Extending a Continent: Architecture, Rheology and Heat Budget (Geological Society Special Publication)

Extending a Continent: Architecture, Rheology and Heat Budget

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) RANDELL STEPHENSON (UK)

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk.

It is recommended that reference to all or part of this book should be made in one of the following ways: RING , U. & WERNICKE , B. (eds) 2009. Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321. BOULTON , C., DAVIES , T. & MC SAVENY , M. 2009. The frictional strength of granular fault gouge: application of theory to the mechanics of low-angle normal faults. In: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 9–31.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 321

Extending a Continent: Architecture, Rheology and Heat Budget

EDITED BY

U. RING University of Canterbury, New Zealand

and B. WERNICKE California Institute of Technology, USA

2009 Published by The Geological Society London

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Preface This book is the outcome of a Penrose Conference (8–12 October 2007) that examined processes that contribute to horizontal extension of continental lithosphere and the origin of oceanic basins. Over the last three decades, there has been a growing appreciation of the role of extensional tectonics in convergent orogens and Wernicke’s Introduction to this book provides a flavour of how this era changed our views on tectonometamorphic relationships in orogenic belts. This trend was initiated by the discovery of highly attenuated crustal sections in the Basin and Range province and the recognition that the attenuation was caused by regional-scale horizontal extension, as manifested by low-angle normal faults or detachments. Soon afterward extensional detachments were recognized as a global phenomenon in most orogens. The Aegean Sea is another well-known example of an orogen that has been extremely modified by large-scale continental extension. The majority of papers of this volume are devoted to the Aegean extensional province. Horizontal extension in the Aegean occurs directly above the Hellenic subduction zone and affects the forearc as well as the backarc region. Lithospheric extension in the Aegean is widely thought to be caused by slab rollback and started across the entire area rather abruptly at the beginning of the Miocene. New research frontiers in the field of continental extensional tectonics are seen in the application of geodesy for addressing steady/non-steady-state behaviour of the lithosphere during extension. Ideally the emerging new results from this approach should be combined/compared with studies on deeply exhumed fossil extensional complexes. A key issue in the latter is to merge our understanding and gain more information on the exact timing of deformation with magmatic, metamorphic and tectonic microstructures. A general problem is that the new developments and achievements in numerical modelling far outstrip natural observations, stressing the importance of more and detailed field-based research in much closer interaction with numerical models than has previously been done. The volume contains 11 papers. The opening contribution by Wernicke provides a historic view our ideas about large-scale tectonic contacts in mountain belts have changed over the years. Wernicke concludes that controversy still persists over the existence and mechanics of slip on shallowly dipping extensional detachments. Boulton et al. address this issue. They summarize recent geological, experimental, and numerical research into the frictional behaviour of rocks.

They posit that dynamic fragmentation in granular fault cores may result in low static and dynamic frictional strengths, thereby allowing faults to fail under low resolved shear stresses. Marotta et al. report a 2D thermo-mechanical modelling study that aims at reducing the ambiguity on the geodynamic significance of Permotriassic high-temperature/low-pressure metamorphism, igneous activity and sedimentary records in the pre-Alpine continental crust in the Alps. The modelling predictions are compared with: (1) Peak PT conditions of Permotriassic metamorphism in the continental crust of the Penninic and Southalpine domains; (2) mafic intrusions and associated intermediate to acidic magmatism occurring mainly in the Austroalpine and Southalpine domains; and (3) structural and volcanic activity and coeval formation of sedimentary basins. The agreement between field data and model predictions supports the idea that large-scale lithospheric extension was responsible for the Permotriassic geological record in the Alps. The authors develop a model of bottom-up periodic strain localization at 5–10 Ma intervals that are comparable with volcanic pulses, episodic basin deepening, periodic sedimentary facies fluctuations and related episodic faulting. Stern compares short- and long-term deformation in the Central Volcanic Region in central North Island of New Zealand. Both rotation and translation of the backarc can be described by observations of volcanic arc migration, GPS, geodesy and paleomagentism. Short- (10 years) and long-term (5Ma) measures of extension (6–17 mm a1) and rotation (c. 6 Ma1) are surprisingly consistent. Heat flux in the Central Volcanic Region is of the order of 4.3 GW, which when averaged over the region of surface geothermal activity gives an effective heat-flow of c. 860 mWm2: a value many times greater than other backarc basins. Stern argues that other continental backarcs have histories of episodic and rapid extension, and that the central North Island is in one of these phases. Tulloch et al. use high precision geochronology to reveal a strong episodicity in silicic magmatism in the interval between cessation of East Gondwana arc magmatism at 105 Ma and rifting of Zealandia at c. 83 Ma. They show that both 101 and 97 Ma groups of rhyolites and tuffs occur across the entire width of Zealandia from near the paleotrench to the continental interior, indicating distinct phases of widespread, near instantaneous, and subparallel, extension. The authors interpret these features to be most consistent with rifting being caused either by

