of the analysed band), constructed via manual EBSD analysis and the program Pf2k (D. Mainprice; see also Mainprice 1990 and Mainprice & Humbert 1994). (c) Correlated (nearest neighbour) misorientation angle frequency histogram and expected random distribution curve for D3-32 symmetry class. Note distinctive peaks near 60° and at low angles (typically 5.5),

ACTIVE LOW-ANGLE NORMAL FAULTS in both compressional (Sibson & Xie 1998) and extensional environments (Collettini & Sibson 2001). Fluid overpressure in the study area might be the cause of the fault weakening mechanism, as also suggested by the presence of the extensional fractures within the ZF. Bearing in mind the difficulty in invoking fluid overpressures (Pf > 1). Note shortening normal to the deformation zone and significant zone boundary stretching in all models.

to restore the Goochland terrane to its paleogeographic position prior to dextral transpression. Consider George's Tavern (GT) in the Chopawamsic terrane and Goochland Courthouse (GCH) in the Goochland terrane (Figs 9 & 16). These locations are separated by the approximately 15 km wide SHSZ. Keeping George's Tavern fixed, Goochland Courthouse is retrodeformed to a predeformation position 80 to 300 km NE of its present location (Fig. 16). The greatest displacements are produced by the high-strain, pure shear dominated deformation (Model 4: Rs = 20, Wm = 0.4). Although the overall shear strain for Model 4 is less than for Model 2 (Rs = 20, Wm = 0.8), the pure shear dominated deformation produces the most zone normal thinning which results in significant offset when Goochland Courthouse is restored to its predeformation geometry (Figs 15 & 16). Displacement estimates are minimum values because strains calculated from boudinaged and folded dykes are minimum values and ultramylonite layers are likely to record strain ratios much greater than 20:1. In summary, the Goochland terrane, relative to the more western elements in the Virginia Piedmont, experienced significant southwestern translation during the Alleghanian orogeny.

Strain compatibility Non-plane strain, general shear deformations lead to compatibility problems with rigid wall

rocks because material is typically shortened normal to the zone and stretched parallel to zone boundaries. Bulk compatibility in general shear zones may be accomplished by: (1) deformable wall rocks; (2) sectional area changes along curved, non-parallel sided boundaries; (3) arrays of anastomosing high-strain zones; and (4) the development of discrete faults or dilational gaps (Simpson & De Paor 1993; Hudleston 1999; Bhattacharyya & Hudleston 2001). Away from Piedmont high-strain zones, rocks are penetratively deformed (albeit at lower strains) in complex patterns (Gates 1987; Gates & Glover 1989; Spears & Bailey 2002). In the Chopawamsic terrane, Goodman et al. (2001) recognized domains of constrictional and flattening strains with strike-parallel elongation lineations. The entire Chopawamsic terrane, between the BHSZ and the SHSZ, experienced elongation that is parallel and kinematically similar to deformation in the high-strain zones (Goodman et al 2001; Spears & Bailey 2002). With the exception of the northwestern boundary of the BHSZ in the Melrose pluton, the BHSZ and SHSZ have gradational boundaries with their wall rocks, such that there is a transition from high-strain to lower strain domains. Bulk compatibility is maintained between Piedmont high-strain zones and their wall rocks because deformation in both domains is kinematically related and occurs across a distinct strain gradient. Stretching faults occur in flowing rock bodies in which wall rocks lengthen or

TRANSPRESSION IN THE VIRGINIA PIEDMONT

261

Fig. 16. Present day map with the 'palaeogeographic'position of Goochland Courthouse prior to dextral transpressive deformation in the SHSZ. Goochland Courthouse retrodeformed using kinematic models in Fig. 15. Piedmont to the NW of the SHSZ is assumed to be fixed. Kinematic models require 80 to 300 km of SW displacement of the Goochland terrane relative to Piedmont terranes to the NW.

shorten while slip accumulates (Means 1989). Kinematically, Piedmont high-strain zones and their surrounding lower strain domains behaved as positive (material lengthening) stretching faults during deformation.

Tectonic significance of Piedmont high-strain zones Dextral transpressive high-strain zones in the Appalachian hinterland form a megaduplex structure (Costain et al 1987; Gates et al 1988; Hatcher 1989) active between about 320 and 280 Ma (Dallmeyer et al 1986; Horton et al 1987; Gates & Glover 1989; Wortman et al 1998). Strike-slip offset was important in the Piedmont, but the kinematics of deformation in the BHSZ and SHSZ reveals that material was also elongated parallel to the strike of the Appalachian orogen (NE-SW). Although collisional orogens are commonly viewed as thickening the crust and shortening the crust normal to the orogen, Piedmont high-strain zones demonstrate the significance of the three-dimensional (non-plane) strains and orogen-parallel material elongation. An unresolved controversy in southern Appalachian tectonics concerns the origin of the Goochland terrane. Mesoproterozoic rocks and A-type Neoproterozoic granitoids of the

Goochland terrane are similar to native Laurentian rocks in the Virginia Blue Ridge and have led many workers to conclude that the Goochland terrane is of Laurentian origin (Farrar 1984; Glover et al 1989; Aleinikoff et al 1996; Owens & Tucker 2000). Others suggest the Goochland terrane may be of peri-Gondwanan affinity and was accreted to Laurentia during the Appalachian orogen (Rankin et al 1989; Hibbard & Samson 1995). The Goochland terrane is currently separated from the Virginia Blue Ridge by the Mountain Run, Brookneal/Shores, and Spotsylvania high-strain zones (Fig. 2). All of these structures record late Palaeozoic dextral transpression (Gates et al 1986; Bobyarchick 1999; this work). If displacement on the Mountain Run and Brookneal/Shores high-strain zones are of similar magnitude to that of the SHSZ (over 100 km), the Goochland terrane was located at least 300 km NE of the Virginia Blue Ridge prior to the Alleghanian orogen. Tectonic models directly linking the Goochland terrane to the Virginia Blue Ridge are untenable because of significant dextral translation across Alleghanian high-strain zones. Quantitative understanding of the kinematics does not resolve whether the Goochland is a native Laurentian or exotic terrane, but does place meaningful limits on its pre-Alleghanian position in the Appalachian orogen. If the

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Goochland terrane is Laurentian, it originated somewhere between the Pennsylvania reentrant and the New York promontory, not outboard of the Virginia Blue Ridge. Conclusions Although naturally deformed rocks experience complex strain paths that may not be fully recorded in rock structures, strain and vorticity analyses provide reasonable limits on the nature of deformation in high-strain zones. Analysis of two Palaeozoic high-strain zones in the Virginia Piedmont reveals that they experienced nonplane strain general shear (Wm c) has a rvalue of +1; deformation (e.g. Ghosh & Ramberg 1976; a sphere has a lvalue equal to zero. If a > b > c Jezek et al 1994,1996). The effect of orientation then T lies between 0 and +1; alternatively if a < (Fig. 4), aspect ratio (Fig. 5), and vorticity (Fig. b & VAN DEN DRIESSCHE, J. 1993. La zone de cisaillement de Quiberon: temoin d'extension de la chaine varisque en Bretagne meridionale au Carbonifere. Comptes Rendus de VAcademie des Sciences, II, 1123-1129. GAPAIS, D. & LE CORRE, C. 1980. Is the Hercynian belt of Brittany a major shear zone? Nature, 288, 574-576. GODARD, G. 2001. Les Essarts eclogite-bearing Complex (Vendee). Geology of France 'Special Vendee', 1-2,19-51. GORE, B. & LE CORRE, C. 1987. Cinematique hercynienne du cisaillement nord-armoricain a la bordure du granite syntectonique de Saint Renan-Kersaint (Finistere). Bulletin de la Societe Geologique de France, 3, 811-819. GUILLOCHEAU, F. & ROLET, J. 1982. La Sedimentation Paleozoi'que Ouest-Armoricaine. Bulletin de la Societe Geologique et Mineralogique de Bretagne, 14, 45-62. HANMER, S. & VIGNERESSE, J.L. 1980. Mise en place de diapirs syntectoniques dans la chaine hercynienne: exemple des massifs leucogranitiques de Locronan et de Pontivy (Bretagne Centrale). Bulletin de la Societe Geologique de France, 22, 193-202. HANMER, S., LE CORRE, C. & BERTHE, D. 1982. The role of Hercynian granites in the deformation and metamorphism of Brioverian and Palaeozoic rocks of Central Brittany. Journal of the Geological Society of London, 139, 85-93. HERROUIN, Y., DADET, P., GUIGUES, I, LAVILLE, P. & TALBO, H. 1989. Geological map (1/50 000) - Bain de Bretagne (388). BRGM, Orleans. HERROUIN, Y. & RABU, D. 1990. Geological map (1/50 000) - Chateaubriant (389). BRGM, Orleans. HIRBEC, Y. 1979. Le complexe basique de Belle-Isle-enTerre (Cotes du Nord). Sa place dans revolution geodynamique du nord du Massif Armoricain. Troisieme cycle universitaire thesis, Rennes 1, Rennes. JANJOU, D., LARDEUX, H., CHANTRAINE, J., CALLIER, L. & ETIENNE, H. 1998. Geological map (1/50 000) Segre (422). BRGM, Orleans. JEGOUZO, P. 1980. The South Armorican Shear Zone. Journal of Structural Geology, 2, 39-47. JEGOUZO, P., PEUCAT, J.-J. & AUDREN, C. 1986. Caracterisation et signification geodynamique des orthogneiss calco-alcalins d'age ordovicien de Bretagne meridionale. Bulletin de la Societe Geologique de France, 2, 839-848. JEGOUZO, P. & ROSELLO, E.A. 1988. La branche nord du cisaillement sud-armoricain (France): un essai devaluation du deplacement par 1'analyse des

mylonites. Comptes Rendus de VAcademie des Sciences de Paris, 2,1825-1831. LE CORRE, C. 1977. Le Brioverien de Bretagne centrale: essai de synthese lithologique et structurale. Bulletin du B.R.G.M. Section I, 3, 219-254. LE CORRE, C. 1978. Approche quantitative des processus synschisteux. L'exemple du segment Hercynien de Bretagne Centrale. Etat Thesis, Rennes 1, Rennes. LE CORRE, C, AUVRAY, B., BALLEVRE, M. & ROBARDET, M. 1991. Le Massif Armoricain. Scientific Geological Bulletin, 44, 31-103. LE HEBEL, F, VIDAL, O., KIENAST, J.R. & GAPAIS, D. 2002. Les 'Porphyroides' de Bretagne meridionale: une unite de HP-BT dans la chaine hercynienne. Comptes Rendus Geoscience, 334, 205-211. LE METOUR, J. & AUDREN, C. 1977. Relations structurales entre 1'orthogneiss ordovicien de Roguedas et son encaissant migmatitique. Consequences sur 1'age des evenements tectonometamorphiques en Bretagne meridionale. Bulletin de la Societe Geologique et Mineralogique de Bretagne, 9,113-123. LE THEOFF, B. 1977. Marqueurs ellipso'idaux et deformation finie. Troisieme Cycle Universitaire Thesis, Rennes 1, Rennes. LEDRU, P., MAROT, A. & HERROUIN, Y. 1986. Le synclinorium de Saint-Georges-sur-Loire: une unite ligerienne charriee sur le domaine centre armoricain. Decouverte de metabasite a glaucophane sur la bordure sud de cette unite. Comptes Rendus de VAcademie des Sciences, 303-11, 963-968. LOPEZ-MUNOZ, M. 1981. Analyse structurale de la partie centrale du synclinorium de Saint-Julien de Vouvantes et de 1'axe Lanvaux-les-Ponts-de-Ce (Massif Armoricain). Bulletin de la Societe Geologique et Mineralogique de Bretagne - serie C, 13,117-123. MATHERON, G. 1955. Application des methodes statistiques a 1'evaluation des gisements. Annales des Mines, 144, 50-75. MATHERON, G. 1962. Traite de Geostatistique Appliquee. Technip, Paris. PARIS, F. & DADET, P. 1988. Geological map - Combourg (1/50 000). BRGM, Orleans. PERCEVAULT, M.-N. & COBBOLD, P.R. 1982. Mathematical removal of regional ductile strains in Central Brittany: evidence for wrench tectonics. Tectonophysics, 82, 317-328. PEUCAT, J.-J., CHARLOT, R., MIDFAL, A., CHANTRAINE, J. & AUTRAN, A. 1979. Definition geochronologique de la phase bretonne en Bretagne Centrale. Etude Rb/Sr de granites du domaine centre armoricain. Bulletin du BRGM, 1, 349-356. PIVETTE, B. 1978. Le synclinorium de Saint-Georges sur Loire, Massif Armoricain. Sa place dans revolution geodynamique de la Bretagne meridionale au Paleozoique. Troisieme cycle universitaire thesis, Rennes 1, Rennes. PLAINE, J. 1976. La bordure sud du synclinorium paleozo'ique de Laval Troisieme cycle universitaire thesis, Rennes 1, Rennes. PLAINE, I, HALLEGOUET, B. & QUETE, Y. 1984.

STRAIN REMOVAL IN CENTRAL BRITTANY Geological map (1/50000) - Questembert (418). BRGM, Orleans. PLUSQUELLEC, Y., ROLET, J. & DARBOUX, J.R. 1999. Geological map - Chdteaulin (1/50,000). BRGM, Orleans. QUETE, Y, PLAINE, J. & HALLEGOUET, B. 1981. Geological map (1/50 000) - Malestroit (386). BRGM, Orleans. RAMSAY, J.G. 1967. Folding and fracturing rocks. McGraw-Hill. RAMSAY, J.G. & GRAHAM, R.H. 1970. Strain variations in shear belts. Canadian Journal of Earth Sciences, 7, 786-813. REGNAULT, S. 1981. Stratigraphie et structure du Paleozoi'que dans le Menez-Belair occidental (Synclinorium median armoricain). Bulletin de la Societe Geologique et Mineralogique de Bretagne -C, 13,1-111.

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TRAUTMANN, F. 1988. Geological map (1/50000) Nozay (420). BRGM, Orleans. TRAUTMANN, E, BECQ-GIRAUDON, J.F & CARN, A. 1994. Geological map (1/50000) - Janze (353). BRGM, Orleans. UPTON, G.J.G. & FINGLETON, B. 1989. Spatial Data Analysis by Example, vol. 2. John Wiley & Sons, New York. VIDAL,P 1972. L'axe granitique Moelan-Lanvaux (sud du Massif Armoricain): mise en evidence par la methode Rb-Sr de trois episodes de plutonisme pendant le Paleozoi'que Inferieur. Bulletin de la Societe Geologique et Mineralogique de Bretagne, 4, 75-89. VIDAL, P. 1976. L'evolution polyorogenique du Massif Armoricain. Apport de la geochronologie et de la geochimie isotopique du strontium. Troisieme cycle universitaire thesis, Rennes 1, Rennes. ROBARDET, M., BONJOUR, J.L., PARIS, E, MAORZADEC, VIDAL, P. 1980. L'evolution polyorogenique du Massif P. & RACHEBOEUF, PR. 1994. Ordovician, SilArmoricain. Apport de la geochronologie et de la urian, and Devonian of the Medio-North-Armorgeochimie isotopique du strontium. Memoires de ican Domain. In: KEPPIE, J.D. (ed.) Pre-Mesozoic la Societe Geologique et Mineralogique de Bregeology in France and related areas. Springertagne, 21,1-162. Verlag, Berlin. VIGNERESSE, J.L. 1978. Gravimetrie et granites armoriROLET, J. & THONON, P. 1979. Mise en evidence de trois cains. Structure et mise en place des granites Hercomplexes volcano detritiques d'age Devonien cyniens. Troisieme cycle universitaire thesis, inferieur a moyen, Strunien et Viseen inferieur Rennesl, Rennes. sur la bordure nord du bassin de Chateaulin VIGNERESSE, J.L. & BRUN, J.P. 1983. Les leucogranites (feuille Huelgoat 1/50000, Finistere). Impliarmoricains marqueurs de la deformation cations paleogeographiques et tectoniques. regionale: apport de la gravimetric. Bulletin de la Bulletin du BRGM, 1, 303-315. Societe Geologique de France, 25, 357-366. ROLIN, P. & COLCHEN, M. 2001. Les cisaillements her- WACKERNAGEL, H. 1998. Multivariate geostatistics. Springer, Berlin. cyniens de la Vendee au Limousin. Geologic de la France, 1-2,15-44. WATTS, M.S. & WILLIAMS, G.D. 1979. Faults rocks as ROMAN-BERDIEL, T, GAPAIS, D. & BRUN, J.-P. 1997. indicators of progressive shear deformation in the Granite intrusion along strike-slip zones in Gingamp region, Brittany. Journal of Structural experiment and nature. American Journal of Geology, 1, 323-332. Sciences, 297, 651-678. WEBER, C. 1967. Le prolongement des granites de SAGON, J.P 1976. Contribution a I'etude geologique de Lanvaux d'apres la gravimetric et Faeromagla partie orientale du bassin de Chateaulin. Stratinetisme. Memoires du BRGM, 52, 83-90. graphie, volcanisme, metamorphisme et tectonique. Etat thesis, Paris 6, Paris.


Strain and deformation history in a syntectonic pluton. The case of the Roses granodiorite (Cap de Creus, Eastern Pyrenees) J. CARRERAS, E. DRUGUET, A. GRIERA & J. SOLDEVILA Departament de Geologia, Universitat Autdnoma de Barcelona, 08193 Bellaterra, Barcelona, Spain (e-mail: [email protected]) Abstract: The Roses granodiorite is a Variscan stock with well developed syn- and postmagmatic deformation structures that crops out in the Pyrenean Axial Zone. Analysis of structures reveals a continuous deformation history during and after magma cooling. The deformation history is divided on the basis of mechanical behaviour into two stages: an early one with the development of magmatic structures and a late stage with the development of mylonitic fabrics along shear zones. Both stages are separated in time by the emplacement of aplite-pegmatite dykes. Time of dyke emplacement is thought to coincide with a sudden change in rheology of the granodiorite. The abundance of quartz dioritic enclaves permits the use of shape analysis to characterize the magmatic fabric as a homogeneous deformation. Later solid-state deformation led to the development of an inhomogeneous deformation pattern with different sizes of anastomosing shear zones wrapping around lozenge-shaped domains. The displacement/width ratio measured in shear zones ranges between one and two orders of magnitude. The Roses granodiorite is thought to be a synkinematically emplaced stock which records a continuous deformational history with two distinct deformation stages, both recording bulk finite strains of similar order of magnitude but with a marked difference in finite strain distribution.

Granitoid batholiths and stocks are abundant in the Variscan basement of the Pyrenees (Fig. 1). Although these were initially referred to in the literature as Variscan late-tectonic intrusions (e.g. Maladeta Massif; Zwart 1979), more recent studies have revealed the syntectonic nature of many of these plutons (e.g. Bassies Granite, Gleizes et al 1991). Such studies have been carried out both on the internal fabrics, by the use of the AMS method (Leblanc et al 1996; Gleizes et al. 19980), and on the structures present in the country rock aureoles (Evans et al. 1998). Thus, most Variscan plutons in the Pyrenees are now well characterized by means of deformational features and relative time of intrusion, being generally accepted as emplaced during a main Variscan deformational event. Furthermore, some of these syntectonic plutons are also affected by shear belts developing mylonite bands (Fig. 1). The age and geotectonic significance of the mylonite belts is still under debate (Guitard 1970; Carreras etal 1980; Lamouroux et al. 1980; Saillant 1982; Delaperriere et al 1994). The Roses and Rodes massifs are two synkinematically emplaced stocks of mainly granodioritic composition, located on the Cap de Creus Peninsula, which forms the easternmost outcrop of the Palaeozoic basement in the Pyrenean Axial Zone (Fig. 1). The Roses stock, elongated in a NW-SE direction, was emplaced

into low grade Cambro-Ordovician metasediments, developing a narrow contact aureole of spotted phyllites and hornfelses. The enclosing metasediments exhibit a polyphase history with two main deformation events. The first is responsible for a layer-subparallel penetrative cleavage referred to as the regional foliation. The second is represented by an inhomogeneously distributed crenulation cleavage which postdates the contact metamorphism. This deformation history is revealed by contact metamorphic porphyroblasts that grew over the main foliation, but show a crenulation cleavage related to the late folding phase wrapping around them. The emplacement of the Roses granodiorite started before the late folding phase. Folds related to the crenulation in the enclosing metasediments were contemporaneous with shear zone development in the granodiorite (Carreras & Losantos 1982). These two distinct types of structures, folds in the metasediments and shear zones in the granodiorite, possibly reflect how the two lithological units with different rheological properties responded to deformation. The orientation of all Variscan structures in the Roses granodiorite is affected by a local Alpine overturning. This overturning occurs all along the southern border of the Pyrenean Axial Zone (Vergely 1970; Munoz et al 1986) and


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J. CARRERAS ETAL.

Fig. 1. Sketch of the main lithological units and structures in the Variscan of the eastern Pyrenees and location of the Roses granodiorite.

causes originally NE-dipping dextral shear zones to appear as sinistral ones (Carreras 2001).

Progressive development of structures in the Roses granodiorite Like some other Variscan granitoid massifs of the Pyrenean basement (e.g. Querigut Massif; Marre 1973 and Bassies granodiorite; Gleizes et al. 1991), the Roses granodiorite shows an early homogeneous magmatic fabric and a network of low-temperature shear zones (Fig. 2) that gave rise to inhomogeneous mylonitization at low greenschist facies, presumably of Variscan age (Carreras & Losantos 1982). This was followed by late cataclasis developed in narrow bands of millimetre thickness (Simpson et al. 1982). A range of magmatic-state/pre-full crystallization to solid-state/crystal plastic strain fabrics developed throughout the cooling history from hightemperature to the low-temperature regimes, in a manner similar to that proposed by Gapais (1989) in a general model. The existence of this continuous gradation of structures and the difficulty in establishing a clear distinction between the early synmagmatic structures and the fabrics related to the late

shear zones will be discussed below. This will be followed by an analysis of the magmatic fabrics and the structures related to late solid-state shearing. For the sake of simplicity and objectivity, these analyses will be presented separately, using the presence of a swarm of aplite-pegmatite dykes to distinguish between the early structures predating dyke emplacement and those affecting the dykes (Fig. 3).

Magmatic fabric and enclaves The oldest tectonic structure in the granodiorite is a pre-full crystallization fabric (Hutton 1988) or magmatic fabric (Paterson et al. 1989), defined by a preferred orientation of subhedral feldspar, sometimes with tiling between pairs of crystals, along with a weaker alignment of mafic minerals (biotite and amphibole). Additional evidence of magmatic flow is the preferred orientation of elongated enclaves (Fig. 4a) and the presence of schlieren layering in the granodiorite. These are all indicative of synmagmatic deformation. Synmagmatic foliations trend E-W to NW-SE (Figs 3 & 5) in a vaguely curved disposition. The magmatic fabric postdates the regional foliation in the enclosing sediments as is

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Fig. 2. Geological setting of the Roses granodiorite and the studied area, (a) Structural map of the studied area showing traces of the main shear zones. Location of the area is shown in Fig. 1. (b) Stereoplot showing the orientation of the poles to the mylonite foliation and the associated stretching lineation. (c) Section across the Roses granodiorite.

evidenced from the presence of foliated metasedimentary xenoliths in the granodiorite. The Roses granodiorite is characterized by an abundance of enclaves (Figs 4a, b, c, & 5), most of them microquartz dioritic, with abundant mafic minerals (biotite and amphibole). There is also a small proportion of metasedimentary

xenoliths. Enclave distribution is inhomogeneous, with the presence of some dismembered synplutonic microquartz dioritic sheets. The enclaves are predominantly flattened and show a marked preferred orientation subparallel to the magmatic foliation (Fig. 4a, c). This foliation exhibits little or no deflection around

Fig. 3. Schematic qualitative model of the structural history of Roses granodiorite. The progressive development of structures from high to low temperature can be divided in two major stages with regard to the time of dyke emplacement. Stereoplots of orientation of different structures are also shown.


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Fig. 4. Mesoscopic scale structures in the Roses granodiorite. (a) Preferred orientation of elongate microquartz dioritic enclaves of a former synmagmatic dyke, (b) Straight aplite vein cutting across the magmatic fabric marked by orientation of enclaves, (c) Syntectonically emplaced leucocratic dyke forming open folds with axial planes parallel to the magmatic fabric, (d) Microgranite with igneous texture emplaced along a curved shear zone, (e) Leucocratic dyke of aplite in the granodiorite cut by sinistral shear zones (f) Brittle-ductile transition conjugate fractures cutting across mylonites.