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PREFACE

basal traction on a subducted slab that had been captured and pulled oceanwards by the Pacific plate, and/or southwestwards propagation of a ,83Ma oceanic ridge between Zealandia and West Antarctica. Skourtsos & Kranis describe the geometry and the kinematics of normal faults exposed onshore along the southern part of the Corinth Rift in Greece, proposing that the southern margin is probably located further to the south of what it is considered to be. They explain the evolution of the Corinth Rift by the presence of a structurally lower detachment fault that soles northward into the estimated detachment zone beneath the Gulf of Corinth. Thomson et al. tackle the controversial question over the origin of metamorphic tectonites in core complexes. The combine structural field data with robust geochronologic ages on mylonitization, fission-track and (U –Th)/He cooling ages and PT data to constrain the timing, kinematics and architecture of extension during exhumation of the Ios metamorphic core complex. Their work corroborates that large-scale extensional deformation in the central Aegean did not commence before the early Miocene. Ring et al. use zircon and apatite fission-track ages from the poorly known Amorgos detachment system in the southeastern Aegean for discussing the role the Amorgos detachment played in the opening of the Cretan Sea basin. The authors argue that the Amorgos detachment is associated with the large-magnitude Cretan detachment and both detachments started moving in the early Miocene. Ring et al. also discuss aspects of shortening- and extension-induced normal faults and conclude that the inception of the Amorgos detachment did not result from large-scale horizontal extension of the lithosphere. Seward et al. use a suite of new fission-track ages for revealing the timing of cooling/reheating of the metamorphic sequences on Naxos followed by cooling due to exhumation associated with extension along low-angle shear zones. They show that the late Miocene Naxos granodiorite was intruded into crustal levels where the temperature was already 300 8C as evidenced by the fissiontrack ages on zircon that are comparable to those obtained from higher temperature chronometers. Differential cooling of the pluton, identified through apatite fission-track ages, is related to either variable depth of emplacement and/or to

progressive tectonic exhumation accommodated first by ductile shallow-shear zone and then by high-angle cataclastic faults. Dilek et al. present new geochemical and isotopic data from the syn-extensional granitoid intrusions in the Menderes core complex in western Anatolia, and discuss their petrogenetic evolution in comparison to other late Cenozoic plutons in the Aegean region. Various degrees of partial melting of a mixed source including an enriched mantle lithosphere component and assimilatedmelted lower and middle crustal components were responsible for the production of the magmas of these mostly metaluminous, high-K calc-alkaline granitoids. Partial melting of the subductionmetasomatized lithospheric mantle and the overlying crust are interpreted to be triggered by asthenospheric upwelling caused by lithospheric delamination. The authors argue that both slab retreat generated upper plate deformation and magmatically induced crustal weakening played a major role in the onset of the Aegean extension in the early Miocene. Korja et al. describe a Precambrian example of a laterally spreading orogen with the help of largescale seismic reflection surveys, structural field work and existing geological/geophysical data. As a result of the extensional deformation the exposed bedrock is characterized by networks of orthogonal shear zones and low angle fabrics typically found in deep, lower level sections of core complexes. The authors argue that decoupling of the upper, middle and lower crust during extension resulted in the formation of layered superstructureinfrastructure of the crust. The editors are grateful to the Geological Society of London for making the publication of this book possible and to Angharad Hills, Commissioning Editor of the Geological Society Publishing House, and Jonathan Turner as the Society Books Editor for their continued support and patience. We also wish to thank the following reviewers for helping to improve the quality of this book: Gary Axen, John Bradshaw, Claire Curry, George Davis, Greg Davis, Mary Ford, Kevin Furlong, Klaus Gessner, Jamshid Hassanzadeh, Steve Kidder, Geoffroy Lamarche, Graham Leslie, Georgia Pe-Piper, Christine Siddoway, Christopher Talbot, Brian Taylor, Stuart Thomson. UWE RING & BRIAN WERNICKE

Contents Preface

vii

WERNICKE , B. The detachment era (1977–1982) and its role in revolutionizing continental tectonics

1

BOULTON , C., DAVIES , T. & MC SAVENEY , M. The frictional strength of granular fault gouge: application of theory to the mechanics of low-angle normal faults

9

MAROTTA , A. M., SPALLA , M. I. & GOSSO , G. Upper and lower crustal evolution during lithospheric extension: numerical modelling and natural footprints from the European Alps

33

STERN , T. A. Reconciling short- and long-term measures of extension in continental back arcs: heat flux, crustal structure and rotations within central North Island, New Zealand

73

TULLOCH , A. J., RAMEZANI , J., MORTIMER , N., MORTENSEN , J., VAN DEN BOGAARD , P. & MAAS , R. Cretaceous felsic volcanism in New Zealand and Lord Howe Rise (Zealandia) as a precursor to final Gondwana break-up

89

SKOURTSOS , E. & KRANIS , H. Structure and evolution of the western Corinth Rift, through new field data from the Northern Peloponnesus

119

THOMSON , S. N., RING , U., BRICHAU , S., GLODNY , J. & WILL , T. M. Timing and nature of formation of the Ios metamorphic core complex, southern Cyclades, Greece

139

RING , U., THOMSON , S. N. & ROSENBAUM , G. Timing of the Amorgos detachment system and implications for detachment faulting in the southern Aegean Sea, Greece

169

SEWARD , D., VANDERHAEGHE , O., SIEBENALLER , L., THOMSON , S., HIBSCH , C., ZINGG , A., HOLZNER , P., RING , U. & DUCHEˆ NE , S. Cenozoic tectonic evolution of Naxos Island through a multi-faceted approach of fission-track analysis

179

¨ NER , Z. Syn-extensional granitoids in the Menderes core DILEK , Y., ALTUNKAYNAK , S¸. & O complex and the late Cenozoic extensional tectonics of the Aegean province

197

KORJA , A., KOSUNEN , P. & HEIKKINEN , P. A case study of lateral spreading: the Precambrian Svecofennian Orogen

225

Index

253

The detachment era (1977 – 1982) and its role in revolutionizing continental tectonics B. WERNICKE California Institute of Technology, Division of Geological and Planetary Sciences, 1200 East California Boulevard, Pasadena, CA 91125, USA (e-mail: [email protected]) Abstract: After the discovery of thrust-and-nappe structure near the turn of the twentieth century, mountain belts were viewed as a direct expression of horizontal shortening of the continental crust, and continental rifting was viewed as a phenomenon distinct from it. By mid-century, broad consensus had emerged, mainly on the basis of physical reasoning, that thrust-and-nappe structure instead reflected gravity sliding secondary to vertical motions of the crust, as embodied in the influential stockwerk folding hypothesis. In a noteworthy period from 1977 to 1982, informally referred to here as the ‘detachment era’, not only did the last vestiges of support for the stockwerk hypothesis evaporate, but large-magnitude extension was discovered throughout the Cordillera, manifest primarily by extensional detachments and metamorphic core complexes. Soon afterward extensional detachments were recognized as a global phenomenon, forcing first-order reinterpretation of field relations in most orogens. Although plate tectonics is indisputably the most profound discovery in Earth sciences in the twentieth century, the detachment era arguably had commensurate impact on field-based interpretations of continental tectonics. Three decades later, controversy persists over the origin of metamorphic tectonites in core complexes, and over the existence and mechanics of slip on shallowly dipping extensional detachments.

The discovery and characterization of continental extension, as expressed by extensional detachments and metamorphic core complexes, occurred over a remarkably brief period referred to here as the ‘detachment era’. Its beginning is marked by a May 1977 Penrose Conference held near Tucson, Arizona, that first identified ‘metamorphic core complexes’. Its end is marked more diffusely, but by 1982 the interpretation of core complex detachments as large-displacement normal faults had become common if not universally accepted. Prior to this period, understanding of extensional tectonics evolved far more slowly than understanding of contractile tectonics. Ever since the discovery of thrust-and-nappe structure in the late nineteenth and early twentieth centuries in the Alps, Scotland and Scandinavia, the notion existed that largedisplacement faults (at least tens of kilometres) accommodated contractile strain of the crust (e.g. Ampferer & Hammer 1911). By mid-century, strike-slip faults such as the San Andreas in California were also discovered to have very large displacements (e.g. Hill & Dibblee 1953). The discovery of these elements guided a number of geologists toward mobilistic syntheses of crustal deformation in which mountain belts were regarded as a direct expression of large-scale horizontal motion between continental platforms (e.g. Argand 1916; Carey 1958).

These syntheses included the concept of finite horizontal stretching of the crust (e.g. Argand 1924), perhaps of magnitude sufficient to exhume ductilely stretched lower crust and eventually the upper mantle (e.g. Wegener 1929, chapter 10). Even though features now known to be detachments and core complexes had locally been described even earlier in the century (e.g. Ransome et al. 1910), making the connection between the stretching predicted by the mobilists and its expression in continental geology would have to wait more than half a century.

The fixist era By the advent of plate tectonics in 1968, documentation of large-displacement thrust faults and strikeslip faults in most orogenic belts was extensive. Remarkably, however, few prominent geologists at the time, even including the minority that had earlier advocated continental drift, believed that thrust faults were an expression of crustal contraction, and there was little notion that large-displacement normal faults even existed. Studies of exposed continental rifts were undertaken separately from the study of orogenic belts, and focused on conspicuous regions of active faulting and subsidence such as the Dead Sea rift (e.g. Quennell 1958) and the Basin

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 1–8. DOI: 10.1144/SP321.1 0305-8719/09/$15.00 # The Geological Society of London 2009.

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and Range province (e.g. Thompson 1960). On cross-sections of these rifts, normal faults were generally depicted as steeply dipping structures with displacements of no more than a few kilometres, largely in the vertical component of slip. In light of the thematic focus of many of the papers in this volume – large-displacement normal faults within orogens – at present the pendulum could hardly have swung any farther away from the mid-1960s consensus. In that consensus, the most widely accepted general model for mountain building was the ‘infrastructure-superstructure concept’, or as it was also known, the stockwerk folding hypothesis, in which the development of major tectonic elements results primarily from vertical motions within the crust (Wegmann 1935). According to this hypothesis, the evolution of orogenic belts begins with a long phase of subsidence and sedimentation of cryptic origin, followed by the buoyant rise of a ‘migmatite front’ toward the cold, strong sedimentary cover (superstructure). Finally, the metamorphic complex (infrastructure) incorporates and deforms part of the lower superstructure by a modest amount of horizontal flow along a complex shear zone (abscherungzone). The most cited example in support of this hypothesis was the East Greenland Caledonides, which exhibit spectacular exposures of a superstructure of thick Neoproterozoic –Cambrian sediments sheared off of, and locally incorporated into, their migmatitic substrate. During this era, the upward flux of heat and resulting buoyancy of a ‘granitized’ core was widely regarded as the driving force of tectonics. The model was based partly on the work of C. E. Wegmann and John Haller’s East Greenland studies, which included detailed mapping and spectacular photodocumentation of key structures (synthesized in English in Haller 1971). But its mainstream acceptance was driven mainly by the physical presumption that thrust sheets are internally too weak to be ‘pushed from behind’, leaving downslope movement under the influence of gravitational body forces as the only viable explanation for thrustand-nappe structure. The 1965 edition of the most influential geology text of the twentieth century (Holmes 1965), whose author was an early exponent of continental drift and mantle convection, somewhat ironically argues that horizontal shortening in thrust belts is a surficial response to vertical motions (e.g. Van Bemmelen 1954) and rejects the ‘traditional’ view that Alpine-type thrust-and-nappe structure is an expression of convergence between continental platforms on either side of the orogen. Holmes (1965) quotes Harold Jeffreys (p. 1169), an influential founder of the field of geophysics and (also ironically) a proponent of

the contracting Earth theory, as stating that the origin of thrust-and-nappe structure is ‘something that is not crustal shortening’. The stockwerk hypothesis for East Greenland did not survive geochronological testing, which demonstrated that high-temperature deformation of the deepest exposed infrastructural levels predated deposition of the superstructural sediments by more than a billion years (e.g. Henriksen & Higgins 1976). Modern studies of this complex show that rocks that did experience Caledonian metamorphism and deformation are regionally underlain by Archaean and Proterozoic basement with relatively minor Caledonian overprint, and that the abscherunzone is an extensional detachment (e.g. White & Hodges 2002).