Fig. 5. Pre-dyke finite strains obtained from two-dimensional analysis of enclave shapes in sections close to the XZ plane, in a domain without significant late shearing. For each locality, the mean strain axial ratio (RS) is shown with an ellipse showing mean orientation and shape of the strain ellipse. Location is shown in Fig. 2.


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Fig. 6. Part of an extremely elongated enclave of quartz diorite with an axial ratio of about 150:1 from the inner part of a shear zone.

elongated enclaves, from which it is envisaged that the enclaves were slightly more competent during development of the magmatic fabric (Gay 1968; Paterson & Miller 1998). Later solid-state deformation would probably have caused a further decrease in viscosity ratio, in agreement with the work by Tobisch & Williams (1998).

Leucocratic dykes In most domains the magmatic fabric is cut by straight granophyric aplite-pegmatite veins and dykes (Fig. 4b) that vary in size from a few millimetres to about one metre in width and tens of metres in length. They form a swarm of dykes of predominant NNE-SSW original orientation (Fig. 3), cross-cutting the magmatic fabric and the elongated enclaves at a high angle (70-90°). Thus, dykes are presumed to occur in tension fractures related to the same stress/strain field that produced the magmatic foliation. Furthermore, the dykes are slightly folded locally, with fold axial planes parallel to the earlier magmatic fabric (Fig. 4c), suggesting that high-temperature solid-state deformation occurred with a similar orientation in the granodiorite during and after dyke emplacement. Dykes do not show significant refraction when passing through host/enclave interfaces (Fig. 4b). The emplacement of dykes into a network of brittle fractures affecting the granodiorite is presumed to have been coeval with a sudden change in rheology of the granodiorite stock. The granodiorite, at a submagmatic stage with an overpressured residual leucogranitic melt, would have experienced a 'melt-enhanced embrittle-

ment' (Hollister & Crawford 1986; Davidson et al 1994; Brown & Solar 1998; Handy et al 2001).

Shear zones and associated mylonites Post-dyke deformation is mainly related to the development of shear zones, affecting both granodiorites and leucocratic dykes (Figs 4e & 6). In their present orientations shear zones form a NW-SE anastomosing network with predominant steep dips (Figs 2 & 3) and dominantly sinistral strike-slip components. Associated mylonite bands range in thickness from a few millimetres to more than 500 m. In contrast with magmatic structures, which are widely distributed in the entire stock, the shear zones represent a highly inhomogeneous deformation leaving different-sized elongate domains nearly untouched by mylonitization (Simpson et al. 1982). Movement direction is shown by the disposition of the stretching lineation (Figs 2b & 3). The sense of movement can be depicted easily from the marginal obliquity of the mylonitic foliation and the offset of aplitic dykes in appropiate sections (Fig. 4e). In addition there are abundant shear-band structures with dispositions always coherent with the depicted sense of shear. The anastomosing and fan-like pattern, with dominant sinistral shear zones and less abundant dextral ones, could represent a conjugated system. However, shear directions do not form two maxima but a single one coinciding with the intersection direction of differently orientated planes (Figs 2b & 3). Thus, in a stereoplot, poles of foliation planes define a great circle with its

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pole coinciding with the stretching lineation. This geometrical relation appears to be a common feature of shear belts and is analogous to that described by Ramsay and Allison (1979) in the Maggia Nappe. Most mylonites developed under greenschist facies metamorphic conditions, assisted by strong quartz recrystallization and new growth of chlorite, albite, white mica and epidote. Some broad shear zones contain bands, ranging up to a few metres in width, where the mylonites are completely depleted of quartz, and albitechlorite mylonites form. This mineralogical and chemical transformation of the mylonites may have occurred along zones where fluid was channelled. These zones are probably related to the intrusion of quartz veins and dykes contemporaneous to mylonitization. Although nearly all shear zones formed in the solid-state and under greenschist facies conditions, a peculiar type has been observed. This consists of a complex network of centimetrethick shear zones with fine-grained granitic isotropic material injected along them (Fig. 4d). This particular type of shear zone is considered to represent the earliest stages of localized deformation before the complete crystallization of the granodiorite, and is presented as another argument for the transition from magmatic to solid-state deformation during the cooling of the stock.

fabrics and the superimposed effects of later mylonitization developed under greenschist facies conditions. In addition, rotation and thinning of aplite-pegmatite dykes by shearing and related offset enables shear strain determinations across shear zones. Furthermore, the outcrop conditions along the coastal fringe located in Fig. 2a, with abundant sections close to plane view, enable the establishment of strain profiles across the described structures and the determination of the kinematic pattern associated with shearing.

Deformation predating dykes

Strain analysis has been performed using the Rf/Q technique (Lisle 1985) for enclave populations in different locations. By measuring shape ratio and orientation of the enclaves one can infer information about the total strain and put some constraints on the deformation history. Previously, mafic enclaves have been used for quantitative estimation of finite strain (Ramsay & Rubber 1983; Mutton 1988; Williams and Tobisch 1994; Tobisch & Williams 1998; Wenk 1998). However, the reliability of enclaves as strain markers has been questioned because of the existence of some variables that cannot be related directly to deformation processes (see discussions in Paterson & Vernon 1995 and Tobisch & Williams 1998). Among these variables the most important are: enclave shapes and orientation may reflect other processes Late brittle fractures besides deformation during emplacement, such At the terminations of shear zones, but also as ascent and chamber boundary processes; and cutting across ductile shear zones (Fig. 4f), very the possible contrast in rheology between narrow fractures with associated cataclasites are enclaves and host rocks. present. Although brittle fractures on shear The prevalent subhorizontal sections were zone tips occur in association with shear zone not adequate to perform an exhaustive 3D propagation (Simpson 1983), the cross-cutting analysis. Instead, a general qualitative examinones represent the latest structures formed in ation of enclave population on the entire the granodiorite. These fractures commonly enclave-rich area was made first, from which it form conjugate sets but, in contrast with preced- was established that the XZ section of most ing ductile shear zones, conjugate brittle frac- enclaves is roughly sub-horizontal. After that, a tures always occur at an acute angle to the three-dimensional analysis from three different compressional field, with principal compression sections at one location was performed to estiaxis in an orientation close to north-south. Note mate the orientation and shape of the strain that the final orientation of the compressional ellipsoid. This gave a flattening strain ellipsoid direction is similar and coherent with the orien- with Rxz approximately 3, Ryz about 2.5, a tation required to develop the initial magmatic gently east-dipping XZ plane of finite strain and fabric, direction of dyke emplacement and the east-plunging maximum elongation (X). Then, based on the qualitative assessment and the later shear zones. result from three-dimensional analysis, an extensive two-dimensional strain analysis was Structure and strain profiles carried out on the prevalent sub-horizontal secThe abundance of micro-quartzdioritic enclaves tions, close to the XZ plane. Each location corpermits the use of shape analysis to characterize responds to a homogeneous domain covering a the magmatic to high-temperature solid-state surface ranging from a few square metres up to


about 100 m2, where undeformed dykes provide evidence for the absence of post-dyke strains. In each locality, between 40 and 70 enclaves were measured. The metasedimentary xenoliths and the dismembered microquartz dioritic sheets have been excluded from the shape analysis. Enclaves with unusual shapes and orientations with regard to the main distribution have also been excluded, as they can yield erroneous results. Rf/(j> analyses show a pre-dyke finite strain characterized by axial ratios ranging between 1.1 and 4. Figures 5 and 7 show shapes and orientations of pre-dyke finite strains in two different domains. These finite strains represent the sum of a large magmatic strain and a minor high temperature solid-state deformation. The results probably underestimate the total pre-dyke strain, due to: (1) the presumed viscosity contrast (albeit low) between enclaves and granodiorite; (2) the analysed section, which is not perfectly parallel to the XZ plane of the finite strain ellipsoid; and (3) the earliest deformation in the magmatic history may not have been, or only slightly recorded by the enclaves (Davidson et al. 1994). Not withstanding these problems, the strain measurements using enclave shape give an acceptable estimate of finite strain gradients undergone by the granodioritic stock. Bulk results obtained from different localities indicate that there is a slight inhomogeneity in strain and that this is of low intensity but widespread across the entire stock. Extrapolating the observed mean strain values to the entire pluton, the pre-dyke synmagmatic deformation related to emplacement of the stock represents a bulk horizontal shortening of about 33%.

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ation patterns following a profile which is close to the one expected for simple shear. The shear zone profile analyses (Fig. 8) reveal a high inhomogeneity of strain at different scales. A mean shear strain value from each profile has been calculated by means of the total displacement/total width relationship. Furthermore, an analysis of the displacement/width ratio of shear zones is constant on shear zones of different size and ranges between one and two orders of magnitude (Fig. 9). Although the point distribution in this graph has a similar slope to the plot presented in Mitra (1979), significant differences exist concerning the intersecting point along the displacement axis. In the case of the Roses granodiorite, displacement in each shear zone is generally at least ten times its width, and therefore about one order of magnitude greater than in the relationship shown by Mitra (1979). Averaging the total displacement/total width relationships obtained from different profiles, and evaluating the total width of shear zones versus the total width of unsheared rocks in the area, a bulk shear strain of 1.4 was estimated. This corresponds to a finite strain axial ratio of 3.3. Extrapolating the shear strain values to the entire stock, by means of the movement along the described network of shear zones, we infer a post-dyke bulk horizontal shortening of about 45%. Discussion and conclusions

The Roses case study shows that enclaves and shear zone profiles are powerful tools for the comparison of deformation during two stages of the stock history. Although the magmatic fabric is less conspicuous than the mylonitic one, it Deformation postdating dykes appears that, considering average values, the In broad mylonitic bands, strain analysis was magnitude of horizontal shortening accommoperformed using enclave shape and orientation dated during the early pre-dyke stage is lower (Fig. 7). Enclaves reflect the variable degree of but of similar order of magnitude to the shortdeformation due to the high inhomogeneity of ening accommodated during solid-state mylonistrain distribution. In sections closely parallel to tization. If we superpose the post-dyke the finite XZ plane, RS values greater than 10 shortening (45%) on the pre-dyke shortening are common, with some enclaves reaching (33%), the granodiorite recorded a bulk horivalues up to RS = 150 (Fig. 6). zontal shortening of about 60%. This value is In thin shear zones, the best and most reliable slightly smaller than expected if the two superstrain profiles (Fig. 8) were obtained by applying posed deformations were coaxial. the Ramsay & Graham (1970) method for shear These two deformation events developed strain determination. This technique is based on under different temperature conditions and are the combined use of sigmoidal pattern of clearly separated in time by the emplacement of mylonitic foliation and change in orientations aplite-pegmatite dykes. However, the presence and offsets of dykes, assuming a simple shear of high-temperature fabrics related to open model for the shear zones. This assumption is folding of synkinematic dykes and the presence supported by the fact that the shear zones are of shear zones with igneous material injected localized along narrow, discrete bands, with foli- along them, indicate the continuity of

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Fig. 7. Structural map and strain analysis of pre-dyke structures and post-dyke shear zones along a coastal section east of the Roses Lighthouse. In both cases strain analysis was performed using the Rf/Q technique for enclave populations. For each locality, the mean strain axial ratio (RS) is shown with an ellipse showing mean orientation and shape of the strain ellipse. Location is shown in Fig. 2.

deformation from the magmatic to the mylonitic stages, with the orientation of the regional shortening direction remaining fairly constant. Furthermore, the preferred disposition of dykes, at high angle to the magmatic fabric, suggests that dyke emplacement occurred along tension fractures compatible with the strain field active during preceding and subsequent events. In this way, there is no need for an interkine-

matic regime or a shift in tectonic regime from compressional to extensional to explain dyke emplacement. The structures observed in the Roses granodioritic stock developed during syntectonic cooling and reveal a continuity of deformation from high- to low-temperature conditions. A change of rheology from high- to low-temperature regime is capable of explaining the


317

Fig. 8. Shear strain analysis across three sections in areas affected by inhomogeneous post-dyke shearing. Mean shear strain values have been calculated using total width versus total displacement relationships. All sections are plane-view.

observed structures and strain pattern. At high temperature, strain is distributed more homogeneously, whereas the progressive temperature drop induces a inhomogeneous strain pattern due to high strain localization along shear zones. Two critical points in the rheological history of the stock can be identified: (1) The first one took place at high temperature when the dykes intruded. At this stage, the granodioritic melt was close to completion of crystallization and probably below the critical fraction for magmatic flow. Thus, this first critical point would not be exactly contemporaneous with the 'rheologically critical melt percentage', theoretically and experimentally inferred as a sudden change of rocks strength caused by the crossing of a critical volume fraction of melt (Arzi 1978; Van der Molen & Paterson 1979; Dell' Angelo & Tunis 1988; Rutter & Neumann 1995; Vigneresse and Tikoff 1999). Most probably, high fluid pressures, induced by the presence of water saturated melts, favoured brittle tension fractures filled with

the residual melts at this stage ('melt enhanced embrittlement'). Furthermore, a strain rate increase could also facilitate local brittle failure. (2) The second critical point corresponds to the low temperature ductile-brittle transition in granitic rocks when shearing is replaced by discrete faults. Most of the overall deformation was accommodated during the ductile stages (magmatic fabric, high-temperature solid-state deformation and shear zones), whereas deformation accommodated during brittle stages was negligible relative to the total deformation. Although no conclusive evidence for the strain regime responsible for synmagmatic deformation has been found in Roses, a dextral transpressional regime is inferred from correlation with the nearby northern Cap de Creus tectonometamorphic belt, where the kinematics are well established and documented (Carreras 2001; Druguet 2001). Moreover, this is also in accord with the work by Gleizes et al (19985) that indicates a dextral transpression as a main tectonic

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Fig. 9. Diagram of displacement against width measured in different sized shear zones. The displacement-width relationship line by Mitra (1979) has been drawn for comparison.

setting for the emplacement of granitoid batholiths in the Variscan of the Pyrenees. This work was financed by the BTE2001-2616 project (M.C.Y.T). The manuscript has benefited greatly from thoughtful reviews by B. Miller and S. Johnson. We thank P. D. Bons for his kind comments and help with English. We also thank I. Tribe and C. Simpson for careful review and constructive comments on a previous version of the manuscript. The stereographic analysis was done with Stereonet, a program by R. W. Allmendinger.

References ARZI, A. 1978. Critical phenomena in the rheology of partially melted rocks. Tectonophysics, 44, 173-184. BROWN, M. & SOLAR, G.S. 1998. Granite ascent and emplacement during contractional deformation convergent orogens. Journal of Structural Geology, 20,1365-1393. CARRERAS, J. 2001. Zooming on Northern Cap de Creus shear zones. Journal of Structural Geology, 23,1457-1486 CARRERAS,!. & LOSANTOS,M. 1982. Geological setting of the Roses granodiorite, (E-Pyrenees, Spain). Ada Geologica Hispanica, 17, 211-217. CARRERAS, J., JULIVERT, M. & SANTANACH, P. 1980. Hercynian Mylonite Belts in the Eastern Pyrenees: an example of shear zones associated with late folding. Journal of Structural Geology, 2, 5-9. DAVIDSON, C, SCHMID, S.M. & HOLLISTER, L.S. 1994. Role of melt during deformation in the deep crust. Terra Nova, 1,133-142. DELAPERRIERE, E., DE SAINT BLANQUAT, M., BRUNEL,

M. & LANCELOT, J. 1994. Geochronoloige U-Pb sur zircons et monazites dans le massif du Saint Barthelemy (Pyrenees, France): discussion des ages des evenements varisques et pre-varisques. Bulletin de la Societe Geologique de France, 165, 101-112. DELL'ANGELO, L.N. & TULLIS, J. 1988. Experimental deformation of partially melted granitic aggregates. Journal of Metamorphic Geology, 6, 495-515. DRUGUET, E. 2001. Development of high thermal gradients by coeval transpression and magmatism during the Variscan orogeny: insights from the Cap de Creus (Eastern Pyrenees). Tectonophysics, 332, 275-293. EVANS, N.G., GLEIZES, G, LEBLANC, D. & BOUCHEZ, J.L. 1998. Syntectonic emplacement of the Maladeta granite (Pyrenees) deduced from relationships between Hercynian deformation and contact metamorphism. Journal of the Geological Society, London, 155, 209-216. GAPAIS, D. 1989. Shear structures within deformed granites: Mechanical and thermal indicators. Geology, 17,1144-1147. GAY, N.C. 1968. Pure shear and simple shear deformation of inhomogeneous viscous fluid. 2. The determination of the total finite strain in rocks from objects such as deformed pebbles. Tectonophysics, 5, 295-302. GLEIZES, G, LEBLANC, D. & BOUCHEZ, J.L. 1991. Le pluton granitique de B assies (Pyrenees ariegoises): zonation, structure et mise en place. Comptes Rendus de VAcademie des Sciences (Paris), 312, 755-762. GLEIZES, G, LEBLANC, D., SANTANA, V., OLIVIER, P. & BOUCHEZ, J.L. 19980. Sigmoidal structures featuring dextral shear during emplacement of the Hercynian granite complex of Cauterets-Panticosa,

STRAIN ANALYSIS IN A SYNTECTONIC PLUTON Pyrenees. Journal of Structural Geology, 20, 1229-1245. GLEIZES, G., LEBLANC, D. & BOUCHEZ, J.L. 19986. The main phase of the Hercynian Pyrenees is a dextral transpression. In: HOLDSWORTH, R.E, STRACHAN, R.A & DEWEY, IF. (eds) Continental transpressional and transtensional tectonics. Geological Society, London, Special Publications, 135, 267-273. GUITARD, G. 1970. Le metamorphisme hercynien mesozonal et les gneiss oeilles du massif du Canigou, (Pyrenees Orientales). Memoir-es du Bureau de Recherches Geologiques et Minieres, 63, 353 pp. HANDY, M.R., MULCH, A., ROSENAU,M. & ROSENBERG, C.L. 2001. The role of fault zones and melts as agents of weakening, hardening and differentiation of the continental crust: a synthesis. In: HOLDSWORTH, R.E., STRACHAN, R.A., MAGLOUGHLIN, J.F. & KNIPE, R.J. (eds) The nature and tectonic significance of fault zone weakening. Geological Society, London, Special Publications, 186, 305-332. HOLLISTER, L.S. & CRAWFORD, M.L. 1986. Meltenhanced deformation: a major tectonic process. Geology, 14, 558-561. HUTTON, D.H.W. 1988. Granite emplacement mechanisms and tectonic controls: inferences from deformation studies. Transactions of the Royal Society of Edinburgh, 79, 245-255. LAMOUROUX, C, SOULA, J.C., DERAMOND, J. & DEBAT, P. 1980. Shear zones in the granodiorite massifs of the Central Pyrenees and the behaviour of these massifs during the Alpine orogenesis. Journal of Structural Geology, 2, 49-53. LEBLANC, D., GLEIZES, G, Roux, L. & BOUCHEZ, J.L. 1996. Variscan dextral transpression in the French Pyrenees: new data from the Pic des TroisSeigneurs granodiorite and its country rocks. Tectonophysics, 261, 331-345. LISLE, R.J. 1985. Geological strain analysis: a manual for the Rf/(/) technique. Pergamon, Oxford. MARRE, J. 1973. Le complexe eruptif de Querigut. Petrologie, Structurologie, cinematique de mise en place. These Toulouse, 543 pp. MITRA, G. 1979. Ductile deformation zones in Blue Ridge basement rocks and estimation of finite strains. Geological Society of America Bulletin, 90, 935-951. MUNOZ, J.A., MARTINEZ, A. & VERGES, J. 1986. Thrust sequences in the Spanish eastern Pyrenees. Journal of Structural Geology, 8, 399-405. PATERSON, S.R. & MILLER, R.B. 1998. Stoped blocks in plutons: paleo-plumb bobs, viscometers, or chronometers? Journal of Structural Geology, 20, 1261-1272. PATERSON, S.R. & VERNON, R.H. 1995. Bursting the

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bubble of ballooning plutons: A return to nested diapirs emplaced by multiple processes. Geological Society of America Bulletin, 107,1356-1380. PATERSON, S.R., VERNON, R.H. & TOBISCH, O.T. 1989. A review of criteria for the identification of magmatic and tectonic foliations in granitoids. Journal of Structural Geology, 11, 349-363. RAMSAY, J.G & ALLISON, 1.1979. Structural analysis of shear zones in an Alpinised Hercynian granite. Schweizerische Mineralogische und Petrographische Mitteilungen, 59, 251-279. RAMSAY, J.G. & GRAHAM, R.D. 1970. Strain variations in shear belts. Canadian Journal of Earth Sciences, 7, 786-813. RAMSAY, J.G. & HUBBER, M.I. 1983. The techniques of modern structural geology, Vol. 1: Strain analysis. Academic Press, London. RUTTER, E.H. & NEUMANN, D.H.K. 1995. Experimental deformation of partially molten westerly granite under fluid-absent conditions, with implications for the extraction of granitic magmas. Journal of Geophysical Research, 100 (B8), 15697-15715. SAILLANT, J.P. 1982. La faille de Merens (Pyrenees Orientales) microstructures et mylonites. These 3eme cycle, 297 pp, Univ. Paris VII. SIMPSON, C. 1983. Displacement and strain patterns from naturally occurring shear zone terminations. Journal of Structural Geology, 5, 497-506. SIMPSON, C., CARRERAS, J. & LOSANTOS, M. 1982. Inhomogeneous deformation in Roses granodiorite. Acta Geologica Hispanica, 17, 219-226. TOBISCH, O.T. & WILLIAMS, Q. 1998. Use of microgranitoid enclaves as solid state strain markers in deformed granitic rocks: an evaluation. Journal of Structural Geology, 20, 727-743. VAN DER MOLEN, I. & PATERSON, M.S. 1979. Experimental deformation of partially-melted granite. Contributions to Mineralogy and Petrology, 70, 299-318. VERGELY, P. 1970. Etude tectonique des structures pyreneenes du versant sud des Pyrenees orientales. These 3eme cycle, Faculte Sciences, Universite de Montpellier. VIGNERESSE, J.L. &TiKOFF,B. 1999. Strain partitioning during partial melting and crystallizing felsic magmas. Tectonophysics, 312,117-132. WENK, H.R. 1998. Deformation of mylonites in Palm Canyon, California, based on xenolith geometry. Journal of Structural Geology, 20, 559-571. WILLIAMS, Q. & TOBISCH, O.T. 1994. Microgranitoid Enclaves shapes and magmatic strain histories: Constraints from drop deformation theory. Journal of Geophysical Research, 99, 24359-24368. ZWART, HJ. 1979. The geology of the central Pyrenees. Leidse Geologische Mededelingen, 50,1-74.


Shear zones and metamorphic signature of subducted continental crust as tracers of the evolution of the Corsica/Northern Apennine orogenic system G. MOLLI1'2 & R. TRIBUZIO3'4 Dipartimento di Scienze della Terra, Universitd di Pisa, Via S. Maria 53, 1-56126 Pisa, Italy (e-mail:[email protected]) 2 CNR Istituto di Geoscienze e Georisorse,Via G. Moruzzi, 1-56124 Pisa, Italy 3 Dipartimento di Scienze della Terra, Universitd di Pavia, Via Ferrata 1, 1-27100 Pavia, Italy (e-mail: [email protected]) 4 CNR Istituto di Geoscienze e Georisorse, Sezione di Pavia, Via Ferrata 1, 1-27100 Pavia, Italy l

Abstract: This paper focuses on new data concerning the deformation and metamorphic history of continental margins of the Mesozoic Ligurian Tethys. In the Tenda massif (NE Corsica), a slice of the European-Iberian continental margin, contractional shear zones show HP/LT metamorphic assemblages and top-to-the-west kinematics. These shear zones are overprinted by greenschist facies exhumation-related structures showing top-to-the-SW sense of transport and then top-to-the-NE extensional shearing. The presence of HP/LT metamorphism, together with the kinematics of syncontractional shear zones, supports the classic view of Cretaceous-Eocene east-vergent 'alpine-subduction' during the early evolution of the Corsica belt. By taking into account structural and metamorphic data on Tuscan continental units belonging to the other side (Adria margin) of the former Mesozoic Ligurian ocean, we ascribe the Corsica/Northern Apennine system to a polycyclic orogen.