The demise of fixism The geochronology of East Greenland notwithstanding, the end of the fixist synthesis of mountain building came not with the advent of plate tectonics even though it confirmed the mobility of lithospheric blocks on either side of mountain belts. Rather, it came from seismic images of crustalscale faults within mountain belts. It began with publication of petroleum industry seismic reflection profiles across the southern Canadian Rockies in the mid-1960s suggesting that the Rocky Mountains foreland fold-and-thrust belt was rooted beneath the crystalline core to the west (Bally et al. 1966), leading many geologists to seriously question the fixist synthesis (e.g. Burchfiel & Davis 1968). Despite these important developments, enthusiasm for vertical tectonics and gravity sliding remained strong through the 1970s in both North America and Europe (e.g. De Jong & Scholton 1973). The coup de grace was administered by the Consortium for Continental Reflection Profiling (COCORP), upon publication of deep seismic reflection profiles across the crystalline cores of two major contractile mountain ranges in the US that vindicated the general conclusions of Bally et al. (1966). A profile across the Laramide Wind River Range in Wyoming demonstrated that its bounding thrust fault cut at moderate to shallow dip to .20 km depth in the crust, and accommodated .20 km of horizontal shortening of the continental basement (Smithson et al. 1978). The profile unambiguously resolved a long-standing debate as to whether or not the Laramide ranges were vertical ‘piston-cored’ uplifts, over which the sedimentary cover had been ‘draped’. A second profile across the southern Appalachian orogen showed that its crystalline core was a detached upper crustal flake that had overthrust the ancient continental platform by at least 260 km (Cook et al. 1979), and so could

INTRODUCTION

not be considered any sort of deep-seated ‘infrastucture’ for the Valley and Ridge province fold and thrust belt. The publication of the COCORP profiles accordingly marked the end of serious speculation that either of the imaged faults, or thrust-and-nappe structure in general, were primarily the result of vertical movements. Despite its demise, the stockwerk folding hypothesis did contain the kernel of the idea that the lower continental crust is hotter and therefore may be weaker than the upper crust, and hence subject to lateral flow, as also anticipated by the early mobilists. During the detachment era (again, coincidentally), this qualitative idea was confirmed and quantified by the synthesis and publication of laboratory data on the brittle and ductile failure strengths of rocks (Brace & Kohlstedt 1980). By 1978, when the author entered graduate school, the kinematic relationship between upper crustal strain via faulting and deep crustal strain via flow mechanisms was a much-debated subject for strike-slip, thrust and normal fault systems, about which there was, and still remains, much controversy.

Impact of the 1977 Penrose Conference near Tucson, Arizona Even with the demise of fixism and acceptance of continental rifting and sea-floor spreading, the consensus view remained that continental extension was accommodated by steep normal faults, and hence no field evidence or mechanism was known for the mechanical extension of the upper crust much beyond 10% increase in original width (Stewart 1971; Thompson & Burke 1974). Further, it was tacitly assumed by most geologists that the formation and exposure of metamorphic rocks occurred in contractile regimes. However, a handful of field and geochronological studies in the Basin and Range province (Anderson 1971; Armstrong 1972; Wright & Troxel 1973; Proffett 1977), having little to do with the plate tectonics revolution, quietly began to mount a strong challenge to both of these views of continental tectonics. Put into practice for Basin and Range field relations, the first implied that all low-angle faults, either on geological maps or in subsurface images, were thrust faults. The second implied that all Phanerozoic metamorphism in the Cordillera occurred during either Palaeozoic or Mesozoic episodes of horizontal contraction. Because thrust faulting in the retroarc Cordilleran fold-and-thrust belt ended by Early Tertiary time, major low-angle faults now recognized as Tertiary extensional detachments in the Basin and Range were initially interpreted as either Mesozoic thrusts (e.g. Thorman 1970; Drewes 1978) or alternatively as gravity slides of

3

Mesozoic (Hose & Danes 1973) or Tertiary age (e.g. Compton et al. 1977). As a typical example, in the Snake Range area of east-central Nevada, a number of large low-angle faults are developed within thick continental shelf deposits of Palaeozoic age that consistently place younger rocks on top of older. Among the best examples of these faults is the Snake Range decollement, which juxtaposes steeply tilted strata as young as mid-Tertiary over Cambrian strata metamorphosed to the amphibolite facies. The fault was originally discovered by Misch (1960) and interpreted to be a Mesozoic thrust fault. Armstrong’s (1972) geometric analyses of relationships between Tertiary strata and a number of these faults showed that virtually none of the large-offset low-angle faults in the region could be Mesozoic in age, and that the consistency of younger-on-older relationships were best attributed in some way to extension. (A detailed analysis of this landmark paper may be found in Wernicke & Spencer 1999.) The 1977 Penrose Conference convened near Tucson, Arizona (Max D. Crittenden, Peter J. Coney and George H. Davis, convenors) marks the beginning of the detachment era. It was attended by a large fraction of the geologists who over the ensuing five years would field-test the premise of the conference: that the Basin and Range was literally awash in both metamorphic tectonite and large faults that post-dated regional contraction of the continental interior. With the noteworthy exception of the mobilistic synthesis of Hamilton & Myers (1966), prior to this era few geologists believed that Cordilleran extension amounted to anything other than modest horizontal stretching along moderate to steeply dipping high-angle faults. This view was not substantially different from structural sections drawn near the turn of the century (e.g. the much-reproduced fig. 1 in Davis 1903), except perhaps for quantification of horizontal extension as greater than zero, but relatively small. By 1982, it was clear on the basis of field relations and geochronology that a profound mid-Tertiary period of extension resulted in the development of detachments with displacements large enough to exhume broad tracts of mid-crustal metamorphic tectonite and associated magmas in their footwalls, and that the net horizontal extension across the province was probably much larger than 10% (e.g. Davis & Coney 1979; Crittenden et al. 1980; Reynolds & Rehrig 1980; Wernicke 1981; Armstrong 1982; Frost & Martin 1982). The generic template for core complexes that quickly developed was that of a domiform fault surface separating footwall metamorphic tectonites (or in some cases, unmetamophosed, deeper crustal levels) from distended hanging-wall rocks. In this template (e.g. Davis 1980), the maximum

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elongation direction in the tectonites trends parallel to antiformal axes defined by the detachment, spectacularly exposed in ranges such as the Snake Range, Whipple Mountains of southeastern California, and Catalina-Rincon Mountains of southeastern Arizona, the latter of which was visited and examined by participants of the 1977 conference. The maximum elongation direction in footwall tectonites also tends to parallel the maximum elongation direction of hanging-wall normal fault blocks. Footwall tectonites are often syn-tectonically intruded by granitic magmas, and exhibit progressive overprinting of higher temperature deformation fabrics by lower. Hanging-wall fault blocks are typically steeply tilted, and contain mafic to silicic volcanics interstratified with fluvial–lacustrine sediments and scarp-induced rock avalanche deposits. The export of this generic template from the Cordillera to the rest of the world was swift. It is so distinctive that the presence of core complexes often became apparent simply by reading existing literature. For example, the first core complexes identified in the Alpine– Himalayan chain were in a previously well studied portion of the Cyclades Islands in the Aegean Sea. These core complexes were identified primarily by review of existing structural and geochronological data, and as was the case in the Cordillera, decades of interpreting low-angle structures in the Aegean region as thrust faults came to an abrupt end (Lister et al. 1984). Comparable revolutions also soon followed in the Alps (e.g. Selverstone 1988), Himalaya–Tibet (e.g. Burchfiel & Royden 1985), Caledonides (e.g. Serrane & Siguret 1987), and for accretionary settings in general (Platt 1986). Core complex elements were also soon identified in passive margin settings (Wernicke 1985; Lister et al. 1986) and in slow-spreading mid-ocean ridges (e.g. Karson & Dick 1983).

Uniform stretching and its exceptions During the detachment era, a parallel revolution occurred in understanding the physical relationship between extension and vertical motions of the crust. The key breakthrough came from consideration of the isostatic consequences of mechanical extension and ensuing conductive cooling of newly thinned lithosphere. In general, if ‘typical’ continental lithosphere is stretched uniformly, it will lose elevation, because the negative effect of thinning the crust outweighs the positive effect of thinning the subcrustal lithosphere. After mechanical stretching is over, conductive cooling of the boundary layer predicts elevation loss at a rate proportional to the square root of time, analogous to

oceanic lithosphere as it moves away from spreading centres (McKenzie 1978). Adding to this framework the effect of sedimentary compaction and flexural strength, relatively simple physical models can be developed that predict the subsidence rate in sedimentary basins produced by a given amount of crustal stretching, most usefully in passive margin sedimentary wedges (e.g. Sclater & Christie 1980; Watts et al. 1982). Because of the increasing availability of seismic reflection profiles and well data from petroleum exploration of passive margins, literally hundreds of studies attempting to apply the uniform stretching model to sedimentary basins were published during and after the detachment era. In addition to the powerful ‘exportability’ of their distinctive generic template, the discovery of core complexes in the Cordillera also led to the quantification of the amount of horizontal stretching of the upper crust in many parts of the Basin and Range. It was found that stretching was extremely heterogeneous, such that the province was a complicated patchwork of very highly extended regions set among relatively unextended crustal blocks (e.g. Guth 1981; Miller et al. 1983). It was also realised that although the degree of upper crustal extension varied many-fold, there was little or no commensurate variation in crustal thickness, falsifying the assumption of uniform stretching in dramatic fashion (Wernicke 1985). The simplest way to maintain a flat Moho is to treat the deep crust as a fluid or ‘crustal asthenosphere’ within which the upper crustal fault blocks float, a simple tectonic idea dating back to Taber (1927). In the case of the complex three-dimensional patchwork of highly extended areas in the Basin and Range, flow within an inviscid substrate would precisely complement upper crustal extension so as to minimize relief on the Moho (Block & Royden 1990). Other areas of failure of the uniform stretching model were revealed in detailed analyses of uplift and subsidence patterns within and adjacent to passive margin sedimentary basins, which led many workers to propose a variety of non-uniform stretching models such as simple shear of the whole lithosphere (Wernicke 1985; Ussami et al. 1986) or combinations of simple shear and pure shear (e.g. Kusznir & Ziegler 1992). The notion of an intracrustal asthenosphere initially developed for extensional terrains has since been applied to collisional regimes with thick crust (e.g. Clark & Royden 2000; McQuarrie & Chase 2000). Culshaw et al. (2006) have gone so far as to call for a ‘quantitative revival’ of the infrastructure–superstructure concept, although these authors seemed unaware that the original concept was developed to explain mountain building without resort to horizontal strain of the crust as outlined above.