Analyses of crustal scale shear zones and mesoto microscale kinematic criteria were first applied to unravel plate tectonic scale reconstructions in alpine Corsica, and to support a model of intraoceanic subduction blocked by underthrusting of the continental crust beneath the oceanic lithosphere (e.g. Mattauer et al. 1981; Gibbons & Horak 1984; Harris 1985; Warbourton 1986). Recently, however, these tenets have been challenged, and the importance of exhumation-related greenschist structures and kinematics have been explored (Jolivet et al. 1990,1998; Lahondere 1991; Daniel et al. 1996), resulting in uncertainty about the early contractional history (Jolivet et al. 1998; Rossetti et al. 2002). This debate has important implications for the development of the Corsica/Northern Apennine orogenic system (review in Alvarez 1991). Most recent studies (e.g. Lahondere et al. 1999; Padoa & Durand Delga 2001; Bortolotti et al. 2001; Faccenna et al. 2001; Rossetti et al. 2002), with only a few exceptions (Cello et al. 1996; Doglioni et al. 1998; Malavielle et al. 1998; Michard et al. 2002), follow the proposal of Principi & Treves (1984) in considering the evolution of alpine Corsica and Northern Apennines as the

development of an accretionary wedge formed by a continuous Cretaceous to Oligocene westdipping subduction of the Ligure-Piemontese ocean beneath the Corsica/European continental crust, fitting into a monocyclic-type doublevergent orogenic model (Fig. la). In this model, western 'Hercynian' Corsica is considered as the backstop of the Apennine wedge. Some of the ophiolitic units of alpine Corsica are regarded as the deeper part of the accretionary complex, backthrust on the Corsican crust by corner flow (Cowan & Silling 1978) to the rear of the Apennine accretionary system. This model does not take into account previous studies, which considered the early structures in the alpine Corsica units to be produced by an east-dipping 'alpine' subduction (Fig. Ib) (Mattauer & Proust 1975; Caron 1977; GLOM 1977; Mattauer et al. 1981; Faure & Malavielle 1981; Gibbons et al. 1986; Bezert & Caby 1988; Jolivet etal. 1990). We have tackled the problem by studying the larger slices of continental crust that make up alpine Corsica i.e. the Tenda unit. In order to constrain the early deep kinematics and the possible tectonic setting of the Tenda unit, we have focused our attention mainly on shear


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Fig. 1. Tectonic models proposed for the Corsica/Northern Apennine region: (a) continuous Cretaceous-Oligocene/Miocene west-dipping Apenninic subduction (e.g. Principi & Treves 1984; Jolivet et al. 1998; Rossetti et al. 2002); (b) Alpine-type east vergent intraoceanic subduction (Mattauer et al. 1981) followed by Middle Eocene continental collision (e.g. Gibbons & Horak 1984; Warbourton 1986).

zone structures, peak P-T conditions and metamorphic gradients. Furthermore, we have considered the available data on Adria-derived continental nappes (i.e. Tuscan metamorphic units) of the inner Northern Apennine, to analyse the character of the Corsica/Northern Apennine orogenic system from the Cretaceous to Late Oligocene.

Tectonic setting of alpine Corsica Corsica is located in the northern part of the western Mediterranean and is subdivided into two principal geological domains: a western area mainly formed by Late Hercynian granitoids and minor relicts of host rock-basement; and an eastern area characterized by continental and oceanic-derived units deformed during the Alpine orogeny (alpine Corsica) (Fig. 2). Corsica and Sardinia are traditionally regarded as a microblock welded to southern France-northern Iberia until the Middle Oligocene, when rifting and then drifting in the Ligure-Provencal basin took place (Guegen et al 1997; Carminati et al. 1998; Speranza et al 2002; Rollet et al 2002 and references therein). The western Hercynian domain is correlated with the Maures-Esterel basement in southern France, whereas alpine Corsica is regarded as the southern prolongation of the western Alps,

Fig. 2. (a) Tectonic setting of Corsica within the western Mediterranean, (b) Geological setting of the studied areas within the Corsica/Northern Apennine framework.

through the Western Liguria and Voltri Group Alpine units (Durand Delga 1984 and references therein). Corsica (Fig. 3) comprises a series of largescale nappes and can be subdivided into four major composite units (Nardi 1968; Caron 1977; Mattauer et al 1981; Durand Delga 1984; Warbourton 1986). From bottom to top these units are: (1) the autochthon 'Hercynian' Corsica; a mainly undeformed series of the Corsica plate (basement rocks and their sedimentary cover up to Eocene); (2) the deformed part of the Corsica continental margin, i.e. the Tenda massif and the more internal Serra di Pigno/Farinole units (mainly Hercynian granitoids with Permian to Mesozoic volcano-sedimentary and sedimentary cover);

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1984; Lahondere 1991) and by a (retrograde) epidote-blueschistfaciesmetamorphism (Harris 1985; Fournier et al 1991; Lahondere 1991; Caron 1994). No general agreement exists regarding the metamorphic evolution of the Tenda unit. Peak pressure conditions are particularly uncertain. Gibbons & Horak (1984) gave pressure estimates of 0.6-0.9 GPa, based on the celadonite content of phengites. In contrast, Lahondere (1991), Jolivet etal (1998), and Lahondere et al (1999) give values lower than 0.5 GPa, suggesting that the Tenda massif was not involved in the subduction process (Tribuzio & Giacomini 2002).

Structural and metamorphic history of the Tenda massif Geological outline

Fig. 3. (a) Tectonic map of northern Corsica showing the main tectonic units, (b) Geological section through northern Corsica (modified after Jolivet et al 1990; Dalian & Puccinelli 1995; Malavielle et al. 1998).

(3) the 'Schistes Lustres' composite nappe formed by ophiolitic sequences (mantle ultramafics, gabbros, pillow lavas and associated Jurassic to Cretaceous metasediments); and (4) the Balagne/Nebbio/Macinaggio system, i.e. ophiolitic and continental units of internal 'Ligurian-Adria' affinity. These uppermost nappes and their contact with the underlying units are unconformably overlain by Early Miocene sediments (e.g. the St Florent limestone, Dalian & Puccinelli 1995; Ferrandinieffl/. 1998). The external autochthon basement and cover are basically unmetamorphosed or only locally affected by subgreenschist facies recrystallization along fault zones (Egal 1992). The uppermost Balagne/Nebbio/Macinaggio units, show prehnite/pumpellyite facies assemblages in basic rocks. The 'Schistes Lustres' nappe and the Serra di Pigno/Farinole gneissic units are characterized by relict eclogitic association (Dal Piaz & Zirpoli 1979; Caron et al. 1981; Pequignot et al

The Tenda massif (Figs 3 & 4) is exposed for nearly 200 km2 from St Florent-Ostriconi southward to Ponte Leccia, and is characterized by 280-300 Ma old Palaeozoic granitoids, mainly amphibole/biotite granodiorites and leucomonzogranites (Rossi et al 1994). In the southern part, the massif is intruded by a gabbroic complex (the Bocca di Tenda gabbro), which shows well preserved magmatic features (Ohnestetter & Rossi 1985). This complex was dated by single zircon lead evaporation at 274 ± 4 Ma (Lahondere et al 1999) and consists mainly of olivine-gabbronorites, gabbronorites and hornblende-bearing diorites/tonalites (Ohnestetter & Rossi 1985; Tribuzio etal 2001). Granitoids and gabbros are cross-cut by doleritic and peralkaline rhyolite dykes, which both show chilled margins against the host rocks (Tribuzio et al 2001; Tribuzio & Giacomini 2002). Remnants of a pre-Carboniferous basement occur locally in the southern and northwestern side of the massif (Delcey & Meunier 1966; Nardieffl/. 1978; Jourdan 1988; Rossi ef al 1994). The basement rocks are associated with a Permian volcano-sedimentary cover, well exposed in the northern part of the massif (Agriates's Desert). A thin Mesozoic metasedimentary cover (Mattauer et al 1981; Durand Delga 1984) crops out at the eastern and southern sectors of the massif. These sediments are mainly represented by Triassic quarzites and micaceous arkoses, Jurassic dolomitic and calcitic marble, and probable Cretaceous polygenic metaconglomerates, principally formed from basement-derived debris with subordinate carbonatic clasts (Jourdan 1988). The entire

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Fig. 4. Geological map of the Tenda massif with the location of relict HP shear zones and related kinematics, (a) our data (b) after Mattauer et al. (1981) and Jourdan (1988). Map compiled after this work, Delcey & Meunier (1966), Nardi et al. (1978), Jourdan (1988), Dalian & Puccinelli (1995), Rossi et al (1994) and Lahondere et al. 1999.

of the autochthon basement on the Tenda massif. Slickensides on minor fault planes testify to the predominant sinistral strike-slip character of this fault zone, which at least locally appears Deformation history of the Tenda massif to be reworked with normal movement The Tenda massif represents a regional open (Jourdan 1988; Daniel et al. 1996). In the central/southern Tenda (south of antiform (Fig. 3b), bounded by high angle faults in its western and eastern sides. This antiform Urtaca), the regional scale antiform is formed by has an amplitude of nearly 10 km in the northern kilometre scale antiform/synform pairs, locally sector of the massif, with its western limb associated with a steep crenulation to disjunctive deformed by a fault zone in which Tenda- cleavage observable in orthogneiss, as well as in derived rocks (basement and cover) and cover metasediments. The development of these autochthon series (basement and cover) are fold structures, showing local en-echelon trends, cataclastically deformed. Splays of this fault can be related to the activity of wrench faulting zone locally produce east-vergent overthrusting (whose age can be constrained as post-Eocene cover section shows the same metamorphic imprint as the basement (Mattauer et al. 1981).


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Fig. 5. (a) Casta granodiorite, with synmagmatic deformed mafic xenolite. (b) Mylonitic orthogneiss, bounding the Casta granodiorite, with solid state deformed mafic xenolite. (c) Mylonitic blueschist facies metagranite; shear bands and asymmetrical porphyroclasts define a top towards 270° shearing, (d) Photomicrograph of blueschist granitic mylonite. (e) Asymmetrical porphyroclast with synkynematic Na-amphibole growth (detail of (d)). (f) Photomicrograph of mylonitic metadolerite with a well defined shear band system.

and pre-Burdigalian) affecting the boundaries of the Tenda massif (Maluski et al 1973; Waters 1990; Lahondere et al 1999). The geometry of deformation inside the Tenda massif is controlled by the predominance of granitoids deformed at deep structural levels (Ramsay & Allison 1979; Choukroune & Gapais 1983; Gapais et al 1987). A heterogeneous deformation pattern characterizes the entire massif with domains without fabric (i.e. isotropic) or with only a magmatic grain-shape

fabric locally preserved (e.g. the Casta granite, Fig. 5a) surrounded by mylonitic orthogneisses and/or mylonites (Fig. 5b). Foliated granitoids with S-L fabrics show the development of blueschist and/or greenschist facies minerals. The dominant fabric at regional scale is represented by greenschist facies (GS) tectonites. Based on kinematic criteria (stretching/mineral lineations and the associated shear sense), locally controlled by direct overprint relationships, we subdivided the GS structures

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Fig. 6. (a) Examples of the GS2 high strain zone with shear band systems indicating top-to-the-NE orthogneiss east of M. Buggentone, southeastern Tenda. (b) Shear bands in metasedimentary cover near S. Florent (RN D41) East Tenda Shear zone, (c) Conjugate system of centimetre-scale shear zones related to the GS2 deformation stage. North of M. Astu, central Tenda. (d) Example of orthogneiss with top-to-the-west shear bands (RN D62 south F.di Poragghia). (e) Asymmetrical boudinage of aplitic dyke indicating top-to-the-west shearing in orthogneiss south of Bocca di Tenda. All examples are observed in XZ sections.

into two groups related to distinct stages of deformation. The younger GS2 stage is characterized by a predominant top-to-the-NE shearing (Figs 6a,b, & 7a) localized in the eastern part of the massif (East Tenda Shear zone, ETSZ, of Jolivet et al

1990; Daniel et al 1996, but see also Waters 1990) towards the contact with the overlying Schistes Lustres. In the northern Tenda Massif where the ETSZ was first recognized, it comprises an inhomogeneous shear zone formed by several large scale shear bands up to 100 m thick


Fig. 7. Equal area projection (lower hemisphere) of foliation planes (open circles) and stretching/mineral lineations (black circles), (a) Greenschist facies shear zones GS2 top-to-the-NE (in the ETSZ) extensional stage (mean lineation direction 12/048). (b) Greenschist facies shear zones GS1 top-to-the-west stage (mean lineation direction 6/068). (c) High pressure/low temperature shear zones blueschist stage (mean lineation direction 2/090).

affecting granitoids and their metasedimentary cover (Jolivet et al 1990; Fournier et al. 1991;

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Daniel et al. 1996). In the central-southern Tenda (south of Sorio village), the ETSZ consists of at least two significant high strain zones, several metres thick, separated by low strain domains in which subhorizontally crenulated orthogneiss can be recognized. Inside the massif, centimetre-scale shear zones with top-to-theNE kinematics can be related to the same deformation stage of the ETSZ. Conjugate shear zones associated with top-to-the-NE and top-tothe-SW shear sense (Fig. 6c) are present locally and may be related to partitioned coaxial strain far from the ETSZ (cf. Daniel et al 1996). Available radiometric data (Ar/Ar on phengites) constrain the age of the GS2 stage to between 35 and 25 Ma (Maluski 1977; Carpena et al. 1979; Brunet et al. 2000). The ETSZ has been interpreted by Jolivet et al. (1990) and Daniel et al. (1996) as a major extensional shear zone related to crustal-scale thinning bounding the Tenda massif, which is considered to be a metamorphic core complex. However, the pervasive GS1 deformation stage on the massif scale is associated with an ENE/WSW stretching lineation (Fig. 7b) and top-to-the-west kinematics (Fig. 5c,d,e). During this stage, westward thrusting of the Tenda unit over the western and external 'autochthon' units was realized, as observed in the Urtaca tectonic window (see also Nardi et al. 1975; Jourdan 1988 and discussion below). The heterogeneous deformation pattern (Fig. 4) that characterizes the Tenda massif allows us to recognize and analyse domains of mylonitic orthogneisses and mylonitized doleritic and peralkaline rhyolitic dykes (Fig. 5c,d,f) with a well preserved HP/LT fabric that was unaffected by greenschist facies static and/or dynamic retrogression. The HP microfabric in mylonitized granitoids (Fig. 5c,d,e) and rhyolitic dykes show the typical features of quartz-feldspathic mylonites, suggesting that bulk ductile flow during HP metamorphism was accommodated through dislocation creep of quartz. The HP relict structural domains are characterized by east-west orientated stretching and/or mineral lineations defined by sodic amphibole and are associated with a top-to-the-west shear sense (Fig. 7c). These kinematic indicators were observed at kilometre distances (Fig. 4), implying a regional significance to this HP east-west trend and thus strongly constraining the deep deformation history of the Tenda unit. In contrast with previous descriptions (Mattauer et al 1981; Gibbons & Horak 1984), we found HP relicts (structures and/or relict minerals) not only in the eastern upper part of the massif towards the contact of Schistes Lustres,


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Table 1. Significant HP/LT mineral assemblages in different rock types in the Tenda massif (thin lines represent minerals that are only present locally) HIGH PRESSURE ASSEMBLAGES

Qtz-diorites, tonalites and dolerites

Granitoids

Peralkaline rhyolites

Gabbronorites

Na-amphibole (riebeckite-ferroglaucophane) Phengite (Si = 3,5-3,6 apfu) Epidote Chlorite Albite K-feldspar Quartz Na-Cpx (Jd up to 46 mol %) Al-poor (minor than 1,2 apfu) hornblende Titanite Calcite

but distributed throughout the structural thickness of the unit. Where they were still recognizable, for example in the southern area around the Bocca di Tenda gabbro, the geometries of the shear zones fitted the anastomosed distribution pattern expected for crustal scale thrusting.

Metamorphic history of the Tenda massif The wide variety of rocks with different compositions (expecially in the southern part of the massif where the Bocca di Tenda gabbro crops out) allowed the peak metamorphic conditions at 0.8 GPa/300 °C to 1.1 GPa/500 °C to be determined (Tribuzio & Giacomini 2002). The rocks evolved from the gabbroic sequence (quartz diorite/tonalites), basalt doleritic dykes and granitoids are characterized by epidoteblueschist assemblages as they show the coexistence of riebeckite/ferroglaucophane, epidote, celadonite-rich phengite (Si = 3.5-3.6 apfu) and albite (Table 1). The per alkaline rhyolite dykes show an unusual metamorphic paragenesis, defined by jadeite-bearing (up to 46 mol%) aegirine, riebeckite, celadonite-rich phengite (Si = 3.5-3.6 apfu), quartz, albite and K-feldspar. The Mg-rich rocks (olivine gabbronorites to gabbronorites) are characterized by the absence of blue amphibole. Deformed gabbronorites show the development of a mineral association that can be related to the epidote-amphibolite facies, as it displays the coexistence of Al-poor horneblende (Al < 1.2 apfu), albite, epidote and celadonite-rich phengite (Si = 3.5 apfu).

The occurrence of epidote-amphibolite facies assemblages in Mg-rich rocks allow the peak P-T metamorphic conditions to be constrained at 1.0 ± 0.1 GPa and 450 ± 50 °C. These values attest to a geothermal gradient (dT/dP) of 10/13 °C km"1, thus suggesting a subductionrelated tectonic setting (Tribuzio & Giacomini 2002) of 'slow-type' (see discussion below). The age of//P/Lrmetamorphism in the Tenda massif is not well defined. A separate of celadonite-rich phengite (Si = 3.5 apfu) from a deformed granitoid of the Northern Tenda massif has yielded a discordant 39Ar/40Ar spectrum that regularly increases during step-heating, from about 25 Ma to 47 Ma (Brunet et al 2000). This might suggest that the high-pressure metamorphism had a minimum age of 47 Ma. Phengite compositions in deformed GS2 granitoids (Si = 3.2-3.3 apfu) suggest a decompression at pressure lower than 0.5 GPa (Fig. 8). The decompression was most likely coupled with a temperature decrease, as suggested by the amphibole compositional variations (e.g. outward decrease of Al, Na and Ti) in deformed gabbronorites. Ar/Ar investigations on white micas from deformed granitoids show that the youngest greenschist facies recrystallization occurred at around 25 Ma (Brunet etal 2000). In addition, phengitic micas with intermediate Si compositions have given values of 35-37 Ma (Brunet et al 2000). We therefore consider the 47 to 37 Ma period as related to the cooling and concomitant exhumation of the Tenda unit, in agreement with the work of Malavielle et al. (1998).


Fig. 8. Inferred pressure-temperature-time path of the Tenda massif. Gin-out taken from Maresch (1977), the reaction pumpellyite + chlorite —> actinolite + epidote after Liou et al (1983), lawsonite-clinozoisite transition after Barnicoat & Fry (1986), the lower stability limit of barroisite from Ernst (1979), oligoclase in reaction from Maruyama et al (1983), the reaction curve for Na-clinopyroxene (Jd46) + quartz -> albite was calculated with the 3.1 version of THERMOCALC program (Holland & Powell 1998, and references therein). The reaction curve for Mg-phengite (Si = 3.6 apfu) —> quartz + Kfeldspar + phlogopite + H2O) is after Massone & Szpurska (1997). Ar/Ar data after Brunet et al. (2000), Jourdan (1988) and Maluski (1977). Fission track apatite ages (Ap FT) after Janki et al (2000) and Cavazza et al (2001).

To sum up, three major deformation events are recorded in large scale shear zone development of the Tenda massif: (1) early stages of deformation under HPILT metamorphic conditions, recorded by localized shear zones showing top-to-trie-west kinematics. The age of this event is possibly older than 47 Ma; (2) syncontractional exhumation related to greenschist-facies retrogression and with westward thrusting (GS1). This history can be constrained between 47 and 37 Ma; (3) top-to-the-NE extensional shearing (GS2), between 35 and 25 Ma old, with partial reactivation and overprinting of previous fabrics.

Structural and metamorphic evolution of the inner Tuscan metamorphic units A recent synthesis of data and interpretations of the regional geology of the Northern Apennines is reported in Carmignani et al (1995), Jolivet et

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al (1998) and Cerrina Feroni et al (2002). The former western margin of the Adria plate is exposed below the remnants of the Ligurian accretionary wedge on the Thyrrenian side of the Northern Apennines (Ligurian and sub-Ligurian units). This margin is represented by different thrust sheets forming the so-called Tuscan units (Elter 1975). Part of these continental-margin units, mainly associated with cover sequences of Triassic/Late OligoceneMiocene age (e.g. the Tuscan nappe), remained at high structural levels during the whole Apennine history. Other portions of the same continental margin were more deeply underthrusted and are now exposed in tectonic windows (Tuscan metamorphic units, see Fig. 2b) below the overlying lower grade composite nappesystem (Ligurian accretionary wedge units and Tuscan nappe). Some of the Tuscan metamorphic units show high-pressure greenschist-facies peak assemblages (Mg-chloritoid and kyanite in metapelites) that developed at 0.6-0.8 GPa and 400-500 °C (Massa unit in the Alpi Apuane region, Franceschelli et al 1986; Jolivet et al 1998; Molli et al 20000, b). In southern Tuscany, high-pressure/low-temperatureassemblages (Fe-Mg carpholite in metapelites) were recognized. Estimated pressure and temperature conditions vary from 0.6-1.0 GPa and 350-380 °C for the M.Leoni/Monticiano Roccastrada in the Montagnola Senese area (Giorgetti et al 1998) to 1.0-1.2 GPa and 350-420 °C for the Verrucano of Monte Argentario (Theye et al 1997; Jolivet et al. 1998). In these continental units, the structural history is characterized by an early generation of syn- to late-peak metamorphic structures related to contraction and nappe stacking, deformed by younger exhumationrelated structures (Molli et al 2000a,5 and references therein). The early stages of contractional deformation, 27-20 Ma ago (Kligfield etal. 1986; Deino et al 1992; Brunet et al 2000) resulted in SW to NE directed overthrusts and NE-facing tight to isoclinal recumbent folds at regional scale with flat-lying axial surfaces. These structures are traditionally related to the underthrusting of the Adria continental margin following the west-vergent subduction of the relict of the Ligurian Tethys ocean (Carmignani et al 1978; Carmignani & Kligfield 1990).

Discussion The deformation of continental (and oceanic) crust is characteristically heterogeneous in nature and localized linked faults and shear zone systems represent a common tectonic setting

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(Rutter et al. 2001). The results of our analyses in the Tenda massif show how a slice of continental crust was internally deformed during a subduction/exhumation cycle. The tectonic history has been unravelled by studying incomplete shear zone reactivation, associated with an increase in partitioning and localization of the strain during the exhumation, these have been constrained by metamorphic assemblages, kinematics (shear direction and sense of transport), overprinting relationships and radiometric ages. The results of this study carries implications for the tectonic evolution of the Corsica/Northern Apennine system with regard to monocyclic or polycyclic orogenic processes. This study shows that the Corsican crust of the Tenda massif: underwent peak metamorphism in epidote/blueschist facies at about 1 GPa and 450 °C, possibly earlier than 47 Ma; was deformed during subduction with top-to-thewest kinematics; and presently crops out as a core of antiformal stack overthrusted on autochthonous Corsica basement. These features support an east-dipping 'alpine' subduction of the Corsica basement. However, the inner Northern Apennine ophiolitic units (Gorgona, Roselle, Argentario, Giglio) and continental Adria-derived units underwent highpressure/low-temperature metamorphism or high-pressure greenschist facies metamorphism at 27-20 Ma. This supports a polycyclic development of the Corsica/Northern Apennine orogenic system as suggested by Boccaletti et al. (1971), Elter & Pertusati (1973), Dal Piaz (1974), Reutter et al. (1978), Doglioni et al (1998) and Michard et al. (2002). For the Corsica/Apennine area we suggest the following tectonic evolution (Fig. 9): (1) a Cretaceous stage of east-dipping intraoceanic subduction beneath the Nebbio microcontinent, which was related to ESE movement of the Iberian plate. This is recorded in the Corsican Schistes Lustres by development of eclogites dated at 84 ± 5 Ma (Lahondere & Guerrot 1997). A low to very low geothermal gradient for these ophiolitic units is shown by local occurrence of eclogites with coexisting almost pure jadeite and quartz (Caron et al. 1981; Lahondere 1988) and by lawsonite-bearing eclogites (Pequignot et al 1984; Lahondere 1991; Padoa 2001), which we ascribe to a fast and steep subduction; (2) (Fig. 9a) the subduction pulled down the thinned Corsican continental margin starting from the Late Cretaceous, as testified by 65 Ma old ages of continental derived

eclogites (Farinole-Serra di Pigno unit, Brunet et al 2000). The continental margin was progressively subducted, reducing the subduction rate and increasing the geothermal gradient (Tribuzio and Giacomini 2002); (3) (Fig. 9b) the involvement of thick continental crust (Tenda massif) blocked the subduction during Paleocene-Middle Eocene. Due to buoyancy forces and shearing, the Corsican crust failed and started to be exhumed, whilst break-off and detachment of subducted oceanic lithosphere took place (Van den Beukle 1992; Von Blackenburg and Davis 1995; Chemenda et al 1996; Malavielle et al 1998; Tribuzio and Giacomini 2002). The presence of a still open oceanic domain eastward of the Corsican accretionary wedge (central Tuscany and southern Italy) allowed subduction flip and the beginning of westward subduction to drive the early development of the Apennines (Fig. 9c). This event was followed by the Oligocene-Miocene calc-alkaline volcanism on the western side of Sardinia-southern Corsica, the associated back-arc rifting (30-21 Ma), the formation of oceanic crust (21-16 Ma) in the Ligure-Provencal basin, and the rotation of the Sardinian-Corsica block far from Iberia/Europe mainland (Scandone 1979; Rehault et al 1984; Serri et al. 1993; Speranza et al 2002 and references therein). The exhumation of the Tenda unit is considered as mainly related to a syncontractional tectonic history (Fig. 9b) during the early Alpine cycle, and only the latest stages (realized within the back-arc extensional/transtensional frame) are related to the Apennine subduction (Fig. 9c, d), in agreement with Malavielle etal (1998). In conclusion, our data on shear zones and the metamorphic signature of the Tenda Massif reaffirm alpine Corsica as a case study for intraoceanic subduction blocked by the underthrusting of the continental crust, as suggested more than 20 years ago by Mattauer et al (1981). Our ongoing work is focused on quantitative analyses of structures in terms of volume of deformed domains, evolution of their rheology from rock microfabrics and deformation/metamorphism/rock strength relationships. P. Elter and A. Puccinelli are thanked for their discussion and suggestions on the geology of Corsica. M. Krabbendam and M. Sandiford provided useful comments and suggestions during their review of the submitted manuscript. Editorial handling and remarks


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Fig. 9. Tectonic evolution of Alpine Corsica within the framework of the Corsica/Northern Apennine erogenic system, (a) Subduction of the distal Corsican margin and eclogite metamorphism (Late Cretaceous), (b) Involvement of thick Corsican crust (Tenda), blocking of subduction, slab-break-off below 50 km and flip of subduction polarity, (c) Oligocene regional upper plate crustal extension related to Apenninic subduction; (d) Oligocene/pre-Early Miocene wrench and then late normal faulting possibly connected with Corsica/Sardinia rotation and the beginning of a retreat of Apenninic subduction zone. of I. Alsop are also greatly appreciated. This work was supported by the Universities of Pavia and Pisa, CNR and COFIN funds. G.M is an external collaborator of

RETREAT project, Continental Dynamics Program of National Science Foundation (publ. N°3). This paper is dedicated to the memory of our friend Graziano Plesi.

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References ALVAREZ, W. 1991. Tectonic evolution of the CorsicaApennines-Alps region studied by the method of successive approximations. Tectonics, 10,936-947. BARNICOAT, A.C & FRY, N. 1986. High-pressure metamorphism of the Zermatt-Saas ophiolite zone. Journal of the Geological Society, London, 143, 607-618. BEZERT, P. & CABY, R. 1988. Sur 1'age post-bartonien des evenements tectono-metamorphiques alpines en bourdure orientale de la Corse cristalline. Bulletin de la Societe Geologique de France, 8, 965-971. BOCCALETTI, M., ELTER, P. & GUAZZONE, G.J.P. 1971. Plate tectonic models for the development of the western Alps and Northern Apennines. Nature, 234,108-111. BORTOLOTTI, V, FAZZUOLI, M., PANDELI, E., PRINCIPI, G, BABBINI, A. & CORTI, S. 2001. Geology of central and eastern Elba island, Italy. Ofioliti, 26, 97-150. BRUNET, C, MONIE, P., JOLIVET, L. & CADET, IP. 2000. Migration of compression and extension in the Thyrrhenian Sea, insights from 40Ar/39Ar ages on micas along a transect from Corsica to Tuscany. Tectonophysics, 321,127-155. CARMIGNANI, L. & KLIGFIELD, R. 1990. Crustal extension in the Northern Apennines: the transition from compression to extension in the Alpi Apuane Core Complex. Tectonics, 9,1275-1303. CARMIGNANI, L., GIGLIA, G & KLIGFIELD, R. 1978. Structural evolution of the Apuane Alps: an example of continental.margin deformation in the Northern Apennine. Journal of Geology, 86, 487-504. CARMIGNANI, L., DECANDIA, F.A., DISPERATI, L., FANTOZZI, PL., LAZZAROTTO, A., LIOTTA, D. & OGGIANO, G 1995. Relationships between the Tertiary structural evolution of the SardiniaCorsica-Provengal Domain and the Northern Apennines. Terra Nova, 7,128-137. CARMINATI, E., WORTEL, M.J.R., SPAKMAN, W. & SABADINI, R. 1998. The role of slab detachment process in the opening of the western-central Mediterranean basins: some geological and geophysical evidence. Earth and Planetary Science Letters, 160, 651-665. CARON, J.M. 1977. Lithostratigraphie et tectonique des Schistes Lustres dans les Alpes Cottiens septentrionales et en Corse orientale. Memoire Universite Louis Pasteur Strasbourg, 326 pp. CARON, J.M. 1994. Metamorphism and deformation in Alpine Corsica. Schweizerische Mineralogische und Petrographische Mitteilungen, 7,105-114. CARON, J.M., KIENAST, J.R. & TRIBOULET, C. 1981. High-pressure/low-temperature metamorphism and polyphase Alpine deformation at Sant'Andrea di Cotone (Eastern Corsica). Tectonophysics, 78, 419-451. CARPENA, J., MAILLHE, D., NAESER, C.W. & POUPEAU, G 1979. Sur la datation par traces de fission d'une phase tectonique d'age Eocene superieur en Corse. Compte Rendu Academic Science Paris, serie II, 289, 829-832.

CAVAZZA, W., ZATTIN, M., VENTURA, B. & ZUFFA, G.G 2001. Apatite fission-track analysis of Neogene exhumation in northern Corsica (France). Terra Nova, 13, 51-57. CELLO, G, INVERNIZZI, C. & MAZZOLI, S. 1996. Structural signature of tectonic process in the Calabrian Arc, southern Italy: Evidence from oceanic-derived Diamante-Terranova unit. Tectonics, 15,187-200. CERRINA FERONI, A., MARTELLI, L., MARTINELLI, P., OTTRIA, G. & CATANZARITI, R. 2002. Carta Geologico-Strutturale dell'Appennino EmilianoRomagnolo-Note Illustrative. Regione EmiliaRomagna, Selca Firenze. CHEMENDA, A.I., MATTAUER, M. & BOKUN A.N. 1996. Continental subduction and a mechanism for exhumation of high-pressure metamorphic rocks: new modelling and field data from Oman. Earth and Planetary Science Letters, 143,173-182. CHOUKROUNE, P. & GAPAIS, D.C. 1983. Strain pattern in Aar granite (Central Alps): orthogneiss developed by bulk inhomogeneous flattening. Journal of Structural Geology, 5, 411-418. COWAN, D. & SILLING, R.M. 1978. A dynamic scaled model of accretion at trenches and its implications for the tectonic evolution of subduction complexes. Journal of Geophysical Research, 83, 5389-5396. DAL PIAZ, GV. 1974. Le metamorphisme alpine de haute pression et basse temperature dans 1'evolution structurale du bassin ophiolitique alpinoapenninique. le partie. Bollettino della Societd Geologica Italiana, 93, 437-468. DAL PIAZ, GV. & ZIRPOLI, G. 1979. Occurrence of eclogites relics in the ophiolite nappe from Marine d'Albo, Northern Corsica. Neues JahrbuchfurMineralogie, 3,118-122. DALLAN, L. & PUCCINELLI, A. 1995. Geologia della regione tra Bastia e St-Florent (Corsica Settentrionale). Bollettino della Societd Geologica Italiana, 114, 23-66. DANIEL, J.M., JOLIVET, L., GOFFE, B. & POINSOTT, C. 1996. Crustal-scale strain partitioning: footwall deformation below the Alpine Oligo-Miocene detachment of Corsica. Journal of Structural Geology, 18,1841-1859. DEINO, R., KELLER, IV.A., MINELLI, G. & PIALLI, G 1992. Datazioni Ar/Ar del metamorfismo dell'unita di Ortano-Rio Marina (Isola d'Elba). Studi Geologici Camerti, Vol. Spe. CROP 01, 187-192. DELCEY, R. & MEUNIER, R. 1966. Le massif du Tenda (Corse) et ses bourdes: la serie volcano-sedimentaire, les gneiss et les granites; leurs rapports avec les schistes lustres. Bulletin de la carte Geologique de la France, 278, 237-251. DOGLIONI, C., MONGELLI, F. & PIALLI, G. 1998. Boudinage of the Alpine belt in the Apenninic back-arc. Bollettino della Societd Geologica Italiana, 118, 75-89. DURAND-DELGA, M. 1984. Principaux traits de la Corse Alpine et correlations avec les Alpes Ligures. Memorie della Societd Geologica Italiana, 28,285-329. EGAL,E. 1992. Structures and tectonic evolution of the

SHEAR ZONES AND METAMORPHISM external zone of Alpine Corsica. Journal of Structural Geology, 14,1215-1228. ELTER, P. 1975. Introduction a la geologic de 1'Apennin septentrional. Bulletin de la Societe Geologique de France, 7, 956-962. ELTER, P. & PERTUSATI, PC. 1973. Considerazioni sul limite Alpi-Appennino e sulle sue relazioni con 1'arco delle Alpi Occidentali. Memorie delta Societd Geologica Italiana, 12, 359-375. ERNST, W.G. 1979. Coexisting sodic and calcic amphiboles from high-pressure metamorphic belts and the stability of barroisitic amphibole. Miner alogical Magazine, 43, 269-278. EVANS, M. 1986. Phase relationships of epidoteblueschists. Lithos, 25, 3-23. FACCENNA, C, BECKER, T.W., LUCENTE, F.P., JOLIVET, L. & ROSSETTI, F. 2001 History of subduction and back-arc extension in Central Mediterranean. Geophysical Journal International, 145, 809-820. FAURE, M. & MALAVIELLE, J. 1981. Etude structural d'un cisallement ductile: le chariage ophiolitique Corse dans la region de Bastia. Bulletin de la Societe Geologique de France, 23, 335-343. FERRANDINI, M., FERRANDINI, J., LOYE-PILOT, M.D., BUTTERLIN, J. CRAVETTE, J. & JANIN, M.C. 1998. Le Miocene du Bassin de Saint-Florent (Corse): Modalites de la transgression du Burdigalien Superior et mise en evidence du Serravalien. Geobios, 31,125-137. FOURNIER, M., JOLIVET, L., GOFFE, B. & DUBOIS, R. 1991. The Alpine Corsica metamorphic core complex. Tectonics, 10,1173-1186. FRANCESCHELLI, M., LEONI, L., MEMMI, I. & PUXEDDU, M. 1986. Regional destribution of Al-silicates and metamorphic zonation in the low-grade Verrucano metasediments from the Northern Apennines, Italy. Journal of Metamorphic Geology, 4, 309-321. GAPAIS, D., BALE, P., CHOUKROUNE, P., COBBOLD, PR., MAHJOUB, Y. & MARQUER, D. 1987. Bulk kinematics from shear zone patterns: some filed examples. Journal of Structural Geology, 9, 635-646. GIBBONS, W. & HORAK, J. 1984. Alpine metamorphism of Hercynan hornblende granodiorite beneath the blueschist facies schistes lustres nappe of NE Corsica. Journal of Metamorphic Geology, 2, 95-113. GIBBONS, W., WATERS, C & WARBOURTON, J. 1986. The blueschist facies schistes lustres of Alpine Corsica: A review. Geological Society of America Memoir, 164, 301-331. GIORGETTI, G, GOFFE, B., MEMMI, I. & NIETO, F. 1998. Metamorphic evolution of Verrucano metasediments in northern Apennines: new petrological constraints. European Journal of Mineralogy, 9, 859-873. GLOM. 1977.1 complessi ofiolitici e le unita cristalline della Corsica alpina. Ofioliti, 2, 265-324. GUEGUEN, E., DOGLIONI, C. & FERNANDEZ, M. 1997.

Lithospheric boudinage in the western Mediterranean back-arc basin. Terra Nova, 9,184-187. HARRIS, J. 1985. Progressive and polyphase deformation of the Schistes Lustres in Cap Corse,

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Alpine Corsica. Journal of Structural Geology, 7, 637-650. HOLLAND, T.J.B. & POWELL, R. 1998. An internally consistent thermodynamic data set for phases of petrological interest. Journal of Metamorphic Geology, 16, 309-344. JAKNI, B., POUPEAU, G, SOSSON, M., Rossi, P., FERRANDINI, J. & GUENNOC, P. 2000. Denudations cenozoiques en Corse: une analyse thermochronologique par traces de fission sur apatites. Comptes Rendus de I'Academic des Sciences Paris, 331, 775-782. JOLIVET, L., DUBOIS, R., FOURNIER, M., MICHARD, A. & JOURDAN, C. 1990. Ductile extension in Alpine Corsica. Geology, 18,1007-1010. JOLIVET, L., FACCENNA, C., GOFFE, B., MATTEI, M., ROSSETTI, F, BRUNET, C., STORTI, F, FUNICIELLO, R., CADET, J.P., D'AGOSTINO, N. & PARRA,T. 1998. Midcrustal shear zones in postorogenic extension: Example from the northern Tyrrhenian Sea. Journal of Geophysical Research, 103 (B6), 12123-12160. JOURDAN, C. 1988. Balagne orientale et massif du Tenda (Corse septentrionale): etude structurale, interpretation des accidents et des deformatios, reconstritutions geodynamiques. These de 3 erne cycle, Universite de Paris-Sud, Orsay. KLIGFIELD, R., HUNZIKER, J, DALLMEYER, R.D. & SCHAMEL, S. 1986. Dating of deformational phases using K-Ar and 40Ar/39Ar techniques: results from the Northern Apennines. Journal of Structural Geology, 8, 781-798. LAHONDERE, D. 1988. Le metamorphisme eclogitique dans les orthogneiss, et les metabasites ophiolitiques de la region de Farinole (Corse). Bulletin de la Societe Geologique de France, 4, 579-585. LAHONDERE, D. 1991. Les schistes blues et les eclogites a lawsonite des unites continentals et oceanique de la Corse alpine: Nouvelles donnee petrologique et structurales. (Corse). These de 3 erne cycle, Universite de Montpellier II, France, 262 pp. LAHONDERE, D. & GUERROT, C. 1997. Datation Nd-Sm du metamorphisme eclogitique en Corse alpine: un argument pour 1'existence, au Cretace superieur, d'une zone de subduction active localisee le long du block corse-sarde. Geologic de la France, 3, 3-11. LAHONDERE, D., Rossi, P. & LAHONDERE, J.C. 1999. Structuration alpine d'une marge continentale externe: le massif du Tenda (Haute-Corse). Implications geodynamiques au niveau de la transversale Corse-Apennines. Geologic de la France, 4, 27-44. Liou, J.G., KIM, H.S. & MARUYAMA, S. 1983. Prehniteepidote equilibria and their petrologic applications. Journal of Petrology, 24, 321-342. MALAVIELLE, J., CHEMENDA, A. & LARROQUE, C. 1998. Evolutionary model for Alpine Corsica: mechanism for ophiolite emplacement and exhumation of high-pressure rocks. Terra Nova, 10, 317-322. MALUSKI, H. 1977. Application des methods de datation 39Ar/40Ar aux mineraux des roches cristallins perturbes par les evenementes thermiques et

334


tectoniques en Corse. These de 3 erne cycle, Universite Montpellier, France. MALUSKI, H., MATTAUER, M. & MATTE, P. 1973. Sur la presence de decrochement alpins en Corse. Comptes Rendus de I'Academie des Sciences Paris, serie D, 276, 709-712. MARESCH, W.V. 1977. Experimental studies on glaucophane: an analysis of present knowledge. Tectonophysics, 43,109-125. MARUYAMA, S., SUZUKI, K. & Liou, J.G. 1983. Greenschist-amphibolite transition equilibria at low pressures. Journal of Petrology, 24, 583-604. MASSONE, H.J. & SZPURSKA, Z. 1997. Theromodynamic properties of white micas on the basis of high-pressure experiments in the systems K2OMgO-Al2O3-SiO2-H2O and K2O-FeO-Al2O3Si02-H20. Lithos, 41, 229-250. MATTAUER, M. & PROUST, F. 1975. Sur quelques problemes generaux de la chaine alpine en Corse. Bulletin de la Societe Geologique de France, 18, 1177-1178. MATTAUER, M., FAURE, M. & MALAVIELLE, J. 1981. Transverse lineation and large-scale structures related to Alpine obduction in Corsica. Journal of Structural Geology, 3, 401-409. MICHARD, A., CHALOUAN, A., FEINBERG, H., GOFFE, B. & MONTIGNY, R. 2002. How does the Alpine belt end between Spain and Morocco? Bulletin de la Societe Geologique de France, 173, 3-15. MOLLI, G, CONTI, P., GIORGETTI, G, MECCHERI, M. & OESTERLING, N. 2000. Microfabric studies on the deformational and thermal history of the Alpi Apuane marbles (Carrara marbles), Italy. Journal of Structural Geology, 22,1809-1825. MOLLI, G, GIORGETTI, G. & MECCHERI, M. 2000. Structural and petrological contraints on the tectono-metamorphic evolution of the Massa Unit (Alpi Apuane, NW Tuscany, Italy). Geological Journal, 35, 251-264. NARDI, R. 1968. Le unita alloctone della Corsica e la loro correlazione con le unita delle Alpi e dell' Appennino. Memorie della Societd Geologica Italiana, 1, 323-344. NARDI, R., PUCCINELLI, A. & VERANI, M. 1978. Carta geologica della Balagne "sedimentaria" (Corsica) alia scala 1:25.000 e note illustrative. Bollettino della Societd Geologica Italiana, 97, 3-22. OHNESTETTER, M. & Rossi, P. 1985. Reconsitution d'une paleochambre magmatique expectionalle dans le complexe basique-ultrabasique du Tenda. Comptes Rendus de I'Academie des Sciences Paris, serie II, 300, 853-858. PADOA, E. 2001. Lawsonite-bearing eclogites from the Northern Castagniccia area (Alpine Corsica, France). 3° Forum FIST, Abstract volume, 561-563. PADOA, E. & DURAND-DELGA, M. 2001. L'unite du Rio Magno en Corse Alpine: elements des ligurides de 1'Appennin septentrional. Comptes Rendus de I'Academie des Sciences Paris, 333, 285-293. PEQUIGNOT, G, LARDEAUX, J.M. & CARON, J.M. 1984. Recrystallization d'eclogites de basse temperature dans les metabaltes corse. Comptes Rendus Academic des Sciences Paris, serie II, 299,871-874.

PRINCIPI, G. & TREVES, B. 1984. II sistema corso-appenninico come prisma d'accrezione. Riflessi sul problema generale del limite Alpi-Appennini. Memorie della Societd Geologica Italiana, 28, 549-576. RAMSAY, J.G. & ALLISON, J. 1979. Structural analysis of shear zones in an Alpinised Hercynian granite. Schweizerische Mineralogische und Petrographische Mitteilungen, 59, 251-279. REHAULT, J.P, MASCLE, J. & BOILLOT, G. 1984. Evolution geodynamique de la Mediterranee depuis 1'Oligocene. Memorie della Societd Geologica Italiana, 27, 85-96. REUTTER, K.J., GUNTHER, K. & GROSCURTH, J. 1978. An approch to the Geodynamics of the CorsicaNorthern Apennines double orogene. In: CLOSS, H., ROEDER, D. & SCHMIDT, K. (eds) Alps, Apennines, Hellenides, IUGG Scientific Report, 299-311. ROLLET, N, DEVERECHERE, J., BESLIER, M.-O., GUENNOC, P., REHAULT, J.-P, SOSSON, M. & TRUFFET, C. 2002. Back arc extension, tectonic inheritance, and volcanism in the Ligurian Sea, Western Mediterranean. Tectonics, 21,1-23. ROSSETTI, F, FACCENNA, G, JOLIVET, L., GOFFE, B. & FUNICIELLO, R. 2002. Structural signature and exhumation P-T-t paths of the blueschist units exposed in the interior of the Northern Apennine chain, tectonic implication. Bollettino della Societd Geologica Italiana, volume speciale n.l, 20,829-842. Rossi, P., COCHERIE, A. & LAHONDERE, D. 1992. Relations entre les complex mafiques-ultramafiques et le volcanisme andesitique stefanopermian de Corse occidentale, temoins des phenomenes d'amicissement crustal neovarisques. Comptes Rendus de I'Academie des Sciences Paris, serie II, 315, 1341-1348. Rossi, P., LAHONDERE, J.C., LUCH, D., LOYE-PILOT, M.D. & JACQUET, M. 1994. Carte geologique de la France a 1/50.000, feuille Saint-Florent. BRGM. RUTTER, E.H., HOLDSWORTH, R.E. & KNIFE, E.J. 2001. The nature and tectonic significance of fault-zone weakening: an introduction. In: HOLDSWORTH, R.E., STRACHAN, R.A., MAGLOUGHLIN, J.F. & KNIPE, R.J. (eds) The nature and tectonic significance of fault zone weakening. Geological Society, London, Special Publication, 186, 65-84. SCANDONE, P. 1979. Origin of the Thyrrenian Sea and the Calabrian arc. Bollettino della Societd Geologica Italiana, 98, 27-34. SERRI, G, INNOCENTI, F. & MANETTI, P. 1993. Geochemical and petrological evidence of the subduction of delaminated Adriatic continental lithosphere in the genesis of the Neogene-Quaternary magmatism of central Italy. Tectonophysics, 223,117-147. SPERANZA, F, VILLA, I.M., SAGNOTTI, L., FLORINDO, F, COSENTINO, D., ClPOLLARI, P. & MATTEI, M. 2002.

Age of the Corsica-Sardinia rotation and LigureProvencal Basin spreading: new paleomagnetic and Ar/Ar evidence. Tectonophysics, 6571, 112-133. THEYE, T, REINHARDT, J., GOFFE, B., JOLIVET, L. &

SHEAR ZONES AND METAMORPHISM BRUNET, C. 1997. Ferro and magnesiocarpholite from Mt Argentario (Italy): first evidence for high-pressure metamorphism of the metasedimentary Verrucano sequence, and significance for P-T path reconstruction. European Journal of Mineralogy, 9, 859-873. TRIBUZIO, R. & GIACOMINI, F. 2002. Blueschist facies metamorphism of peralkaline rhyolites from the Tenda crystalline massif (northern Corsica): evidence for involvement in the Alpine subduction event? Journal of Metamorphic Geology, 20, 513-526. TRIBUZIO, R., PATTI, G. & DALLAI, L. 2001. Interaction between mantle and continental crust liquids in the Bocca di Tenda Gabbroic complex. Journal of Conference Abstracts, EUG XI, 817.

335

VAN DEN BEUKEL J. 1992. Some thermomechanical asepcts of the subduction of continental lithosphere. Tectonics, 11, 316-329. VON BLACKENBURG, F. & DAVIES, J.H. 1995. Slabbreakoff: a model for syncollisional magmatism and tectonics in the Alps. Tectonics, 14,120-131. WARBOURTON, J. 1986. The ophiolite-bearing Schistes Lustres nappe in Alpine Corsica: A model for emplacement of ophiolites that have suffered HP/LT metamorphism. Geological Society of America Memoir, 164, 313-331. WATERS, CN. 1990. The Cenozoic tectonic evolution of alpine Corsica. Journal of the Geological Society, London, 147, 811-824.


Crenulation-slip development in a Caledonian shear zone in NW Ireland: evidence for a multi-stage movement history D. M. CHEW1, J. S. DALY2, M. J. FLOWERDEW3, M. J. KENNEDY2 & L. M. PAGE4 5 l Departement de Mineralogie, Universite de Geneve, Rue des Maraichers 13, CH-1205 Geneve, Switzerland (e-mail: [email protected]) ^Department of Geology, University College Dublin, Belfield, Dublin 4, Ireland ^British Antarctic Survey, c/o NERC Isotope Geosciences Laboratory, Kingsley Dunham Centre, Keyworth, Nottingham NG12 5GG, UK ^Laboratory of Isotope Geology, Vrije Universiteit, De Boelelaan 1085,1081 HV Amsterdam, the Netherlands 5 Present address: Department of Geology, Lund University, Solvegatan 13, 223 62 Lund, Sweden Abstract: In Scotland and Ireland, a Laurentian passive margin sequence, the Dalradian Supergroup, was deformed during the c. 470-460 Ma Grampian orogeny, resulting in the formation of crustal-scale recumbent nappes. In Ireland, this passive margin sequence is in general bounded to the SE by the Fair Head-Clew Bay Line (FHCBL), a segment of a major lineament within the Caledonides. Adjacent to the FHCBL, Dalradian metasediments in two separate inliers have undergone post-Grampian strike-slip movement, with the initially flat-lying Grampian nappe fabric acting as a decollement-like slip surface in both cases. As the orientation of these foliation slip surfaces was oblique to the local shear plane in both inliers, displacement along these pre-existing foliation surfaces was also accompanied by crenulation slip. However, the crenulation-slip morphologies produced imply the opposite sense of movement in the two inliers. 40Ar-39Ar dating of muscovite defining the crenulation-slip surfaces indicates that post-Grampian dextral displacement took place along the FHCBL at 448 ± 3 Ma. A subsequent phase of sinistral movement along the FHCBL took place at c. 400 Ma, based on previously published Rb-Sr muscovite ages for synkinematic pegmatites. The kinematic information obtained from crenulationslip morphologies combined with geochronology can thus be used to constrain the reactivation history of a major crustal-scale shear zone.