INTRODUCTION

The way ahead The breakthroughs in extensional tectonics during the detachment era resulted primarily from professionally courageous interpretations of field relations that strongly contradicted community consensus, and from the advent of geochronological studies of crystalline rocks in core complex footwalls, which demonstrated their mid-Tertiary age. They occurred in the context of a wider consensus that all low-angle faults were either thrust faults or surficial gravity slides, normal faults were of limited displacement and restricted to continental rifts, and metamorphic tectonites form in contractile tectonic settings. The wholesale demolition of this consensus during the detachment era was profoundly transformative of continental tectonics, arguably the single most important advance in the post-plate tectonics era. (A geophysical perspective that instead emphasizes active tectonics and various continuum theories may be found in Molnar 2001.) The objective of identifying and describing this transformation here is to highlight the importance of scepticism in moving the field ahead. In contrast to most community endeavours where the objective is to forge consensus (e.g. in religious, political or family matters), in science the highest calling of community members is to disembowel it, the wider the better. Further progress in understanding extensional detachments and metamorphic core complexes should therefore focus on scepticism of major hypotheses that in one way or another have weaknesses. A top candidate for scrutiny is the notion that metamorphic tectonites in core complexes represent the down-dip continuation of the brittle shear plane of the overlying detachment fault (e.g. Wernicke 1981). In most instances, the situation appears to be more complex, to the point that it is questionable whether even a small fraction of footwall tectonite has this origin. Flow within a relatively inviscid fluid layer is a distinct alternative. However, as is clear from several papers in this volume, the formation of footwall tectonites in many instances occurs well before displacement on the overlying detachments, suggesting more complex mechanisms for the development and ‘capture’ of tectonites by the fault may be in play. Improved constraints on the timing of tectonite formation, as amply represented by the research published in this volume, will likely lead to new and surprising conclusions regarding the origin of footwall tectonite. These terrains are clearly the most important laboratories for testing models of the kinematic relationship between upper crustal and deeper crustal layers during tectonism, and the lack of consensus over how they evolve makes them all the more interesting.

5

An equally important candidate that merits scepticism is the concept that some detachments are active at dips of less than 308, as suggested by seismic reflection profiles (most notably, the COCORP profile across the Sevier Desert basin in the Utah Basin and Range; Allmendinger et al. 1983) and field relations between detachments and hanging wall fault blocks exposed in a number of core complexes (e.g. Lister & Davis 1989). The close of the detachment era might also be marked by the emergence of a spirited, contrarian minority arguing that even though the extension is large, core complex detachments are not large-displacement faults (e.g. Miller et al. 1983). Although most geologists today, including the author, remain comfortable drawing reconstructions that include low-angle normal faults with large displacements, over the last quarter century a ‘loyal opposition’ has persistently challenged evidence in support of active slip on low-angle planes (e.g. Jackson & White 1989; Anders & Christie-Blick 1994; Anders et al. 2006; Wong & Gans 2008). These authors have properly stressed that an inventory of earthquake focal mechanisms on a par with that of low-angle thrust faults, strike-slip faults, and moderate- to high-angle normal faults has yet to materialize, and that slip on low-angle normal faults is in any event extremely difficult mechanically (e.g. Wills & Buck 1997). A novel modelling approach that accounts for thermal–mechanical feedbacks between brittle and viscous layers demonstrates that elastic strain energy becomes focused into the strongest layers of the lithosphere, producing primary low-angle normal faults without any special assumptions or anisotropy (e.g. Weinberg et al. 2007), and thus may explain the mechanics. Nonetheless, after many decades of seismic monitoring in regions of active extension, the community still awaits ‘the big one’ on an active low-angle normal fault. The controversy, and its implications for the mechanics of brittle faulting in the crust, has recently stimulated international interest in the establishment of a borehole observatory through the very low-angle (10 –128) detachment imaged beneath the Sevier Desert basin (Christie-Blick et al. 2007), which geodetic data suggest is currently active (Niemi et al. 2004). The establishment of a borehole observatory across this fault is particularly timely in that the mechanical paradox of brittle slip on fault planes oriented at high angle to the maximum principal stress direction is common to both the Sevier Desert detachment and the San Andreas fault. The latter is developed within a transpressional tectonic regime and is currently being monitored by the San Andreas Borehole Observatory at Depth (SAFOD). Preliminary data suggest that the San Andreas is a weak zone of

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high pore-fluid pressure and strongly variable permeability, embedded within an otherwise much stronger crust (Hickman et al. 2008), along the lines suggested for the San Andreas by Rice (1992) and extended to low-angle normal faults by Axen (1992). Verifying that the Sevier Desert detachment is active, determining the orientations and magnitudes of the stresses within and adjacent to the fault zone and characterizing the pore fluid pressure and composition of fault rocks will provide significant new insight into fault mechanics. In particular, it will provide an opportunity to determine whether the physical processes that enable slip on the San Andreas are the same as those in the radically different tectonic setting of the Basin and Range. Preparation of this account was supported by the Caltech Tectonics Observatory and NSF Grant EAR-0810328. Comments by Gregory A. Davis, George H. Davis and Uwe Ring greatly improved the balance and clarity of the manuscript.

References A MPFERER , O. & H AMMER , W. 1911. Geologischer querschnitt durch die Ostalpen. Jahrbuch Geologische Reichsanst, 61, 531– 711. A NDERS , M. H. & C HRISTIE -B LICK , N. 1994. Is the Sevier Desert reflection of west-central Utah a normal fault. Geology, 22, 771–774. A NDERS , M. H., C HRISTIE -B LICK , N. & W ALKER , C. D. 2006. Distinguishing between rooted and rootless detachments: a case study from the Mormon Mountains of southeastern Nevada. Journal of Geology, 114, 645– 664. A NDERSON , R. E. 1971. Thin-skin distension in Tertiary rocks of southeastern Nevada. Geological Society of America Bulletin, 82, 43–58. A RGAND , E. 1916. Sur l’arc des Alpes Occidentales. Eclogae Geologicae Helvetiae, 14, 145– 191. A RGAND , E. 1924. Des Alpes et de l’Afrique. Bulletin de la Socie´te´ vaudoise des sciences naturelles, 55, 233– 236. A RMSTRONG , R. L. 1972. Low-angle (denudation) faults, hinterland of Sevier orogenic belt, Eastern Nevada and western Utah. Geological Society of America Bulletin, 83, 1729– 1754. A RMSTRONG , R. L. 1982. Cordilleran metamorphic core complexes – from Arizona to southern Canada. Annual Review of Earth and Planetary Sciences, 10, 129– 154. A XEN , G. J. 1992. Pore pressure, stress increase, and fault weakening in low-angle normal faulting. Journal of Geophysical Research– Solid Earth, 97, 8979–8991. B ALLY , A. W., G ORDY , P. L. & S TEWART , G. A. 1966. Structure, seismic data, and orogenic evolution of southern Canadian Rocky Mountains. Bulletin of Canadian Petroleum Geology, 14, 337–381. B LOCK , L. & R OYDEN , L. H. 1990. Core complex geometries and regional scale flow in the lower crust. Tectonics, 9, 557–567.

B RACE , W. F. & K OHLSTEDT , D. L. 1980. Limits on lithospheric stress imposed by laboratory experiments. Journal of Geophysical Research, 85, 6248– 6252. B URCHFIEL , B. C. & D AVIS , G. A. 1968. Two-sided nature of the Cordilleran orogen and its tectonic implications. International Geological Congress, 23rd Proceeding, 3, 175–184. B URCHFIEL , B. C. & R OYDEN , L. H. 1985. North–south extension within the convergent Himalayan region. Geology, 13, 679 –682. C AREY , S. W. 1958. A tectonic approach to continental drift. In: C AREY , S. W. (ed.) Continental Drift: A Symposium. University of Tasmania, Hobart, 117–355. C HRISTIE -B LICK , N., A NDERS , M. H., D RESEN , G., L ISTER , G. S., M ANATSCHAL , G. & W ERNICKE , B. P. 2007. Testing the extensional detachment paradigm: scientific drilling in the Sevier Desert basin, western United States. Geological Society of America Abstracts with Programs, 39, 227– 227. C LARK , M. K. & R OYDEN , L. H. 2000. Topographic ooze: building the eastern margin of Tibet by lower crustal flow. Geology, 28, 703– 706. C OMPTON , R. R., T ODD , V. R., Z ARTMAN , R. E. & N AESER , C. W. 1977. Oligocene and Miocene metamorphism, folding, and low-angle faulting in northwestern Utah. Geological Society of America Bulletin, 88, 1237–1250. C OOK , F. A., A LBAUGH , D. S., B ROWN , L. D., K AUFMAN , S., O LIVER , J. E. & H ATCHER , R. D. 1979. Thin-skinned tectonics in the crystalline southern Appalachians – COCORP seismic-reflection profiling of the Blue Ridge and Piedmont. Geology, 7, 563– 567. C RITTENDEN , M. D., C ONEY , P. J. & D AVIS , G. H. (eds) 1980. Cordilleran Metamophic Core Complexes. Geological Society of America, Memoirs, 153. C ULSHAW , N. G., B EAUMONT , C. & J AMIESON , R. A. 2006. The orogenic superstructure–infrastructure concept: revisited, quantified, and revived. Geology, 34, 733–736. D AVIS , G. H. 1980. Structural characteristics of metamorphic core complexes, southern Arizona. In: C RITTENDEN , M. D., C ONEY , P. J. & D AVIS , G. H. (eds) Cordilleran Metamorphic Core Complexes. Geological Society of America Memoirs, 153, 35–77. D AVIS , G. H. & C ONEY , P. J. 1979. Geologic development of the Cordilleran metamorphic core complexes. Geology, 7, 120– 124. D AVIS , W. M. 1903. The mountain ranges of the Great Basin. Bulletin of the Museum of Comparative Zoology, 42, 129– 177. D E J ONG , K. A. & S CHOLTON , R. 1973. Gravity and Tectonics. Wiley-Interscience, New York. D REWES , H. 1978. Cordilleran orogenic belt between Nevada and Chihuahua. Geological Society of America Bulletin, 89, 641– 657. F ROST , E. G. & M ARTIN , D. L. (eds) 1982. Mesozoic– Cenozoic Evolution of the Colorado River Region, California, Arizona, and Nevada (Anderson–Hamilton Volume). Cordilleran Publishers, San Diego. G UTH , P. L. 1981. Tertiary extension north of the Las Vegas Valley shear zone, Sheep and Desert Ranges, Clark County, Nevada. Geological Society of America Bulletin, 92, 763– 771.