Deforming materials are seldom isotropic, and hence anisotropy plays an important role in partitioning strain in shear zones. Common examples of anisotropy encountered in midcrustal shear zones include planar elements such as sedimentary layering or foliations. Such planar anisotropies should act as decollementlike surfaces during shear deformation when they are suitably orientated (i.e. subparallel to the shear plane). Renewed movement along pre-existing foliations (foliation reactivation) is thus likely to be a feature of many mid-crustal shear zones. Kinematic models have been used (e.g. Dennis & Secor 1987) to predict the structural features produced when slip occurs along foliation surfaces which are oblique to the walls of a shear zone. Oblique slip produces a displacement component normal to the zone wall, which is inconsistent with plane strain, simple shear

deformation (Ramsay & Graham 1970). In order to maintain the initial thickness of the shear zone and thus preserve a simple shear deformation path, crenulation slip has been interpreted to compensate for this normal displacement component (Dennis & Secor 1987). This paper presents structural and 40Ar-39Ar and Rb-Sr isotopic data from a major shear zone within the Caledonides of NW Ireland, the Fair Head-Clew Bay Line. Detailed field mapping has demonstrated that the main regional foliation developed within Dalradian Supergroup metasediments adjacent to the Fair Head-Clew Bay Line is used as a slip surface within this shear zone, where it is accompanied by crenulation slip. The age of the main regional foliation is also well constrained by previous geochronological studies based immediately outside this shear zone (e.g. Flowerdew et al. 2000; Chew et al. 2003), where it is unaffected by later


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D.M. CHEW ETAL.

Fig. 1. (a) Regional geology of NW Ireland displaying localities referred to in the text, (b) Location map of Ireland within the Caledonides.

deformation (i.e. shearing). Isotopic dating of both the crenulation-slip fabrics and the reactivated foliation within the shear zone enable individual phases of reactivation to be constrained temporally. The crenulation morphologies predicted by the model described above can thus be used to identify not only shear sense on a major crustal-scale shear zone, but also to establish the timing of movement.

Geological significance of the Fair Head-Clew Bay Line In NW Ireland, a Laurentian passive margin sequence, the Neoproterozoic-Cambrian Dalradian Supergroup (Fig. 1), was deformed during the c. 470-460 Ma Grampian orogeny (Friedrich et al 19990, b\ Flowerdew et al 2000). This orogenic episode is believed to be related to the collision of the Laurentian margin with an outboard oceanic arc and associated forearc ophiolite (Dewey & Shackleton 1984), which resulted in metamorphism and the production of crustalscale recumbent nappes within the Dalradian sequence.

The Dalradian Supergroup in NW Ireland (with the exception of the allochthonous Connemara terrane) is bounded to the SE by the Fair Head-Clew Bay Line (Fig. 1). This structure is believed to be equivalent to the Highland Boundary Fault of Scotland and the Baie Verte-Brompton Line of Newfoundland and as such is a significant lineament within the Caledonides. It is believed to represent the original collisional suture between the deformed and metamorphosed Laurentian margin sequences and the outboard oceanic arc (Dewey & Shackleton 1984). The Fair Head-Clew Bay Line itself is defined by a conspicuous magnetic lineament (Max & Riddihough 1975) from northeastern Ireland (Fig. 1) to the north shore of Clew Bay on the western Irish coast (Fig. 2a), whereas the main surface expression of the Fair Head-Clew Bay Line is a fault zone which in general lies about 10 km to the south of the magnetic lineament (Fig. 1). Both are often collectively referred to as the Fair Head-Clew Bay Line (Ryan et al 1995). Throughout most of the southeastern margin of the Dalradian Supergroup in Scotland and Ireland, ductile structures related to movement

REACTIVATION OF A CALEDONIAN SHEAR ZONE

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Fig. 2. (a) Regional geology of Co. Mayo displaying localities referred to in the text, (b) North-south crosssection (X-Y) across the NW Mayo inlier displaying major F2 folds, (c) NW-SE cross section across the Central Ox Mountains inlier adapted from MacDermot et al. (1996) and displaying major F3 folds.

along the Highland Boundary Fault or the Fair Head-Clew Bay Line are rarely exposed. However, in Dalradian Supergroup metasediments on Achill Island in NW Mayo and in the Central Ox Mountains (Fig. 1), ductile structures which can be related to strike-slip motion along the Fair Head-Clew Bay Line (FHCBL) are clearly observed.

Evidence for strike-slip motion along the FHCBL in NW Mayo (South Achill) Achill Island is situated on the southern margin of the NW Mayo inlier (Figs 1 & 2a). The NW Mayo inlier preserves an excellently exposed transect from Laurentian basement, the Annagh Gneiss Complex (Daly 1996), through presumed para-autochthonous (Winchester 1992) Laurentian cover, the Dalradian Supergroup, to outboard oceanic elements located immediately to the south of the FHCBL (e.g. the Clew Bay Complex; Williams et al. 1994). Polyphase deformation is pervasive throughout the Dalradian Supergroup outcrop in Ireland and Scotland. In the NW Mayo inlier, a D! deformation event is responsible for the bulk of the high-strain observed, and is associated with the development of tectonic slides and locally-developed isoclinal Fj folds (Kennedy 1980). The D2 deformation event in the NW Mayo inlier is the main nappe-forming phase. The D2 nappes in general plunge gently east and 'root' in the basement core, the Annagh Gneiss Complex (Fig. 2b). Adjacent to, and directly

above the Annagh Gneiss Complex, the D2 nappes are upward-facing (Fig. 2b); to the south of this 'root zone' recumbent D2 folds face south (Fig. 2b; Kennedy 1980) and are rotated into a downward-facing orientation approaching the southern margin of the inlier (Fig. 2b; Chew 2003). The production of downward-facing D2 folds and the associated S2 strike-swing (Fig. 2a) in the southern part of the inlier has been attributed to either later modification of the D2 nappe pile by an east-west dextral shear zone running along the north margin of Clew Bay (Sanderson et al. 1980; Chew 2003), or a steep zone of D2 dextral transpression contemporaneous with the development of flat-lying D2 nappes to the north (Harris 1993,1995). The model of Sanderson et al. (1980) assumed that the S2 nappe fabric was progressively rotated into the proposed shear zone rather than new fabrics forming. However, detailed mapping of the southern part of Achill Island (South Achill) and the island of Achill Beg (Fig. 3) on the southern margin of the inlier demonstrates that the S2 nappe fabric is modified by 03 structures consistent with dextral strike-slip displacement. The northern boundary of the shear zone to the north is not sharply defined, but D3 structures gradually become less abundant to the north of Ashleam Bay in South Achill (Fig. 3). Two discrete elements have been recognized in the D3 deformation episode, asymmetrical buckle folds with axial planes anticlockwise to the S2 foliation, and extensional crenulation cleavages which cut the S2 foliation in a clockwise sense.

340

D.M. CHEW^r^L.

Fig. 3. Geological map of South Achill and Achill Beg.


341

Fig. 4. Angular relationships predicted between the shear zone wall and the foliation-slip and crenulation-slip surfaces, (a) Reverse-slip crenulations. (b) Normal-slip crenulations. Reprinted from Journal of Structural Geology, 9, Dennis & Secor: A model for the development of crenulations in shear zones with applications from the southern Appalachian Piedmont, pp. 809-817. Copyright 1987, with permission from Elsevier.

Crenulation-slip morphologies produced by oblique foliation-slip The kinematic models of Dennis & Secor (1987, 1990) predict the crenulation morphologies that develop in order to compensate for the displacement component of foliation slip normal to the shear zone wall. In a dextral shear zone, when the pre-existing foliation is at an acute, clockwise angle to the shear zone wall (Fig. 4a), movement away from the shear zone wall due to oblique foliation slip is compensated by reverseslip crenulations (RSC), which transfer slip up to 'higher' foliation planes. When the slipping foliation is at an acute anticlockwise angle to the shear zone wall in a dextral shear zone (Fig. 4b), movement normal to the shear zone wall is compensated by normal-slip crenulations (NSC).

Asymmetrical buckle folds (reverse-slip crenulations) The most common crenulation-slip morphologies in South Achill and Achill Beg are asymmetrical buckle folds (Fig. 5a). The strike of the F3 axial planes makes an angle of approximately 27° with the strike of the S2 foliation in an anticlockwise direction (Fig. 6a). F3 fold axes plunge moderately to the NE (Figs 3,6a). The largest F3 folds have wavelengths of only a few metres, and typically the smaller F3 folds can be observed to 'root' in the S2 foliation surface. These features are typical of the reverse-slip crenulations of Dennis & Secor (1987). With progressive dextral

shear, the S3 foliation which initiates anticlockwise to S2 is progressively brought into parallelism (Fig. 5a). This is particularly evident where there is a large competence contrast across a bedding surface. Relatively rigid psammitic layers respond to D3 shear by folding with S3 usually oblique to S2. If the S3 foliation continues out into an adjacent graphitic pelite layer, then commonly the S3 foliation swings clockwise into parallelism with S2 and hence the weak graphitic pelite layers are accommodating the bulk of the displacement. The weak graphitic pelite layers also often display evidence of slip along the S2 foliation surfaces (Fig. 5b). On a vertical surface, the S3 foliation commonly rotates into the vertical parallel to S2, with a down-to-the-south shear sense.

Extensional crenulation cleavages (normalslip crenulations) Extensional crenulation cleavages (Platt & Vissers 1980) are relatively common within pelitic lithologies in the South Achill sequence. They make an angle of approximately 29° (Fig. 6b) with the S2 foliation in a clockwise direction, and consistently give a dextral sense of shear (Fig. 5c). Identical in style to the normal-slip crenulations of Dennis & Secor (1987), they are believed to be broadly contemporaneous with the F3 asymmetrical buckle folds based on the absence of apparent overprinting relationships. The orientation of the earlier S2 foliation controls the morphology of the later crenulations

342

D. M. CHEW ETAL.

Fig. 5. (a) Asymmetrical buckle folds (F3), interpreted as reverse slip crenulations (RSC). Later rotation of S3 cleavage is due to progressive dextral shear. Hammer 40 cm long. South Achill Dalradian [L69329495]. (b) Graphitic pelite illustrating decollement-like slip along the 82 foliation surface. Plane-polarized light, scale bar 1000 urn. South Achill Dalradian [L69039540]. (c) D3 dextral shear bands cutting the S2 foliation, interpreted as normal slip crenulations (NSC). Hammer 40 cm long. South Achill Dalradian [L69029544]. (d) Sinistral extensional crenulations affecting the composite S2/S3 foliation, interpreted as normal slip crenulations (NSC). Coin 2.2 cm in diameter. Central Ox Mountains Dalradian [G323026]. (e) Photograph of a polished rock slice used for 40Ar-39Ar in situ laserprobe dating of muscovite defining both the S2 and S3 foliations. Sample DC-79, scale bar 1000 um. South Achill Dalradian [L69089511].

(e.g. RSC vs. NSC). On a vertical surface, the extensional shear bands give a down-to-thesouth shear sense. Orientation of the D3 dextral shear zone From the angular relationships proposed by Dennis & Secor (1987), the shear zone wall is

expected to lie between the axial planes of the asymmetrical buckle folds (Fig. 6a) and the extensional shear bands (Fig. 6b). At localities which display both RSC and NSC fabrics, the dominant slip foliation (S2) modified by the RSC makes a very small clockwise angle with the S2 foliation affected by the NSC, similar to the geometry predicted by Fig. 4. However, on a


343

Fig. 6. (a) Stereographic plot of the orientation of RSC-related structures in South Achill (F3 fold hinges and poles to the S3 foliation) along with the mean orientation of the S2 foliation, (b) Rose diagram illustrating the mean orientation of NSC-related structures in South Achill (dextral shears measured on horizontal surfaces). The mean orientation of the S2 foliation is also illustrated, (c) Stereographic plot of the orientation of the preexisting foliation-slip surface in South Achill (poles to the S2 foliation), (d) Stereographic plot of the orientation of the pre-existing foliation-slip surface in the Central Ox Mountains (poles to the composite S2/S3 foliation), along with the mean orientation of NSC-related structures (sinistral extensional crenulation cleavages). regional scale this relationship is not apparent and hence the dominant slip foliation (S2) data for both the RSC and NSC sets are presented together (Fig. 6c). Combining data from both horizontal and vertical surfaces, the shear zone

would therefore be an approximately east-west trending, sub-vertical structure, similar to the geometry proposed by Sanderson et al. (1980). No L3 elongation lineations have been observed and hence although shear sense has been

344

D. M. CHEW ETAL.

determined reliably from the crenulation-slip morphologies on both horizontal and vertical surfaces, the exact shear direction remains uncertain.

Evidence for strike-slip motion along the FHCBL in the Central Ox Mountains

Glennawoo Slide and the Callow Shear Zone (Taylor 1969). In the Callow Loughs region (Fig. 7), the Lough Talt and Glennawoo Slides may be adequately represented as discrete shear zones, but the Callow Shear Zone is significantly wider (more than 750 m across strike) and is regarded as a substantial mylonite belt. Both the Glennawoo Slide and the Callow Shear Zone display abundant evidence of sinistral extensional crenulation cleavage development (Fig. 7).

The Central Ox Mountains inlier consists of a sequence of Dalradian metasediments (Long & Max 1977; Alsop & Jones 1991) intruded by a Caledonian granite, the Ox Mountains granodi- Extensional crenulation cleavages (normalorite (Fig. 2a, c). The intrusion age of the Ox slip crenulations) Mountain granodiorite has proved controversial in the past (e.g. Kennan 1997), and it is discussed In the Callow Loughs region, the main foliation in detail later. This granodiorite is intruded into is close to the vertical and trends NE (Figs 6d the core of a significant upright D3 antiform & 7). It is regarded as a composite S2/S3 fabric as which trends NE-SW, subparallel to the length S2 and S3 are usually coplanar (Jones 1989; Macof the inlier (Fig. 2c; Taylor 1969). High strain Dermot et al. 1996) and the S3 fabric can only be zones are well developed on the limbs of the conclusively identified when F3 folds are main D3 antiformal structure, are parallel to the present. F3 folds plunge gently to the NE and vertical, axial planar S3 fabric and kinematic verge towards the Ox Mountains granodiorite indicators such as rotated porphyroblasts and (Fig. 7). Extensional crenulation cleavages make extensional crenulation cleavages display abun- an angle of approximately 28° (Fig. 6d) with the dant evidence for sinistral shear (e.g. Hutton & main (composite S2/S3 foliation) in an anticlockDewey 1986; Hutton 1987; Jones 1989; McCaf- wise direction, and consistently give a sinistral frey 1992, 1994). These shear zones have been sense of shear (Fig. 5d). On a vertical surface, regarded as contemporaneous with the develop- the extensional shear bands give a down to the ment of the main E>3 antiform and the Central south shear sense. Ox Mountains has thus been regarded a trans- , The composite (S2/S3) main foliation is usually pressive sinistral shear zone during D3 (Hutton defined by muscovite, chlorite and equigranular & Dewey 1986; Hutton 1987; Jones 1989; quartz grains with interlobate grain boundaries. McCaffrey 1992,1994). These quartz grains are typically of the order of The Ox Mountains granodiorite itself also dis- 100-200 um in diameter and display undulose plays abundant evidence of sinistral strike-slip extinction. MP3 porphyroblasts of albite and deformation. The main solid-state foliation is almandine garnet overgrow the main foliation subvertical, strikes NE-SW and is accompanied and staurolite is locally developed (MacDermot by a stretching lineation which plunges gently to et al. 1996). Along the extensional crenulation the NE or SW (McCaffrey 1992, 1994). NNE cleavage planes, quartz has undergone signifitrending sinistral S-C fabrics are commonly well cant dynamic recrystallization. It is conspicudeveloped within the granodiorite and are sub- ously finer grained (10-20 um) than that parallel to the asymmetrical extensional crenu- defining the composite S2/S3 foliation, and has a lation cleavages developed in the country rock weak shape-preferred orientation (aspect ratios (Hutton & Dewey 1986; McCaffrey 1992,1994). of up to 3:1). Additionally, phyllosilicates within The Ox Mountains granodiorite has thus been the shear bands (chlorite and muscovite) are regarded as being emplaced synkinematically significantly finer grained than those defining the with respect to D3 sinistral transpressive defor- S2/S3 foliation. The extensional crenulation mation in the country rock (Hutton & Dewey cleavage planes are short and anastomosing, commonly rooting in the pre-existing composite 1986; Jones 1989; McCaffrey 1992,1994). S2/S3 foliation, and are very similar in morphology to the normal-slip crenulations (NSC) of Dennis & Secor (1987). High strain zones (tectonic slides) in the

Central Ox Mountains Three high strain zones are particularly well developed on the SE flank of the Central Ox Mountains inlier - the Lough Talt Slide, the

Isotopic dating of crenulation-slip surfaces Many deformed rocks contain evidence (e.g. multiple fabrics) for having experienced more


Fig. 7. Geological map of the Callow Loughs area in the Central Ox Mountains.

345

346

D.M. CHEW ETAL.

than one deformation event. The timing of growth of multiple fabrics can be constrained by dating fabric-forming minerals which crystallize below their closure temperatures (e.g. Cliff 1985). One of the most common fabric-forming minerals is muscovite, and Rb-Sr dating and 40 Ar-39Ar laserprobe dating of multiple generations of muscovite has been employed in many studies (e.g. Mtiller et al 1999). Both the Rb-Sr and Ar-Ar systems have shown that muscovite grown below its closure temperature during deformation (e.g. in mylonites) may record the age of neocrystallization (e.g. Dunlap 1997; Freeman et al. 1997, Miiller et al. 1999). Additionally, samples containing earlier generations of white mica (e.g. as porphyroclasts or older foliations) record the crystallization age of these early fabrics where they have not been rejuvenated by later deformation (West & Lux 1993; Freeman et al. 1997, Miiller et al. 1999). In this study, fabric-forming muscovite has been dated using the 40Ar-39Ar and Rb-Sr systems in order to constrain both the age of shearing and the age of the reactivated foliationslip surface within samples that have undergone shear-related deformation along the FHCBL. The age of the pre-existing foliation which is exploited by shearing along the FHCBL is also well constrained by previous geochronological studies based immediately outside of this shear zone (e.g. Flowerdew et al. 2000; Chew et al. 2003), where it is unaffected by later deformation. Thus the age of this foliation in undeformed samples and in samples where it has been reactivated due to subsequent shearing can be compared.

Pre-existing constraints on the age of the foliation-slip fabric in both inliers Recent geochronological studies in the Dalradian of NW Ireland (Flowerdew et al. 2000; Chew et al. 2003) has shown that main foliation development in both South Achill and the Central Ox Mountains inlier is Grampian (c. 470-460 Ma) in age in samples which are unaffected by later shearing along the FHCBL. Knowledge of the age of this main foliation in samples unaffected by later deformation is essential, as it enables us to assess whether the age of the reactivated foliation is partially reset when it is rejuvenated by later shearing along the FHCBL. Four Rb-Sr S2 muscovite ages from the Dairadian of South Achill range from 460-458 Ma (± 7 Ma), and four 40Ar-39Ar S2 muscovite step-

heating plateaux from the same samples range from 463-457 Ma (± 4 Ma) (Chew et al. 2003). Two Rb-Sr S2 muscovite ages of 460 ± 7 Ma and 461 ± 7 Ma have been obtained from the low greenschist facies Clew Bay Complex and probably record crystallization (Chew et al. 2003). As this outboard terrane (Figs 2 & 3) is in structural continuity with the Dalradian Supergroup (Chew 2003), it is thought likely that the c. 460 Ma Rb-Sr S2 muscovite ages for the South Achill Dalradian also record crystallization during the Grampian orogeny, particularly as the peak metamorphic temperature in South Achill is unlikely to have exceeded 450 °C (Chew et al. 2003). The marked similarity with the S2 muscovite 40Ar-39Ar step heating data may imply that the Ar-Ar system is recording either rapid cooling or crystallization, despite growing c. 50 to 100 °C above its closure temperature (c. 350-400 °C; Wijbrans & McDougall 1988). The possibility of extraneous Ar contamination is discussed in detail later. One 40Ar-39Ar hornblende age of 467 ± 3 Ma and one Rb-Sr muscovite age of 472 ± 8 Ma have been obtained from the composite S2/S3 foliation in the Dalradian of Central Ox Mountains inlier (Flowerdew et al. 2000). Younger 40Ar_39Ar and Rb-Sr ages (429-410 Ma) have been strongly influenced by the intrusion of the Ox Mountains granodiorite. Similar Grampian (c. 470-460) ages have also been obtained from Dalradian rocks in the NE Ox Mountains (Fig. 1). Here Dalradian rocks are interleaved during D3 with the Slishwood Division, a unit of psammitic gneisses which has experienced late Precambrian granulite-facies metamorphism (Sanders et al. 1987). Three 40Ar-39Ar hornblende ages and six Rb-Sr muscovite ages defining the S3 foliation within the interleaved Dalradian range from 470-446 Ma. The older S3 mineral ages are indistinguishable from the older ages (472 ± 7 Ma, 467 ± 3 Ma) obtained from the composite S2/S3 fabric in the Central Ox Mountains inlier (Flowerdew et al. 2000).

Sampling In this study we present isotopic data from two Dalradian samples from two localities. DC-79 is a semipelite from South Achill (Fig. 3) in which muscovite defining both the local S2 and S3 foliations was dated in situ by the 40Ar-39Ar laserprobe spot fusion method. DC-03/02-1 is a semipelite from the Central Ox mountains inlier (Fig. 7) from which a bulk mineral separate of muscovite defining the main (composite S2/S3) foliation was analysed by the Rb-Sr method. The whole rock was also analysed.


347

Table 1. 40Ar-39Ar spot fusion data Spot

Laser power

36

Ar(a)

37

Ar(Ca)

38

Ar(Cl)

39

Sample DC-79, S2 and S3 ms. Dalradian, South Achill 2(S 2 ) Fusion 0.0002 0.0000 0.0001 4(S 2 ) Fusion 0.0002 0.0266 0.0000 6(S2) Fusion 0.0000 0.0299 0.0002 10 (S2) Fusion 0.0010 0.0208 0.0000 11 (S2) Fusion 0.0004 0.0308 0.0004 1(S3) Fusion 0.0001 0.0000 0.0000 3(S 3 ) Fusion 0.0000 0.0000 0.0001 5(S 3 ) Fusion 0.0001 0.0345 0.0001 7(S3) Fusion 0.0013 0.0511 0.0002 8(S3) Fusion 0.0002 0.0371 0.0000 9(S 3 ) Fusion 0.0003 0.0178 0.0000 Weighted average* (S2 spots): 451 ± 2 Ma (2a) Weighted average (S3 spots): 448 ± 3 Ma (2a) J value: 0.003987 ± 0.5% (la)

Ar(K)

40

Ar(r) Age (Ma) ±2o

(L69089511). 0.1262 9.0407 0.0320 2.2206 0.0209 1.5135 0.5635 40.1131 0.1893 13.5803 0.1492 10.5535 0.0972 6.8772 0.1325 9.3329 0.0683 4.7528 0.0891 6.2796 0.3099 22.0392

453.4 441.3 457.8 450.7 453.8 448.2 448.2 446.6 441.7 446.9 450.4

3.4 10.1 14.5 1.3 3.5 2.8 3.5 3.6 5.5 4.8 1.9

40

Ar(r) (%)

39

99.2 97.3 99.4 99.3 99.1 99.7 99.8 99.8 92.4 99.0 99.6

Ar(k) (%) 7.1 1.8 1.2 31.7 10.7 8.4 5.5 7.5 3.8 5.0 17.4

*Weighted averages calculated using ISOPLOT (Ludwig 1999) and use the 2a error associated with each analysis.

Analytical methods 40

39

Ar- Ar spot fusion analyses were carried out using the VULKAAN argon laserprobe (Wij brans et al 1995) at the Vrije Universiteit in Amsterdam. Samples were irradiated at the HPPIF facility in the high flux research reactor at Petten, Netherlands. Polished slices were interspersed between the Al tablets containing the flux monitor DRA-1 sanidine (24.99 ± 0.07 Ma; Wijbrans et al. 1995) prior to irradiation. Four flux monitors were used to construct a J-curve with a 0.5% error (lo). Samples were analysed within six months of irradiation to minimize the interference effects produced by radioactive decay after irradiation. The analytical procedure is described in detail by Wijbrans et al. (1995) and is outlined below. Samples were heated using a continuous 18 W argon ion laser (454.5-514.5 nm wavelength). For spot fusion experiments, several short laser pulses (0.1 s) excavated a pit approximately 30 microns in diameter, surrounded by a crater of melt material. The Ar released was cleaned with Fe-V-Zr getters (250 °C), prior to analysis on a MAP-215/50 mass spectrometer. Data reduction was carried out using in-house software, ArArCALC V20. Blank intensities were measured every 3-5 sample runs and mass fractionation was corrected for by regular measurement of shots of clean air argon. For Rb-Sr analyses, standard ion exchange methods were used for chemical separation of elements. Samples were loaded on tantalum filaments and were analysed on a semi-automated

single collector VG Micromass 30 mass spectrometer at the Department of Geology, University College Dublin. During the course of analysis, NBS SRM 987 gave 87Sr/86Sr ratios of 0.71027 ± 5 (n = 8, 2o) and NBS SRM 607 yielded 87Rb/86Sr ratios of 8.005 ± 13 (n = 7,2o). Sr blanks averaged 1.5 ng and are not significant. 2o analytical uncertainties of 1.5% for 87Rb/86Sr and tabulated values (Table 2) for 87Sr/86Sr ratios were used in age calculations which employed a value of 0.0142 Ga"1 for the 87Rb decay constant (Steiger and Jager 1977).

Constraining dextral shear in NW Mayo (South Achill) New S3 mica growth is not commonly associated with the D3 deformation event in South Achill, but locally S3 muscovite is found overgrowing a crenulated S2 fabric in the hinges of asymmetrical buckle folds (reverse-slip crenulations associated with D3 dextral shear). One polished slice from the hinge of an F3 fold was selected for in situ 40Ar-39Ar laserprobe spot fusion analyses (Table 1). Sample DC-79 is semipelitic, with an S2 foliation defined by muscovite and minor chlorite and epidote. Individual muscovite grains are typically around 500 um long and 50 um wide, but both the S2 and S3 fabrics are composed of seams of white mica which are typically several grains in width. Typically S3 seams are approximately 150 um across (Fig. 5e), whereas S2 lithons are larger and may be up to 500 um wide (Fig. 5e). Chlorite grains are

348

D. M. CHEW£r,4L.

typically similar in size to the muscovite grains (c. 500 urn long and 50 urn wide) and the largest epidote needles observed are 200 urn long. The S2 foliation is overgrown by MP2 plagioclase which is augened by the seams of S3 muscovite. The calcic component in some of the Ar analyses (Table 1) is probably derived from the epidote, as the MP2 plagioclase is essentially pure albite. Muscovite defining the S2 foliation is typically more celadonite-rich and paragonitepoor than the later (S3) fabric (Chew et al 2003). In situ 40Ar-39Ar laserprobe dating of the S3 muscovite seams yields a weighted mean age of 448 ± 3 Ma, whereas the older, crenulated S2 muscovite seams yield a weighted mean age of 451 ± 2 Ma (Table 1). The 40Ar-39Ar system is only reliably recording the youngest deformation fabrics present, as undeformed S2 muscovite from South Achill yields consistent c. 460 Ma Rb-Sr and 40Ar-39Ar ages (Chew et al. 2003). The 448 ± 3 Ma age for muscovite within the S3 crenulation-slip fabric is interpreted as a crystallization age based on the low-greenschist facies assemblages observed in the S3 crenulation seams as detailed above. The possibility of extraneous argon cannot be ruled out, particularly in a high-pressure, low temperature terrane such as South Achill. Excess argon (argon with 40Ar/36Ar ratios which differ from the modern atmospheric ratio of 295.5) has been documented in white mica from several high-pressure, low-temperature terranes (e.g. Arnaud & Kelley 1995; Sherlock & Kelley 2002). The presence of excess argon may be evaluated by using an inverse isochron correlation diagram (a plot of 36Ar/40Ar versus 39 Ar/40Ar) as the incorporation of excess argon will result in the intercept of the isochron on the ordinate axis deviating from the modern atmospheric 40Ar/36Ar ratio of 295.5. However, K-rich phases (such as white mica) can produce large quantities of radiogenic 40Ar, and hence the data often cluster close to the 39Ar/40Ar axis. The presence of excess argon is therefore difficult to assess, as the intercept with the 36Ar/40Ar axis is poorly constrained. This is the case with the South Achill Ar data. The presence of excess argon has also been documented in samples which yield intercepts within error of the modern 40Ar/36Ar atmospheric ratio on an inverse isochron correlation diagram (Sherlock & Arnaud 1999). Studies that have documented the presence of excess argon in white mica in high-temperature, low-pressure terranes instead are based on 40 Ar-39Ar phengite ages that are significantly older than either the corresponding Rb-Sr phengite ages or other geochronometers with

significantly higher closure temperatures (e.g. Arnaud & Kelley 1995; Sherlock & Arnaud 1999). This disparity is not observed in South Achill as both the Rb-Sr and 40Ar-39Ar ages for undeformed S2 muscovite ages cluster at c. 460 Ma (Chew et al. 2003) and are mutually within error. The S3 muscovite ages are temporally distinct, and are clearly post-Grampian (c. 470-460 Ma) in age.

Constraining sinistral shear in the Central Ox Mountains The Central Ox Mountains displays extensional crenulation cleavage development superimposed on what is believed to be a pre-existing Grampian (S2/S3) foliation. However, whereas dextral shear in South Achill was constrained by dating both the pre-existing foliation-slip surface and the later crenulation-slip surfaces within the same sample, this strategy has proved impossible in the Central Ox Mountains. Dalradian metasediments in the Central Ox Mountains display only limited growth of extremely fine-grained muscovite on extensional crenulation cleavage surfaces, which is not sufficient for isotopic dating. However, pegmatites and granite sheets associated with the Ox Mountains granodiorite are intruded into the Dalradian metasediments, and in common with the Ox Mountains granodiorite, the pegmatites and granite sheets were emplaced synkinematically with respect to sinistral deformation in the country rocks (McCaffrey 1992, 1994). Sinistral shearing in the Central Ox Mountains can thus be constrained by dating pegmatite crystallization. The age of the foliation surface that is affected by sinistral extensional crenulation cleavages in the Central Ox Mountains has until now remained uncertain, as previous geochronological studies in the inlier (e.g. Flowerdew et al. 2000) were undertaken on foliated samples that were not affected by later deformation. Whereas the reactivated foliation surface is believed to be the regional S2/S3 composite foliation based on detailed field mapping (Fig. 7), it may have developed contemporaneously with sinistral shear development in the Central Ox Mountains, as the timing relationship between mylonitic foliation development and cross-cutting shear bands can be often difficult to establish (e.g. Lister & Snoke 1984). The age of the pre-existing foliation has been constrained by a Rb-Sr muscovite-whole rock age from a semi-pelitic schist sample displaying a pervasive sinistral extensional crenulation cleavage (Fig. 5d). This age of 448 ± 9 Ma (Table 2) is


349

Table 2. Rb-Sr geochronology Sample

Locality and Irish National Grid Ref.

DC-03/02-131

Textural relationship

Mineral

Rb(ppm) Sr(ppm)

Lismoran Main foliation muscovite 338.77 126.60 (G323026) (composite S2/S3) whole rock 77.8571 141.07

87

Rb/86Sr

7.79 1.60

87

Sr/86Sr ± 2a

0.776157 ± 56 0.736629 ± 58

87

Sr/86Sr(i) Age ± 2a (Ma) 0.72641

448 ± 9

Table 3. Rb-Sr geochronology of Ox Mountains pegmatites from Flowerdew et al. (2000) Sample 37 38 39 40 37 40 41

Irish National Grid Ref.

Sample description

Minerals dated

Age ± 2o (Ma)

G270003 M167963 G198956 G242069 G270003 G242069 G198956

Dalradian-hosted pegmatite Dalradian-hosted pegmatite Dalradian-hosted pegmatite Ox Mountains granodiorite Dalradian-hosted pegmatite Ox Mountains granodiorite Dalradian-hosted pegmatite

coarse igneous muscovite-K feldspar coarse igneous muscovite-plagioclase coarse igneous muscovite-K feldspar coarse igneous muscovite-K feldspar recrystallized muscovite-K feldspar recrystallized muscovite-K feldspar recrystallized muscovite-plagioclase

392 ±6 400 ±6 402 ±6 400 ±6 381 ±6 384 ±6 384 ±5

slightly younger than the c. 470-460 Ma age estimates for the Grampian 83 foliation in the Central and NE Ox Mountains inliers (Flowerdew et al. 2000). However, it is broadly coincident with most of the mineral cooling age data from the Central and NE Ox Mountains inliers which cluster at or around 460^50 Ma (Flowerdew et al. 2000), and the reactivated foliation surface is thus thought to represent the regional composite Grampian S2/S3 foliation. The pre-existing foliation surface is markedly older than the age constraints for sinistral extensional crenulation cleavage development in the Central Ox Mountains inlier. Sinistral shearing is constrained by four Rb-Sr muscovite-feldspar ages from pegmatites which range from 402-392 Ma (Table 3) which are interpreted as recording igneous crystallization (Flowerdew et al. 2000). Late sinistral shearing has recrystallized coarse magmatic muscovite within both the Ox Mountains pegmatite suite and the Ox Mountains granodiorite, yielding three Rb-Sr muscovite-feldspar ages of between 385 and 381 Ma (Table 3; Flowerdew et al. 2000).

Discussion on the intrusion age of the Ox Mountains granodiorite Sinistral shearing along the FHCBL in the Central Ox Mountains is effectively constrained by the age of intrusion of the Ox Mountains granodiorite and its presumed coeval pegmatite suite. However, in contrast to the c. 400 Ma age

obtained from the Ox Mountains pegmatite suite (see above), the Ox Mountains granodiorite has yielded c. 480 Ma Rb-Sr whole rock isochrons (Pankhurst et al. 1976; Max et al. 1976). The old Rb-Sr ages suggest that granite emplacement and therefore sinistral strike-slip deformation occurred either before or very early in the Grampian orogeny. However several lines of evidence mitigate against a c. 480 Ma intrusion age. Most of the Ox Mountains granodiorite Rb-Sr whole rock data are characterized by low 87Rb/86Sr values, typically less than 2. Rb-Sr whole rock isochrons characterized by low 87Rb/86Sr values have also been obtained from other Caledonian granites, and these too yield c. 480 Ma intrusion ages even though the intrusion is demonstrably younger (c. 400 Ma) by independent evidence (Kennan 1997). Additionally, unpublished U-Pb multigrain zircon data suggest a c. 415 Ma intrusion age for the Ox Mountains granodiorite (MacDermot et al., 1996). This is consistent with the youngest mineral cooling ages (c. 410 Ma) obtained from the Central Ox Mountains inlier close to the Ox Mountains granodiorite which are likely to be due to thermal resetting, and a c. 400 Ma emplacement age for the pegmatite suite (Flowerdew et al. 2000); a c. 400 Ma intrusion age for the Ox Mountains granodiorite is more likely.

Earlier movements along the FHCBL This study documents two examples of postGrampian strike-slip movement along the

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FHCBL. However, there is evidence for earlier stages of movement along the FHCBL during the Grampian orogeny, illustrating further how this important crustal-scale shear zone has had a long and complicated history of movement. In Tyrone, the D3 Omagh Thrust has translated inverted Dalradian rocks towards the ESE (Fig. 1) over Arenig-Llanvirn shales of the Tyrone volcanics (Alsop & Hutton 1993). In southern Donegal and the NE Ox Mountains inlier, Dalradian rocks were thrust to the SE over granulite-facies basement of the Slishwood Division (Fig. 1) along the Lough Derg Slide (Alsop 1991) and the North Ox Mountains Slide (Flowerdew 1998/9) respectively. Tectonic juxtaposition (D3) of the Dalradian and Slishwood Division is likely to have occurred between 470 and 460 Ma based on 40Ar-39Ar, Rb-Sr and Sm-Nd mineral ages (Flowerdew et al. 2000). Thus in Donegal, the NE Ox Mountains and Tyrone, the earliest constrained phase of movement (c. 470-460 Ma) along the Fair Head-Clew Bay Line involves the translation of the main Dalradian nappes over outboard terranes to the SE.

Conclusions The Fair Head-Clew Bay Line has been shown to have been reactivated several times. It was originally active as a ductile thrust during the (c. 470-460 Ma) Grampian orogeny, where Dalradian nappes adjacent to the Fair Head-Clew Bay Line were thrust over outboard terranes to the SE. The Dalradian metasediments have undergone two separate phases of post-Grampian strike-slip movement adjacent to the Fair Head-Clew Bay Line. These two phases of movement have produced reverse-slip and normal-slip crenulations which modified the earlier Grampian nappe fabrics, and tilted the initially recumbent Grampian nappes into a vertical orientation (Fig. 2b, c). The development of the reverse-slip and normal-slip crenulations produced by these two discrete phases of strike-slip movement has been constrained by isotopic dating. Dextral displacement along the Fair Head-Clew Bay Line in the NW Mayo inlier is constrained to c. 448 Ma, whereas sinistral displacement along the Fair Head-Clew Bay Line in the Ox Mountains inlier is constrained to c. 400 Ma based on previously published pegmatite intrusion ages. Major crustal-scale shear zones may therefore have a long and complicated history of movement, in which pre-existing planar anisotropies (e.g. foliations) act as slip surfaces during later

non-coaxial deformation. Careful analysis of the resulting crenulation morphologies combined with isotopic dating yields a more complete understanding of the reactivation history of major crustal-scale shear zones. D.M.C. gratefully acknowledges a Forbairt Basic Research Grant and a UCD Research Doctoral Scholarship. Barry Long is thanked for many fruitful discussions about Dalradian geology, and the Central Ox Mountains in particular. Sarah Sherlock and Grahame Oliver are thanked for their careful and constructive reviews, which significantly improved this paper.

References ALSOP, G.I. 1991. Gravitational collapse and extension along a mid-crustal detachment: the Lough Derg Slide, northwest Ireland. Geological Magazine, 128, 345-354. ALSOP, G.I. & HUTTON, D.H.W. 1993. Major southeastdirected Caledonian thrusting and folding in the Dalradian rocks of mid-Ulster: implications for Caledonian tectonics and mid-crustal shear zones. Geological Magazine, 130, 233-244. ALSOP, G.I. & JONES, C.S. 1991. A review and correlation of Dalradian stratigraphy in the Ox Mountains and southern Donegal, Ireland. Irish Journal of Earth Sciences, 11, 99-112. ARNAUD, N.O. & KELLEY, S.P. 1995. Evidence for excess argon during high pressure metamorphism in the Dora Maira Massif (western Alps, Italy), using an ultraviolet laser ablation microprobe 40 Ar-39Ar technique. Contributions to Mineralogy and Petrology, 121,1-11. CHEW, D.M. 2003. Structural and stratigraphic relationships across the continuation of the Highland Boundary Fault in western Ireland. Geological Magazine, 140, 73-85. CHEW, D.M., DALY, J.S., PAGE, L.M. & KENNEDY, MJ. 2003. Grampian orogenesis and the development of blueschist-facies metamorphism in western Ireland. Journal of the Geological Society, London, 160, 911-924. CLIFF, R.A. 1985. Isotopic dating in metamorphic belts. Journal of the Geological Society, London, 142, 97-110. DALY, IS. 1996. Pre-Caledonian history of the Annagh Gneiss Complex, north-western Ireland, and correlation with Laurentia-Baltica. Irish Journal of Earth Sciences, 15, 5-18. DENNIS, AJ. & SECOR, D.T. 1987. A model for the development of crenulations in shear zones with applications from the southern Appalachian Piedmont. Journal of Structural Geology, 9, 809-817. DENNIS, AJ. & SECOR, D.T. 1990. On resolving shear direction in foliated rocks deformed by simple shear. Geological Society of America Bulletin, 102,1257-1267. DEWEY, J.F. & SHACKLETON, R.M. 1984. A model for the evolution of the Grampian tract in the early

REACTIVATION OF A CALEDONIAN SHEAR ZONE Caledonides and Appalachians. Nature, 312, 115-121. DUNLAP, W.J. 1997. Neocrystallization or cooling? 40 Ar/39Ar ages of white micas from low-grade mylonites. Chemical Geology, 143,181-203. FLOWERDEW, MJ. 1998/9. Tonalite bodies and basement-cover relationships in the North-eastern Ox Mountains Inlier, north-western Ireland. Irish Journal of Earth Sciences, 17, 71-82. FLOWERDEW, M.J., DALY, IS., GUISE, P.O. & REX, D.C. 2000. Isotopic dating of overthrusting, collapse and related granitoid intrusion in the Grampian orogenic belt, northwestern Ireland. Geological Magazine, 137, 419^35. FREEMAN, S.R., INGER, S., BUTLER, R.W.H. & CLIFF, R.A. 1997. Dating deformation using Rb-Sr in white mica: greenschist facies deformation ages from the Entrelor shear zone, Italian Alps. Tectonics, 16, 57-76. FRIEDRICH, A.M., BOWRING, S.A., MARTIN, M.W. & HODGES, K.V. 19990. Short-lived continental magmatic arc at Connemara, western Ireland Caledonides: implications for the age of the Grampian orogeny. Geology, 27, 27-30. FRIEDRICH, A.M., HODGES, K.V., BOWRING, S.A. & MARTIN, M.W. I999b. Geochronological constraints on the magmatic, metamorphic and thermal evolution of the Connemara Caledonides, western Ireland. Journal of the Geological Society, London, 156,1217-1230. HARRIS, D.H.M. 1993. The Caledonian evolution of the Laurentian margin in western Ireland. Journal of the Geological Society, London, 150, 669-672. HARRIS, D.H.M. 1995. Caledonian transpressional terrane accretion along the Laurentian margin in Co. Mayo, Ireland. Journal of the Geological Society, London, 152, 797-806. HUTTON, D.H.W. 1987. Strike-slip terranes and a model for the evolution of the British and Irish Caledonides. Geological Magazine, 124, 405-425. HUTTON, D.H.W. & DEWEY, J.F. 1986. Palaeozoic terrane accretion in the Irish Caledonides. Tectonics, 5,1115-1124. JONES, C.S. 1989. The structure and kinematics of the Ox Mountains, western Ireland; a mid-crustal transcurrent shear-zone. Unpublished Ph.D. thesis, University of Durham. KENNAN, PS. 1997. Granite: a singular rock. In: Occasional papers in Irish science and technology, 15. Royal Dublin Society, 16. KENNEDY, M.J. 1980. Serpentinite-bearing melange in the Dalradian of County Mayo and its significance in the development of the Dalradian basin. Journal of Earth Sciences of the Royal Dublin Society, 3,117-126. LISTER, G.S. & SNOKE, A.W 1984. S-C mylonites. Journal of Structural Geology, 6, 617-638. LONG, C.B. & MAX, M.D. 1977. Metamorphic rocks in the SW Ox Mountains Inlier, Ireland; their structural compartmentation and place in the Caledonian orogen. Journal of the Geological Society, London, 133, 413-432. LUDWIG, K.R. 1999. Users Manual for Isoplot/Ex, Version 2.10: a geochronological toolkit for

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Microsoft Excel. Berkeley Geochronology Center Special Publication, la, 49 pp. MACDERMOT, C.V., LONG, C.B. & HARNEY, S. 1996. Geology of Sligo-Leitrim: A geological description of Sligo, Leitrim, and adjoining parts of Cavan, Fermanagh, Mayo and Roscommon, to accompany the Bedrock Geology 1:100 000 Scale Map Series, Sheet 7, Sligo-Leitrim. Geological Survey of Ireland. MAX, M.D. & RIDDIHOUGH, R.P 1975. Continuation of the Highland Boundary fault in Ireland. Geology, 3, 206-210. MAX, M.D., LONG, C.B. & SONET, J. 1976. The geological age and setting of the Ox Mountains Granodiorite. Bulletin of the Geologial Survey of Ireland, 2, 27-5. MCCAFFREY, KJ.W. 1992. Igneous emplacement in a transpressive shear zone: Ox Mountains igneous complex. Journal of the Geological Society, London, 149, 221-235. MCCAFFREY, KJ.W. 1994. Magmatic and solid state deformation partitioning in the Ox Mountains granodiorite. Geological Magazine, 131, 639-652. MULLER, W, DALLMEYER, R.D., NEUBAUER, F. & THONI, M. 1999. Deformation-induced resetting of Rb/Sr and 40Ar/39Ar mineral systems in a lowgrade, polymetamorphic terrane (Eastern Alps, Austria). Journal of the Geological Society, London, 156, 261-278. PANKHURST, R.J., ANDREWS, J.R., PHILLIPS, W.E.A., SANDERS, I.S. & TAYLOR, WE.G 1976. Age and structural setting of the Slieve Gamph Igneous Complex, Co. Mayo, Ireland. Journal of the Geological Society, London, 132, 327-36. PLATT, J.P. & VISSERS, R.L. 1980. Extensional structures in anisotropic rocks. Journal of Structural Geology, 2, 397-410. RAMSAY, J.G. & GRAHAM, R.H. 1970. Strain variation in shear belts. Canadian Journal of Earth Sciences, 7,786-813. RYAN, P.D., SOPER, NX, SNYDER, D.B., ENGLAND, R.W. & HUTTON, D.H.W 1995. The Antrim-Galway Line: a resolution of the Highland Border fault enigma of the Caledonides of Britain and Ireland. Geological Magazine, 132,171-184. SANDERS, I.S., DALY, J.S. & DAVIES, G.R. 1987. Late Proterozoic high-pressure granulite facies metamorphism in the north-east Ox inlier, north-west Ireland. Journal of Metamorphic Geology, 5, 69-85. SANDERSON, D.J., ANDREWS, J.R., PHILLIPS, WE.A. & HUTTON, D.H.W 1980. Deformation studies in the Irish Caledonides. Journal of the Geological Society, London, 137, 289-302. SHERLOCK, S.C. & ARNAUD, NO. 1999. Flat plateau and impossible isochrons: Apparent 40Ar-39Ar geochronology in a high-pressure terrain. Geochimica et Cosmochimica Acta, 63, 2835-2838. SHERLOCK, S.C. & KELLEY, S.P. 2002. Excess argon evolution in HP-LT rocks: a UVLAMP study of phengite and K-free minerals, NW Turkey. Chemical Geology, 182, 619-636. STEIGER, R.H. & JAGER, E. 1977. Subcommission on geochronology: convention on the use of decay

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constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. TAYLOR, W.E.G. 1969. The structural geology of the Dalradian rocks of Slieve Gamph, Cos. Mayo and Sligo, western Ireland. Geologische Rundschau, 57, 564-588. WEST, D.RW. & Lux, D.R. 1993. Dating mylonitc deformation by the 40Ar-39Ar method: an example from the Norumbega Fault Zone, Maine. Earth and Planetary Science Letters, 120,221-237. WIJBRANS, J.R. & MCDOUGALL, 1.1988. Metamorphic evolution of the Attic Cycladic Metamorphic Belt on Naxos (Cyclades, Greece) utilizing 40Ar/39Ar age spectrum measurements. Journal of Metamorphic Geology, 6, 571-594.

WIJBRANS, J.R., PRINGLE, M.S., KOPPERS, A.A.P. & SCHEEVERS, R. 1995. Argon geochronology of small samples using the Vulkaan argon laserprobe. Proceedings of the Royal Netherlands Academy of Arts and Sciences, 98,185-219. WILLIAMS, D.M., HARKIN, J., ARMSTRONG, H.A. & HIGGS, K.T. 1994. A late Caledonian melange in Ireland: implications for tectonic models. Journal of the Geological Society, London, 151, 307-314. WINCHESTER, J.A. 1992. Comment on 'Exotic metamorphic terranes in the Caledonides: Tectonic history of the Dalradian block, Scotland'. Geology, 20, 764.

Brittle-ductile shear zone evolution and fault initiation in limestones, Monte Cugnone (Lucania), southern Apennines, Italy S. MAZZOLI1, C. INVERNIZZI2, L. MARCHEGIANI2, L. MATTIONI2 & G. CELLO2 l Facoltd di Scienze Ambientali, Universitd di Urbino, Campus Scientifico Sogesta, 61029 Urbino (PU), Italy (e-mail:[email protected]) 2 Dipartimento di Scienze della Terra, Universitd di Camerino, Via Gentile III da Varano, 62032 Camerino (MC), Italy Abstract: The processes of brittle-ductile shear zone evolution and fault initiation by the coalescence of en echelon arrays of tensile cracks are quantitatively analysed in terms of displacement and temperature conditions at which they took place in very low-grade, well bedded micritic limestones from the southern Apennines, Italy. Three different types of structures are distinguished: (i) conjugate arrays of en echelon, calcite-filled tension gashes, showing extensional shear offsets; (ii) en echelon vein arrays showing incipient development of discontinuous shear-parallel fractures cutting through the tension gashes; and (iii) faulted vein arrays, in which vein array breaching by a continuous, discrete normal fault has occurred. Fluid inclusion microthermometry from vein calcite sampled from the different sets of structures (i) to (iii) above indicates that environmental conditions remained roughly constant during the different stages of vein array evolution and fault development, with average homogenization temperatures from primary fluid inclusions being in the range 130-140 °C. Our results show how displacement accumulation and shear strain essentially control vein array evolution by rotation of en echelon tension gashes, fracture linkage and, eventually, fault nucleation, at approximately constant temperature.

En echelon vein arrays, often organized in more or less well developed conjugate sets, have been intensely investigated by structural geologists during recent decades. The kinematic interpretation of these features has also been controversial (e.g. Roering 1968; Lajtai 1969; Ramsay & Graham 1970; Beach 1975; Durney & Ramsay 1983; Pollard et al 1982; Rothery 1988; Smith 1995,1996,1997), and an important step in their interpretation followed the application of the principles of continuum mechanics. The resulting model by Olson & Pollard (1991) suggested that selective vein growth forming en echelon arrays initially might be controlled by the mechanical interaction of neighbouring fractures. Once formed, macroscopic en echelon vein arrays might act as zones of weakness, localizing the later development of brittle-ductile shear zones (as defined by Ramsay & Huber 1987). The latter structures, in turn, might evolve to faults through fracturing between en echelon veins, with fault length increasing as more en echelon fractures link up (Roering 1968; Segall & Pollard 1983; Martel et al 1988). In recent years, a relevant contribution to the understanding of strike-slip fault nucleation by processes of this type came from the work of Peacock & Sanderson (1995), Willemse et al

(1997) and Kelly et al (1998). In all cases, the field examples were characterized by the occurrence of en echelon vein arrays in conjugate fault zones. The observed evolution along these zones from vein arrays to faults allowed the authors to implement a model of linkage in which pull-apart mechanisms play a primary role. This model can also explain the occurrence of vein arrays showing different structural styles in the same outcrop (Kelly et al 1998), which is observed within the area of the present study. Fault initiation by the coalescence of en echelon arrays of Mode I (tensile) cracks (Lawn & Wilshaw 1975) has been suggested by several authors (e.g. Knipe & White 1979; Pollard et al 1982; Etchecopar et al 1986; Cox & Scholz 19880, b\ Cowie & Scholz 1992). According to Scholz (1990), this process is likely to be important in the early stages of fault development. In this work we report arrays of en echelon tension gashes, occurring along shear zones of extensional type, which appear to have evolved locally into discrete normal faults. The studied structures are hosted in well-bedded, finegrained limestones exposed in a quarry in the southern Apennines in the Lucania province of Italy (Fig. 1). This outcrop displays several useful characteristics (Mazzoli & Di Bucci 2003;

From: ALSOP, G. L, HOLDSWORTH, R. E., MCCAFFREY, K. J. W. & HAND, M. (eds) 2004. Flow Processes in Faults and Shear Zones. Geological Society, London, Special Publications, 224, 353-373. 0305-8719/$15.00 © The Geological Society of London 2004.

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Fig. 1. Location (a), geological setting (b)» and (c) cross-section (no vertical exaggeration) of the study area (modified after Mazzoli et al 2001).

SHEAR ZONE EVOLUTION AND FAULTING

Fig. 2): (i) numerous conjugate sets of vein arrays are well exposed; (ii) some vein arrays show incipient shear fracture development across the tensile cracks; (iii) some are faulted; (iv) in most instances the displacement can be accurately measured; and (v) the displacement associated with vein arrays and faults is quite low (never exceeding 1 m, and mostly below 20 cm). Arrays of fractures, minor faults and veins are commonly observed at the lateral and vertical tips of faults, where they are generally interpreted to record fault propagation rather than initiation/nucleation (e.g. McGrath & Davison 1995; Childs et al 19960; Marchal et al 2003). In these cases, vein arrays and shear zones are interpreted to represent 'process zones' or inelastic strains at the tip of faults that are out of the plane of observation (e.g. Childs et al. 19965). In the latter instance, differently developed structures could represent serial sections through similarly structured fault zones that are dominated by brittle deformation and strain localization (fault slip) over their central portions and brittle-ductile strain towards and around the fault tips. However, in our study area, not only brittle-ductile shear zones dominate the population of structures in terms of their frequency, but also the relatively few discrete brittle faults exposed in the quarry systematically overprint en echelon vein arrays (Fig. 2). Neither a single fault devoid of preexisting en echelon veins, nor vein arrays occurring solely at the tips of discrete fault planes have been observed. These features strongly suggest that the analysed shear zones are indeed precursors to brittle faulting, as opposed to being fault zone-tip related. In summary, evidence of the early stages of normal fault nucleation by the process of coalescence of en echelon vein arrays (e.g. Scholz 1990) is exposed in the studied outcrop. The thermal conditions governing these stages are also constrained by means of fluid inclusion microthermometry, permitting a comprehensive analysis of the early faulting process to be made.

Geological setting The work was carried out in Upper Triassic micritic limestones with chert levels (Calcari con Selce Fm) from the Mesozoic Lagonegro Basin succession (e.g. Scandone 1972) of the southern Apennines fold and thrust belt of peninsular Italy. The southern Apennines include mostly tectonic units derived from the telescoped Apulian (or Adriatic) continental palaeomargin (e.g.

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Cello & Mazzoli 1999 and references therein) together with the remnants of a Cretaceous to Palaeogene accretionary complex (ophiolitebearing Liguride units). Structurally below the internal Liguride units, the thrust belt consists of allochthonous units derived from the deformation of both carbonate platform and pelagic basin successions (Apenninic Platform and Lagonegro Basin, respectively), of passive margin origin and Triassic to Palaeogene age, which are stratigraphically overlain by Neogene foredeep and wedge-top basin deposits (Miocene and Pliocene successions in Fig. la; Carbone et al 1991). These allochthonous units are completely detached from their original substratum and transported onto the foreland sequence of the Apulian Platform. Analysis of synorogenic deposits indicates that thrust accretion of the units derived from the deformation of the Apenninic Platform and Lagonegro Basin successions occurred mainly in Miocene times, while deeper thrusting involving the hinterland part of the Apulian Platform carbonates occurred mainly in Late Pliocene to Early Pleistocene times (e.g. Cello & Mazzoli 1999 and references therein). The late stages of thrusting were partially contemporaneous with kinematically compatible strike-slip faulting along roughly NNE-SSW trending, right-lateral, and WNW-ESE trending, left-lateral, structures (e.g. Monaco et al 1998). Active compression within the southern Apennines appears to have ceased by Middle Pleistocene times (e.g. Hyppoly te et al 1994). Subsequently, the geometry of the orogen has been modified mainly by NE-SW orientated extension (e.g. Cinque et al. 1993). The limestones analysed in this study are exposed in a quarry at Monte Cugnone, in the high Agri River Valley of Lucania (Fig. Ib). Structural analysis carried out by Mazzoli et al (2001) indicates that the Calcari con Selce Formation, a few hundreds metres thick, represented the mechanically dominant member during contraction and buckling of the sedimentary multilayer, which led to the formation of NW-SE to north-south trending, flexural-slip dominated folds. P-T conditions existing during the development of the analysed vein arrays were mostly controlled by the tectonic burial resulting from the previous contractional episodes. The Monte Cugnone regional structure consists of a faulted anticline of about 1 km wavelength, exposed in the footwall to a major thrust within the Lagonegro units (Marsico Nuovo Thrust; Fig. Ib, c). The area of our detailed study is located in the crestal region of the gently

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Fig. 2. Examples of different types of analysed structures, (a) Conjugate sets of vein arrays in the eastern quarry wall, (b) Line drawing from (a), (c) Intact, conjugate arrays of tensile cracks. Note deflection of the bedding (arrowed), (d) Vein array showing incipient development of discontinuous shear-parallel fractures (arrowed) cutting across en echelon tensile cracks, (e) Faulted shear zone, characterized by a continuous, discrete fault zone (arrowed) breaking through an original en echelon vein array, (f) En echelon vein array (E-W striking) cutting across the fold-thrust structure. Note how the vein array passes from thrust hanging wall to footwall without showing any displacement. Since angles in excess of 60° occur here between the quarry wall and the normal to the thrust transport direction, as well as between the strike of the vein array and the thrust transport direction, overprinting relationships are not just apparent, (g) Line drawing from (f).

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NW-plunging Monte Cugnone anticline, in an area characterized by subhorizontal to gently dipping bedding (Fig. Ic). The quarry is developed on three sides, one roughly trending WSW-ENE, the other two NNW-SSE. The latter two quarry sides are approximately perpendicular to the strike of the vein arrays, providing best exposures for their study (Fig. 2). In detail, outcrop surfaces are differently orientated and irregular, permitting correct measurement of veins and arrays as required for geometrical analysis (Smith 1995). The host rock consists of a strongly anisotropic (on metre scale) but rather homogeneous (on 10 m scale) multilayer, characterized by regular limestone beds (mostly 10 to 40 cm thick) containing chert layers, lenses and nodules. The portion of the Calcari con Selce Formation exposed in the quarry lacks in the marly/clayey intercalations that are typical of other parts of this formation (Scandone 1972). This rules out the possibility that the localization of brittle-ductile shear zones and discrete faults is controlled by competence contrasts between layers of differing composition.

intact arrays of en echelon, calcite-filled tension gashes (Fig. 2c); (ii) vein arrays showing discontinuous shear-parallel fractures cutting across the tensile cracks (Fig. 2d); and (iii) faulted vein arrays, characterized by a continuous, discrete fault zone breaking through, and evidently developed from, an original en echelon array of tensile cracks (Fig. 2e). These discrete fault surfaces are very likely to have formed by linkage of pre-existing shear fractures of the type shown in Fig. 2d. Therefore, a progressive evolution appears to occur from (i) to (ii), to (iii) above, with differently developed structures maintaining similar attitudes (Fig. 3). Most of the en echelon vein arrays are arranged in conjugate sets, generally striking NE-SW and steeply dipping (Fig. 3a). The veins belonging to the arrays are filled with calcite, and consist of a thick central portion that tapers off into narrow tails. They are mostly planar and only in scattered instances show sigmoidal or irregular shapes (probably due to rotation and/or folding as a result of shearing; Ramsay & Graham 1970). In profile, veins are generally tens of centimetres in length and have maximum aperture width of a few centimetres. They mostly strike NE-SW and show steep angles of Outcrop data dip (Fig. 3b). Most of the exposed en echelon vein arrays show Calcite shear fibres and striae from fully evidence of shear displacement, with exten- developed, discrete fault planes sometimes indisional offsets consistently defined by bedding cate oblique-slip with a right-lateral component (Fig. 2), irrespective of the different orientations of motion (Fig. 3c). However, structural features of the quarry faces. Kinematic analysis of en such as: overprinting relationships, invariably echelon vein arrays in conjugate shear zones, showing discrete faults developing from precarried out by means of the conjugate bisector existing en echelon vein arrays; fault attitude, method (e.g. Ramsay & Huber 1987), invariably consistently similar to that of pre-existing en show NE-SW orientated, subhorizontal shear echelon vein arrays; and the geometry of conjuzone intersections (representing the inter- gate faults, which maintain the same angular mediate axis of the related finite strain ellip- relationships (dihedral angles, intersection soid). Acute bisectors (i.e. shortening lines) as those of conjugate shear zones, all indidirections) of conjugate shear zone dihedral cate that faults originally formed as extensional angles are generally subvertical, whereas obtuse structures, and that a later, though minor, bisectors (i.e. extension directions) are subhori- oblique-slip reactivation has occurred. zontal and NW-SE orientated on average. Bedding-parallel shear veins are commonly These features clearly suggest a bulk strain observed. Shear fibres tend to be roughly perdominated by NW-SE horizontal extension, pendicular to the main fold axis; therefore, they compatible with purely dip-slip displacements can be best interpreted as a result of fold-related along conjugate shear zones. A few vein arrays flexural-slip processes. Early layer-parallel showing no detectable shear displacement also shortening (LPS) is documented by pressure occur, suggesting that these structures formed solution cleavage perpendicular to bedding and early, in response to bulk extension, and prob- by minor thrusting with associated outcrop-scale ably pre-dated shear strain localization. En folding. Cross-cutting relationships demonstrate echelon vein arrays possibly represented that LPS and flexural folding pre-dated en original zones of weakness that localized later echelon vein arrays (Fig. 2f, g). Therefore, the shear zone development, as suggested by, for structures analysed in this study can best be example, Roering (1968) and Olson & Pollard related to post-buckling extension, roughly par(1991). allel to the regional, NW-SE trending, fold axis. In outcrop, it was possible to distinguish: (i) Such a deformation seems to accommodate only


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Fig. 3. Orientation data (lower hemisphere, equal area projection), (a) Poles to en echelon vein arrays (mean great circles for each of the two conjugate sets are shown, dipping 73° toward 310°N and 74° toward 129°N). (b) Poles to vein planes from en echelon arrays (mean great circles for the two dominant sets are shown, dipping 79° toward 305°N and 78° toward 127°N). (c) Discrete faults cutting through en echelon vein arrays (with slip vector determined from striae/shear fibres on fault plane; see text), (d) Poles to bedding (mean great circle shown, dipping 12° toward 264°N).

a relatively low amount of strain, at least within the studied outcrop (where bulk extension is 2.7) would not substantially affect the main pattern. This suggests that, for /estimates above 2.7, relatively high values of shear strain tend to be taken up by throughgoing, discrete faults. Fig. 9b also emphasizes that vein array breaching by faulting starts to occur for shear strain values around 1.2. For shear strains in the range of 1.2 < /< 2.7, structures of different types (intact, incipiently faulted and faulted vein arrays) appear to define a transition zone to the development of discrete faults. Discussion Based on the results exposed above, it seems that shear strains below 7= 1.2 can be sustained


by brittle-ductile shear zones, at roughly constant temperature conditions around 130 °C, without significant development of shear parallel fractures breaching the tensile crack array. During these early stages of shear zone evolution, strain accumulation is likely to produce rotation of en echelon tension gashes (e.g. Ramsay & Huber 1987), and/or buckling of inter-vein columns of rock (Olson & Pollard 1991). Both of these processes should result in deformation and folding of en echelon veins (e.g. Ramsay & Graham 1970). However, only a few of our en echelon tension gashes show a sigmoidal shape; most of them are planar or nearly so. Therefore, it seems that shear strain accumulation, eventually leading to fault-slip, occurred in the shear zones without significant folding of en echelon veins. On the other hand, the occurrence of vein rotation during shear zone evolution might be suggested by a graph of vein angle (a) vs. zone boundary angle (/?) (Fig. 10). In a graph of this type (Rothery 1988; Kelly et al 1998), the central line (where a = (3) represents conditions of bulk stretching (corresponding to model A of Ramsay & Huber 1987, their fig. 26.42, with no shear zones forming). The two lines where a- ft = ± 45° represent conditions of simple shear (corresponding to model C of Ramsay & Huber 1987). These lines separate domains characterized by a component of extension (or positive dilation according to Ramsay & Huber 1987, their model B) normal to the shear zone walls from fields of contraction (or negative dilation according to Ramsay & Huber 1987, their model D) normal to the shear zone walls. In a different context, similar domains were defined by Kelly et al (1998) as fields of transtension and transpression, respectively (as these authors were actually dealing with strike-slip faults). The graph in Fig. 10 shows that most of the analysed vein arrays cluster in the domains of 'simple shear + bulk stretching'. In particular, the shear zones tend to plot closer to the central line of pure 'bulk stretching' than incipiently or completely faulted shear zones. The latter, in turn, tend to cluster closer to the lines of simple shear. These features suggest that the analysed shear zones were possibly characterized, during the early stages of their development, by a significant component of positive dilation (according to the model of Ramsay & Huber 1987) and/or of horizontal bulk stretching affecting the whole volume of host rock (as homogeneous strain is required to maintain strain compatibility; Ramsay & Huber 1987). This would have led to the formation of vein arrays characterized by an angle 40 cm). Alternatively, maximum fault size could be constrained by mechanical layering. Once completely developed, through-going discrete faults may be able to take up relatively large displacements, leading to progressive strain localization and to the development, in a cumulative frequency distribution, of a non-power-law 'tail' containing the largest faults. Progressive strain localization onto larger faults, accompanied by the abandonment of smaller structures, can in fact produce a change in the active fault population from power-law to scale-bound (Walsh et al 2003). Clearly, the very few data points in the right hand 'tail' of the diagram in Fig. 9(a) are statistically not significant to discriminate between the possible different hypotheses. Conclusions This study emphasizes the fundamental role played by finite strain in the process of fault initiation by the coalescence of en echelon vein arrays in naturally deformed limestones. Fluid inclusion analysis suggests that the different sets of structures recognized in the field developed at approximately constant temperature. Critical values of displacement (D ~ 9 cm) and shear strain (7 ~ 1.2) were determined by Mazzoli & Di Bucci (2003) for the onset of normal fault nucleation from shear zones in the studied limestones. Structures displaying shear displacements below these critical values consist, for the vast majority (Fig. 9, Table 1), of brittle-ductile shear zones as defined by Ramsay

371

& Huber (1987). The deformation processes responsible for their formation are likely to be of dominant viscous type (i.e. intracrystalline deformation, solution mass transfer), accompanied by brittle fracturing (as indicated by the coeval formation of tension gashes). Such deformation processes are likely to become progressively less important as shear-parallel fractures, characterizing the incipient stages of fault nucleation, start to form. Eventually, once throughgoing faults are developed, deformation becomes dominantly frictional. In this context, during progressive strain localization leading to incipient fault nucleation (i.e. for 9 < D < 17 cm and 1.2 < / < 2.7; Fig. 9b), a transition from a dominant viscous to a frictional behaviour would occur. In the studied area, shear zone evolution may be reconstructed as follows: (i) strain localization occurs along en echelon crack arrays (e.g. Olson & Pollard 1991); (ii) en echelon veins, mostly formed originally at angles less than 45° with respect to shear zone walls, progressively rotate within the shear zones; this process is probably accompanied by thickening of the shear zones, mostly without significant folding of the tension gashes themselves; (iii) vein rotation and increasing shear zone displacement produce shear stress localization, leading to linkage of the veins by shear-parallel fractures and eventually producing brittle shear. In conclusion, our results, concerning normal fault initiation from en echelon vein arrays in limestones deformed at very low-grade conditions, should hopefully be applicable to the analysis of fault growth processes within the upper crust. Similar studies, carried out in rocks characterized by various rheological behaviours and/or deformed under different environmental conditions, may be useful to implement existing models of fault nucleation, thereby improving our comprehension of the processes of fault initiation and early development. This study greatly benefited from discussions with Daniela Di Bucci. Thorough and constructive reviews by W. Bailey and M. Barchi substantially helped to improve the paper. Financial support from the Italian MIUR Cofin 2002 (Resp. G. Cello, prot. 2002043912) and Universita di Urbino are gratefully acknowledged.

References BATHURST, R.G.C. 1971. Carbonate sediments and their diagenesis. Elsevier, Amsterdam. BEACH, A. 1975. The geometry of en echelon vein arrays. Tectonophysics, 28, 245-263. BODNAR, R.J. 1993. Revised equation and table for determining the freezing point depression of

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H^O-NaCl solutions. Geochimica et Cosmochimica Acta, 57, 683-684. BODNAR, R.J. & VITYK, M.O. 1994. Interpretation of microthermometric data for H2O-NaCl fluid inclusions. In: DE Vivo, B. & FREZZOTTI, M.L. (eds) Fluid inclusions in minerals: methods and applications. Virginia Tech, Blacksburg, 117-130. BURKHARD, M. 1993. Calcite twins, their geometry, appearance, and significance as stress-strain markers and indicators of tectonic regime: a review. Journal of Structural Geology, 15, 351-368. CARBONE, S., CATALANO, S., LAZZARI, S., LENTINI, F. & MONACO, C. 1991. Presentazione della carta geologica del Bacino del Fiume Agri (Basilicata). Memorie della Societd Geologica Italiana, 47, 129-143. CELLO, G. & MAZZOLI, S. 1999. Apennine tectonics in southern Italy: a review. Journal ofGeodynamics, 27,191-211. CHILDS, C., NICOL, A., WALSH, J.J. & WATTERSON, J. 1996a. Growth of vertically segmented normal faults. Journal of Structural Geology, 18, 1389-1397. CHILDS, C., WATTERSON, J. & WALSH, JJ. 19966. A model for the structure and development of fault zones. Journal of the Geological Society of London, 153, 337-340. CINQUE, A., PATACCA, E., SCANDONE, P. & Tozzi, M. 1993. Quaternary kinematic evolution of the Southern Apennines. Relationships between surface geological features and deep lithospheric structures. Annali di Geofisica, 36, 249-260. COWIE, PA. & SCHOLZ, C.H. 1992. Physical explanation for the displacement-length relationships of faults using a post-yield fracture mechanics model. Journal of Structural Geology, 14, 1133-1148. Cox, SJ.D. & SCHOLZ, C.H. 1988a. An experimental study of shear fracture in rocks: Mechanical observations. Journal of Geophysical Research, 93,3307-3320. Cox, SJ.D. & SCHOLZ, C.H. 19886. On the formation and growth of faults. Journal of Structural Geology, 10,413-430. DURNEY, D.W. & RAMSAY, J.G 1983. Incremental strains measured by syntectonic crystal growth. In: DE JONG, K.A. & SCHOLTEN, R. (eds) Gravity tectonics. Wiley, New York, 67-96. ETCHECOPAR, A., GRANIER, T. & LARROQUE, J.-M. 1986. Origin des fentes en echelon: propagation des failles. Comptes Rendues de I'Academie des Sciences, Paris, 302, 479-484. FROST, H.S. & ASHBY, M.F. 1982. Deformationmechanism maps. Pergamon Press, Oxford. GILLESPIE, P. A., HOWARD, C.B., WALSH, JJ. & WAITERSON, J. 1993. Measurement and characterisation of spatial distribution of fractures. Tectonophysics, 226,113-141. GOLDSTEIN, R.H. & REYNOLDS, TJ. 1994. Systematics of fluid inclusions in diagenetic minerals. SEPM Short Course 31, Tulsa, USA. HYPPOLYTE, J.-C, ANGELIER, J. & ROURE, F. 1994. A major geodynamic change revealed by Quater-

nary stress patterns in the Southern Apennines (Italy). Tectonophysics, 230,199-210. KELLY, P.G., SANDERSON, DJ. & PEACOCK, D.C.P. 1998. Linkage and evolution of conjugate strike-slip fault zones in limestones of Somerset and Northumbria. Journal of Structural Geology, 20, 1477-1493. KNIPE, RJ. & WHITE, S.H. 1979. Deformation in low grade shear zones in Old Red Sandstone, S.W. Wales. Journal of Structural Geology, 1, 53-66. KNOTT, S.D., BEACH, A., BROCKBANK, P.J., BROWN, J.L., MCCALLUM, IE. & WELBON,A.I. 1996. Spatial and mechanical controls on normal fault populations. Journal of Structural Geology, 18, 359-372. LAJTAI, E.Z. 1969. Mechanics of second order faults and tension gashes. Geological Society of America Bullettin, 80, 2253-2272. LAWN, B.R. & WiLSHAW,T.R. 1975. Fracture of brittle solids. Cambridge University Press, Cambridge. MARCHAL, D., GUIRAUD, M. & RIVES, T. 2003. Geometric and morphologic evolution of normal fault planes and traces from 2D to 4D. Journal of Structural Geology, 25,135-158. MARTEL, S.J., POLLARD, D.D. & SEGALL, P. 1988. Development of simple strike-slip fault zones, Mount Abbott Quadrangle, Sierra Nevada, California. Geological Society of America Bulletin, 100,1451-1465. MAZZOLI, S. & Di Bucci, D. 2003. Critical displacement for normal fault nucleation from en-echelon vein arrays in limestones: a case study from the southern Apennines (Italy). Journal of Structural Geology, 25,1011-1020. MAZZOLI, S., BARKHAM, S., CELLO, G, GAMBINI, R., MATTIONI, L., SHINER, P. & TONDI, E. 2001. Reconstruction of continental margin architecture deformed by the contraction of the Lagonegro Basin, southern Apennines, Italy. Journal of the Geological Society of London, 158, 309-319. McGRATH, A. & DAVISON, 1.1995. Damage zone geometry around fault tips. Journal of Structural Geology, 17,1011-1024. MONACO, C, TORTORICI, L. & PALTRINIERI, W 1998. Structural evolution of the Lucanian Apennines: Journal of Structural Geology, 20, 617-638. NEEDHAM,T., YELDING, G. & Fox, R. 1996. Fault population and prediction using examples from the offshore UK. Journal of Structural Geology, 18, 155-167. NICOLAS, A. & POIRIER, J.-P. 1976. Crystalline plasticity and solid state flow in metamorphic rocks. Wiley, New York. OLSON, J.E. & POLLARD, D.D. 1991. The initiation and growth of en echelon veins. Journal of Structural Geology, 13, 595-608. PEACOCK, D.C.P. & SANDERSON, DJ. 1995. Pull-aparts, shear fractures and pressure solution. Tectonophysics, 241,1-13. POLLARD, D.D., SEGALL, P. & DELANEY, P.T. 1982. Formation and interpretation of dilatant echelon cracks. Geological Society of America Bulletin, 93, 1291-1303. PRICE, NJ. & COSGROVE, J.W 1990. Analysis of geological structures. University Press, Cambridge.

SHEAR ZONE EVOLUTION AND FAULTING RAMSAY, J.G. & GRAHAM, R.H. 1970. Strain variations in shear belts. Canadian Journal of Earth Sciences, 7, 786-813. RAMSAY, J.G. & HUBER, M.I. 1983. The techniques of modern structural geology. Volume 1: Strain analysis. Academic Press, London. RAMSAY, J.G. & HUBER, M.I. 1987. The techniques of modern structural geology. Volume 2: Folds and fractures. Academic Press, London. ROERING, C. 1968. The geometrical significance of natural en echelon crack arrays. Tectonophysics, 5,107-123. ROTHERY, E. 1988. En echelon vein array development in extension and shear. Journal of Structural Geology, 10, 63-71. SCANDONE, P. 1972. Studi di geologia lucana: carta dei terreni della serie calcareo-silico-marnosa e note illustrative. Bollettino della Societd dei Naturalisti in Napoli, 81, 225-300. SCHOLZ, C.H. 1990. The mechanics of earthquakes and faulting. University Press, Cambridge. SEGALL, P. & POLLARD, D.D. 1983. Joint formation in granitic rocks of the Sierra Nevada. Geological Society of America Bullettin, 94, 454-462. SMITH, J.V. 1995. True and apparent geometric variability in en echelon vein arrays. Journal of Structural Geology, 17,1621-1626. SMITH, J.V. 1996. Geometry and kinematics of conver-

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gent conjugate vein array systems. Journal of Structural Geology, 18,1291-1300. SMITH, J.V. 1997. Initiation of convergent extension fracture vein arrays by displacement of discontinuous fault segments. Journal of Structural Geology, 19,1369-1373. SRIVASTAVA, D.C. 2000. Geometrical classification of conjugate vein arrays. Journal of Structural Geology, 22, 713-722. WALSH, J.J. & WATTERSON, I 1988. Analysis of the relationship between displacements and dimensions of faults. Journal of Structural Geology, 10, 239-247. WALSH, J.J., CHILDS, C., IMBER, I, MANZOCCHI,T, WATTERSON, J. & NELL,P.A.R. 2003. Strain localisation and population changes during fault system growth within the Inner Moray Firth, Northern North Sea. Journal of Structural Geology, 25, 307-315. WILLEMSE, E.J.M., PEACOCK, D.C.P & AYDIN, A. 1997. Nucleation and growth of strike-slip faults in limestones from Somerset, UK. Journal of Structural Geology, 19,1461-1477. YELDING, G, NEEDHAM,T. & JONES, L. 1996. Sampling of fault population using sub-surface data: a review. Journal of Structural Geology, 18, 135-146.


Index Page numbers in italic refer to figures. Those in bold refer to entries in tables. alpine-type peridotite massifs 11-12 localized deformation 13 Anderson-Byerlee frictional fault mechanics 95 brittle-ductile shear zone evolution and fault initiation at Monte Cugnone, Italy 353-355, 368-372 cross-section 354 examples of structures analysed 356-357 fluid inclusion petrography and microthermometry 363-366 geological setting 354, 355-358 kinematics 366-368 microstructures 363,364 outcrop data 358-363 structural data 360-361 conjugate shearing domain (CSD) 219-221,220 continental crust, metamorphic signature of subduction in Corsica/northern Apennine orogen 321-322, 329-331 structural and metamorphic history of inner Tuscan metamorphic units 329 structural and metamorphic history of Tenda Massif deformation history 324-328,325, 326, 327 geological outline 323-324,324 metamorphic history 328-329,328,329 tectonic setting 322-323,322, 323 crenulation-slip development in NW Ireland 337-338, 350 evidence for strike-slip motion in Central Ox Mountains 344 extensional crenulation cleavages 344 high-strain zones 344 evidence for strike-slip motion in Mayo 339-341 asymmetrical buckle folds 341,342 crenulation-slip morphologies produced by oblique foliation-slip 341 extensional crenulation cleavages 341-342 orientation of D3 dextral shear zone 342-344, 343 predicted angular relationships 341 geological significane of the Fair Head-Clew Bay Line 338-339 isotopic dating of crenulation-slip surfaces 344-346 analytical methods 347 40 Ar/39Ar spot fusion data 347 constraining dextral shear in NW Mayo 347-348 constraining sinistral shear in Central Ox Mountains 348-349 intrusion age of Ox Mountains granodiorite 349 pre-existing age constraints on foliation-slip fabric 346 Rb-Sr geochronology 349 sampling 346 regional geology 338

dauphine twinning and misorientation 39, 54, 58-59 microstructural evolution 54-55 microstructural stability 56-57 misorientation angle distributions 57-58 study details crystallographic misorientation analysis 43-44, 44 grain boundary (mis)orientation analysis 44-45 relationship between crystal slip and boundary orientation 45 relationships between quartz crystal slip systems 46 relationships between specific quartz crystal slip systems 45 summary of Loch Torridon shear zone data 43 study results and interpretations boundary formation 52-54,53 dauphine twinning and twin boundaries 54 microstructure and LPO 47 misorientation analysis 47-52 petrofabric and misorientation analysis 51 SEM/EBSD analysis of dauphine twin microstructures 48-49 (sub)grain boundary formation 55-56 deformation in a complex crustal-scale shear zone 229-230, 246-247 Archaean granitic gneiss 230-231,233 Errabiddy Shear Zone 230,231 evolution of Capricorn Orogen 232 felsic gneiss and Erong Shear Area 234-237,237, 238, 239 kinematic evolution of Errabiddy Shear Zone 244 palaeoproterozoic metasedimentary rocks - Camel Hills 239 deformation in migmatized pelitic schist and gneiss 241-242,242 deformation in psammitic gneiss 239-241,240, 241 structural geometry within Errabiddy Shear Zone 243-244 summary 243 temporal and tectonic evolution of Errabiddy Shear Zone 244-246,245 ductile shearing 161-162,173-174 basement lithology on Sikinos 163-164,164 basement-cover contact on Sikinos 172 geology of cover sequence on Sikinos 162-163 maps 162,163 high-pressure metamorphic imprint in Cycladic basement 172-173 pressure-temperature conditions of metamorphism on Sikinos 170-172,171 feldspar porphyroclast populations, kinematics and strain 265-266 application to western Idaho shear zone 277-278, 279

376

INDEX

assumptions 279-281 feldspar shape preferred orietation data 281 field conclusions 284 implications for shear zone studies 284 location map 275 model conclusions 283-284 quantification of field data 281-282 study results 282-283,283 forward model of clast rotation 266-268 construction of fabric ellipsoid 269 three-dimensional description of clast orientation 268-270 model results 270 fabric ellipsoid versus finite strain ellipsoid 277 orientation of fabric ellipsoid to shear sense 276-277,277 populations of oblate clasts 276 populations of prolate clasts 272-276,273, 274, 275 rotation of single clasts 270 rotations of populations of clasts 271-272 single oblate clasts 270-271 single prolate clasts 270,271 flattening strain 252-253,253, 253 fluid-rock interactions in West Fissure Zone, Chile 141-142 comparison with San Andreas Fault, California 157-158 description of fault rocks 143-147,147 sampling profiles 146 fluid sources and fluid composition 156-157 geochemistry of fault rocks carbon and oxygen isotope relationships in calcite 148-149,149,150 fluid inclusions 147-148,148, 155-156 major elements 150-154,153,156,156 oxygen versus distance relationships in monzodiorite 154-155,154,155 trace elements 149-150,150-154,151,152,153, 156,156 geological setting 142-143 regional map 144-145 regional overview 142 stratigraphy 143 sampling and analytical methods 143 variations in fluid-rock interaction 157 Geographic Information Systems (GIS) applied to deformation patterns 73-76 aeromagnetic dataset 68-70,69-70 combined and integrated dataset combination of all available datasets 72-73, 73 combined directional structural data and shaded total magnetic field map 70-72, 71 vertical gradient of total magnetic field, foliation trends and metamorphic data combined 72, 73 database 65 directional structural dataset 66-68,67 fabric type dataset 68 lithological dataset 65 metamorphic dataset 65-66 proposed indentor model 74, 75 west Greenland case study 64-65,64,66 geostatistical analysis

kriging interpolation 305 variogram computation and interpretation 303-305,304 grain-size sensitive (GSS) flow 32,33, 34,34 granulites, instability and deformation localization in the lower crust 25-26, 35-36 Clarke Head megabreccia 26-27,26 deformation microstructures cherty ultramylonite 30,31 host mylonite 27,28 ultramylonite 27-29,29, 30 experimental procedures 27 interpretation of microstructures and deformation deformation environment 31-32 deformation mechanisms 32-33,33, 34 deformation partitioning and localization 34-35 mechanism transitions 33-34 high-pressure metamorphism 161-162,173-174 basement lithology on Sikinos 163-164,164 basement—cover contact on Sikinos 172 geology of cover sequence on Sikinos 162-163 maps 162,163 high-pressure metamorphic imprint in Cycladic basement 172-173 pressure-temperature conditions of metamorphism on Sikinos 170-172,171 hydrous fluid channelling 161-162,173-174 basement lithology on Sikinos 163-164,164 basement-cover contact on Sikinos 172 geology of cover sequence on Sikinos 162-163 maps 162,163 high-pressure metamorphic imprint in Cycladic basement 172-173 pressure-temperature conditions of metamorphism on Sikinos 170-172,171 indentor tectonics 73-76 aeromagnetic dataset 68-70,69-70 case study in west Greenland 64-65, 64,66 combined and integrated dataset combination of all available datasets 72-73, 73 combined directional structural data and shaded total magnetic field map 70-72, 71 vertical gradient of total magnetic field, foliation trends and metamorphic data combined 72, 73 directional structural dataset 66-68,67 fabric type dataset 68 lithological dataset 65 metamorphic dataset 65-66 proposed indentor model 74, 75 kriging interpolation 305 lattice preferred orientation (LPO) 39,54,58-59 dauphine twinning and (sub)grain boundary formation 55-56 dauphine twinning and microstructural evolution 54-55 dauphine twinning and microstructural stability 56-57 misorientation angle distributions 57-58 shear zone grain size reduction model 55 study details

INDEX crystallographie misorientation analysis 43-44, 44 grain boundary (mis)orientation analysis 44-45 relationship between crystal slip and boundary orientation 45 relationships between quartz crystal slip systems 46 relationships between specific quartz crystal slip systems 45 samples 39-41, 40, 42 SEM/EBSD technique 41-43, 42 summary of Loch Torridon shear zone data 43 study results and interpretations boundary formation 52-54,53 dauphine twinning and twin boundaries 54 microstructure and LPO 47 misorientation analysis 47-52 petrofabric and misorientation analysis 51 SEM/EBSD analysis of dauphine twin microstructures 48-49 low angle normal faults (LANFs) 95-97,105-109 active versus exhumed LANFs Altoberina Fault in Umbria region 97-102, 98 Zuccale Fault in Isle of Elba 102-105,105 regional setting of Northern Apennines 97 lower crust granulites, instability and deformation localization 25-26, 35-36 Clarke Head megabreccia 26-27,26 deformation microstructures cherty ultramylonite 30,31 host mylonite 27,28 ultramylonite 27-29,29, 30 experimental procedures 27 interpretation of microstructures and deformation deformation environment 31-32 deformation mechanisms 32-33,33, 34 deformation partitioning and localization 34-35 mechanism transitions 33-34 microstructure evolution during deformation lower crust granulites 33 mylonitic quartz simple shear zone 39 study details 39-46 study results and interpretation 47-54 misorientation analysis 39, 54, 58-59 crystallographie relationships 43-44 crystal slip systems and boundary orientations 45 quartz 44, 45,46 dauphine twinning and (sub)grain boundary formation 55-56 dauphine twinning and microstructural evolution 54-55 dauphine twinning and microstructural stability 56-57 grain boundary analysis 44-46 microstructure and LPO 47 misorientation angle distributions 57-58 SEM/EBSD technique 41-43,42 localized dauphine microstructures 48-49 shear zone grain size reduction model 55 study results 47-50 boundary formation 52-54,53 dauphine twinning and twin boundaries 54 localized dauphine microstructures 48-49

377

misorientation angle distributions 50 misorientation axis/angle pairs 50-52 petrofabric and misorientation analysis 51 study samples 39-41,40 Nabarro-Herring creep 81,81 ophiolite-type peridotite massifs 11-12 localized deformation 13 orthorhombic fabrics, development within a simple shear sinistral transpression zone 215-216 Arronches gneisses, structural analysis conjugate shearing domain (CSD) 219-221,220 grain-size reduction and deformation mechanisms in fabric formation 221 intermediate sinistral domain (ISD) 221 peralkaline gneisses 216-219 sinistral domain (SD) 221 Arronches Tectonic Unit regional setting 216 structure and metamorphism 216 study area 27 7, 218, 219 conjugate shear band formation 224-226,225 dynamic recrystallization and development of fabric and texture 221-224,224, 225 relative timing of orthorhombic and monoclinic fabric formation 224 partially molten rocks (PMR), application of twophase rheology 79-81, 91 experiments 89-90 development of instabilities 90-91 non-linear effects 90 implications for other two-phase systems 87-89 importance of shear deformation 89 main principles 81-82 schematic map of plastic deformation 81 stress versus strain diagram 82 Theological responses 84-87,86,87 rheology of two-phase materials 82-84,83, 84 pelitic rocks, shear deformation 113,121-124 composition of sheared clays 116,116 srnectic/illite transformation 116-117,117 geological framework of Scorciabuoi Fault (SBF) 113-115,774,775,71(5 grain size analysis 118-119,118 shear zone fabric 119, 719 fine scale analysis 119-121,122,123 mesoscale observations 119,120,121 peridotite mylonites 16,17-19 plane strain 252-253,253, 253 plastic deformation, application of two-phase rheology 79-81, 91 experiments 89-90 development of instabilities 90-91 non-linear effects 90 implications for other two-phase systems 87-89 importance of shear deformation 89 main principles 81-82 schematic map of plastic deformation 81 stress versus strain diagram 82 rheological responses 84-87,86,87 rheology of two-phase materials 82-84,83 plate convergence, shear and fluid flow 127

378

INDEX

comparisons and contrasts between the study sites 135 deformation structures with faults 136-137 fault thickness 136 fault-zone margins 137 hydrogeology 137 internal geometry of fault zones 137 lithological influence of propagation 136 summary of features 138 deformation features Barbados 129-130, 729 Costa Rica 130-131,130 Nankai 131,131 fluid transport Barbados 131-133,132,133,134 Costa Rica 133-134,135 Nankai 134-135,136 implications from study sites for other mega-shear zones 137-138 tectonic settings Barbados 128,128 Costa Rica 128-129,128 Nankai 128,129 rheology of two-phase materials 79-81, 91 experiments 89-90 development of instabilities 90-91 non-linear effects 90 extrapolating the end-members 83 implications for other two-phase systems 87-89 importance of shear deformation 89 main principles 81-82 schematic map of plastic deformation 81 stress versus strain diagram 82 rheological responses 84-87,86,87 thermodynamic considerations 83-84,83, 84 rigid percolation threshold (RPT) 80 shear zone folds 177-178,196-197 Caledonian Moine Nappe, Sutherland 179-180, 180,181 curvilinear fold patterns and evolution 189 fold evolution model 189-194 fold inheritance model 194-195,195,196 hybrid fold model 195-196 fold types 178 sheath folds 178-179 synshearing flow folds 179 Melness folds case study 181-185,183,185 topological relationships between sheath folds and synshearing folds 186-189,188,189, 190-191,192-193,194 transection relationships between sheath folds and synshearing folds 185-186,186,187 shear zones 1, 8 fault controls and shear zone development 4 anastomosis around low-strain augen 5 grain-scale controls 4 lithospheric-scale controls 4 network geometry-scale processes 4-5 histories 7-8 lithosphere deformation and rheology of shear zones 5-6,5

occurrence on different scales 1, 2 partitioning processes in shear zones 6-7, 7 strength, strain rate histories and fault rocks at depth 1-4 deformation regimes and typical fault rocks 3 schematic strength profile through crust and upper mantle 4 strain and deformation history in a syntectonic pluton, Roses granodiorite 307-308, 315-318 displacement versus width diagram 318 main lithological units 308 progressive development of structures in Roses granodiorite 308 elongated enclave of quartz diorite 313 geological setting 309 late brittle fractures 314 leucocratic dykes 313 magmatic fabric and enclaves 308-313 mesoscopic scale structures 311 pre-dyke finite strains 312 qualitative model and structural history 310 shear zones and associated mylonites 313-314 shear strain analysis 317 structural map and strain analysis 316 structure and strain profiles 314 deformation postdating dykes 315 deformation predating dykes 314-315 strain removal within Hercynian Shear Belt, methodology and tectonic implications 287, 287, 300 data processing cleavage trajectory model 292-293,293 domainal distribution of cleavage directions 294-295,294 geostatistical analysis of cleavage directions 291-292,292 geological setting 288 lithologies 288-289 structures 289 kinematic data cleavage and finite strain ellipsoid 289,290, 291 deformation regime 289-291,297 model validation and regional implications 296 at the boundaries 298-299,298 within restored area 296-298,297 restoration of eastern Central Brittany 295-296, 296 strike-slip deformation 250,251 tectonites relative softening fine grained 16-17 medium-to-coarse grained 16 structures and microstructures fine grained 15 medium-to-coarse grained 15 transpression terrane boundaries, geometric and kinematic analysis 201-202, 213 fault rocks 203 fault zone deformation mechanisms 211-212 fault zone kinematics 210-211 framework of the Minas fault system 202-203 location of the Minas fault system 202

INDEX strain partitioning and localization 212-213 structural elements and geometric relationships 204, 206, 207, 208, 209, 210 crenulation cleavage 208-209 faults 209-210 folds 203 foliations 203-205 lineations 205 S-C fabrics 208 shear bands 209 veins 205-208 transpressional high-strain zones, strain and vorticity analysis 249-250, 262 interpretation at study areas Brookneal high-strain zone (BHSZ) 258-259, 259 Spotsylvania high-strain zone (SHSZ) 259-260, 259, 260 strain compatibility 260-261 tectonic significance of Piedmont high-strain zones 261-262 kinematic deformation models 249 kinematic vorticity and vorticity analysis 251-253, 252, 253, 253 transpression and general shear 250-251,251

379

upper mantle shear zones 11-12,19-20, 21 features 12 implications for mantle strength 21 possible tectonite and mylonite shear zones 20 relative softening mechanisms 16 fine grained tectonites 16-17 mantle cross-section at Hilti, Oman 17 medium-to-coarse grained peridotite tectonites 16 olivine deformation mechanism map for Othris, Greece 17 peridotite mylonites 17-19 P-Tgridl9 SEM image of grain boundary alignments for Turon de Tecouere, France 17 SEM images of fine grain production 18 structures and microstructures 12-15 fine grained tectonites 15 localized deformation 13 medium-to-coarse grained tectonites 15 peridotite mylonites 16 photomicrographs 14 variogram computation and interpretation 303-305, 304 vorticity 251-253,252

Flow Processes in Faults And Shear Zones (Geological Society Special Publication No. 224)

Early Precambrian Processes (Geological Society Special Publication)

Hydrothermal Vents and Processes (Geological Society Special Publication No. 87)

Geoarchaeology: Exploration, Environments, Resources (Geological Society Special Publication, No. 165) (Geological Society Special Publication, No. 165)

Geological Data Management (Geological Society Special Publication)

Geomaterials in Cultural Heritage (Geological Society Special Publication No. 257)

Hydrocarbons in Crystalline Rocks (Geological Society Special Publication No. 214)

Climates: Past and Present (Geological Society Special Publication No. 181)

Australian Landscapes (Geological Society Special Publication 346)

Martian Geomorphology (Geological Society Special Publication 356)

Palaeozoic Reefs and Bioaccumulations: Climatic and Evolutionary Controls - Special Publication no 275 (Geological Society Special Publication)

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Sustainable Groundwater Development (Geological Society Special Publication,)

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Fractured Reservoirs (Geological Society Special Publication no 270)

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Deformation of Glacial Materials (Geological Society Special Publication No. 176)

Devonian Events and Correlations : Special Publication no 278 (Geological Society Special Publication)

The Evolving Continents: Understanding Processes of Continental Growth - Special Publication 338 (Geological Society Special Publication)

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Modern Ocean Floor Processes and the Geological Record (Geological Society Special Publication)

Fractal Analysis for Natural Hazards - Special Publication No 261 (Geological Society Special Publication)

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Geological Processes on Continental Margins: Sedimentation, Mass- Wasting And Stability (Geological Society Special Publication)

Structurally Complex Reservoirs - Special Publication no 292 (Geological Society Special Publication)

Flow Processes in Faults And Shear Zones (Geological Society Special Publication No. 224)

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Hydrothermal Vents and Processes (Geological Society Special Publication No. 87)

Geoarchaeology: Exploration, Environments, Resources (Geological Society Special Publication, No. 165) (Geological Society Special Publication, No. 165)

Geological Data Management (Geological Society Special Publication)

Geomaterials in Cultural Heritage (Geological Society Special Publication No. 257)

Hydrocarbons in Crystalline Rocks (Geological Society Special Publication No. 214)

Climates: Past and Present (Geological Society Special Publication No. 181)

Australian Landscapes (Geological Society Special Publication 346)

Martian Geomorphology (Geological Society Special Publication 356)

Palaeozoic Reefs and Bioaccumulations: Climatic and Evolutionary Controls - Special Publication no 275 (Geological Society Special Publication)

Communicating Environmental Geoscience - Special Publication no 305 (Geological Society Special Publication) (No. 305)

Sustainable Groundwater Development (Geological Society Special Publication,)

Dating and Duration of Fluid Flow and Fluid-Rock Interaction (Geological Society Special Publication No. 144)

Orogenic Processes: Quantification and Modelling in the Variscan Belt (Geological Society Special Publication, No. 179)

Phanerozoic Ironstones (Geological Society Special Publication)

Slope Tectonics (Geological Society Special Publication 351)

Volcanic Degassing (Geological Society Special Publication No. 213)

Fractured Reservoirs (Geological Society Special Publication no 270)

Confined Turbidite Systems (Geological Society Special Publication No. 222)

Deformation of Glacial Materials (Geological Society Special Publication No. 176)

Devonian Events and Correlations : Special Publication no 278 (Geological Society Special Publication)

The Evolving Continents: Understanding Processes of Continental Growth - Special Publication 338 (Geological Society Special Publication)

Non-Marine Permian Biostratigraphy and Biochronology - Special Publication No. 265 (Geological Society Special Publication)

Sedimentary Processes in the Intertidal Zone (Geological Society Special Publication No. 139)

Modern Ocean Floor Processes and the Geological Record (Geological Society Special Publication)

Fractal Analysis for Natural Hazards - Special Publication No 261 (Geological Society Special Publication)

Fine-Grained Sediments: Deep-Water Processes and Facies (Geological Society special publication No. 15)

Geological Processes on Continental Margins: Sedimentation, Mass- Wasting And Stability (Geological Society Special Publication)

Structurally Complex Reservoirs - Special Publication no 292 (Geological Society Special Publication)

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