INTRODUCTION H ALLER , J. 1971. Geology of the East Greenland Caledonides. Interscience, London. H AMILTON , W. & M YERS , W. B. 1966. Cenozoic Tectonics of Western United States. Reviews of Geophysics, 4, 509– 549. H ENRIKSEN , N. & H IGGINS , A. K. 1976. East Greenland Caledonian fold belt. In: E SCHER , A. & W ATT , W. S. (eds) Geology of Greenland. Geological Survey of Greenland, Copenhagen, 183– 246. H ICKMAN , S. H., Z OBACK , M. ET AL . 2008. Structure and composition of the San Andreas fault in central California: recent results from SAFOD sample analyses. EOS Transactions of AGU, 89(53), T53F-01. H ILL , M. L. & D IBBLEE , T. W. 1953. San Andreas, Garlock, and Big Pine Faults, California – a study of the character, history, and tectonic significance of their displacements. Geological Society of America Bulletin, 64, 443–458. H OLMES , A. 1965. Principles of Physical Geology (2nd edn). The Ronald Press Company, New York. H OSE , R. K. & D ANES , Z. F. 1973. Development of the Late Mesozoic to Early Cenozoic structures of the eastern Great Basin. In: DE J ONG , K. A. & S HOLTON , R. (eds) Gravity and Tectonics. WileyInterscience, New York, 429– 441. J ACKSON , J. A. & W HITE , N. J. 1989. Normal faulting in the upper continental crust – observations from regions of active extension. Journal of Structural Geology, 11, 15–36. K ARSON , J. A. & D ICK , H. J. B. 1983. Tectonics of ridge– transform intersections at the Kane fracture zone. Marine Geophysical Researches, 6, 51– 98. K USZNIR , N. J. & Z IEGLER , P. A. 1992. The mechanics of continental extension and sedimentary basin formation – a simple-shear pure-shear flexural cantilever model. Tectonophysics, 215, 117–131. L ISTER , G. S., B ANGA , G. & F EENSTRA , A. 1984. Metamorphic core complexes of Cordilleran type in the Cyclades, Aegean Sea, Greece. Geology, 12, 221–225. L ISTER , G. S. & D AVIS , G. A. 1989. The origin of metamorphic core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, USA. Journal of Structural Geology, 11, 65–94. L ISTER , G. S., E THERIDGE , M. A. & S YMONDS , P. A. 1986. Detachment faulting and the evolution of passive continental margins. Geology, 14, 246 –250. M C K ENZIE , D. 1978. Some remarks on the development of sedimentary basins. Earth and Planetary Science Letters, 40, 25–32. M C Q UARRIE , N. & C HASE , C. G. 2000. Raising the Colorado Plateau. Geology, 28, 91–94. M ILLER , E. L., G ANS , P. B. & G ARING , J. 1983. The Snake Range decollement – an exhumed mid-Tertiary ductile–brittle transition. Tectonics, 2, 239–263. M ISCH , P. 1960. Regional structural reconnaissance in central-northeast Nevada and some adjacent areas – observations and interpretations. In: Intermountain Association of Petroleum Geologists 11th Annual Field Conference Guidebook, 17–42. M OLNAR , P. 2001. From plate tectonics to continental tectonics. In: O RESKES , N. & L E G RAND , H. (eds) Plate

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Tectonics: An Insiders’s History of the Modern Theory of the Earth. Westview Press, Boulder. N IEMI , N. A. ET AL . 2004. BARGEN continuous GPS data across the eastern Basin and Range province, and implications for fault system dynamics. Geophysical Journal International, 159, 842– 862; doi:10.1111/ j.1365-246X.2004.02454.x. P LATT , J. P. 1986. Dynamics of orogenic wedges and the uplift of high-pressure metamorphic rocks. Geological Society of America Bulletin, 97, 1037–1053. P ROFFETT , J. M. 1977. Cenozoic geology of Yerington District, Nevada, and implications for nature and origin of Basin and Range faulting. Geological Society of America Bulletin, 88, 247– 266. Q UENNELL , A. M. 1958. The structural and geomorphic evolution of the Dead Sea rift. Quarterly Journal of the Geological Society of London, 114, 1–24. R ANSOME , F. L. ET AL . 1910. Geology and ore deposits of the Bullfrog District, Nevada. U.S. Geological Survey Bulletin, 407, 1– 130. R EYNOLDS , S. J. & R EHRIG , W. A. 1980. Mid-Tertiary plutonism and mylonitization, South Mountains, central Arizona. In: C RITTENDEN , M. D., C ONEY , P. J. & D AVIS , G. H. (eds) Cordilleran Metamophic Core Complexes. Geological Society of America Memoirs, 153, 159–176. R ICE , J. R. 1992. Fault stress states, pore pressure distributions, and the weakness of the San Andreas fault. In: W ONG , T. F. & E VANS , B. (eds) Fault Mechanics and Transport Properties of Rocks. Academic Press, London, 475–503. S CLATER , J. G. & C HRISTIE , P. A. F. 1980. Continental stretching – an explanation of the post-mid-Cretaceous subsidence of the central North-Sea basin. Journal of Geophysical Research, 85, 3711– 3739. S ELVERSTONE , J. 1988. Evidence for east–west crustal extension in the Eastern Alps – implications for the unroofing history of the Tauern window. Tectonics, 7, 87–105. S ERRANE , M. & S EGURET , M. 1987. The Devonian basins of western Norway: tectonics and kinematics of an extending crust. In: C OWARD , M. P., D EWEY , J. F. & H ANCOCK , P. L. (eds) Continental Extensional Tectonics. Geological Society, London, Special Publications, 28, 537–548. S MITHSON , S. B., B REWER , J., K AUFMAN , S., O LIVER , J. & H URICH , C. 1978. Nature of Wind River thrust, Wyoming, from COCORP deep-reflection data and from gravity data. Geology, 6, 648– 652. S TEWART , J. H. 1971. Basin and Range structure – system of horsts and grabens produced by deep-seated extension. Geological Society of America Bulletin, 82, 1019– 1044. T ABER , S. 1927. Fault troughs. Journal of Geology, 35, 577– 606. T HOMPSON , G. A. 1960. Problem of Late Cenozoic structure of the Basin Ranges. In: Proceedings of the 21st International Geological Congress, Copenhagen, 18, 62–68. T HOMPSON , G. A. & B URKE , D. B. 1974. Regional geophysics of the Basin and Range province. Annual Review of Earth and Planetary Sciences, 2, 213– 238. T HORMAN , C. H. 1970. Metamorphosed and nonmetamorphosed Paleozoic rocks in Wood-Hills and

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Pequop-Mountains, northeast Nevada. Geological Society of America Bulletin, 81, 2417–2447. U SSAMI , N., K ARNER , G. D. & B OTT , M. H. P. 1986. Crustal detachment during South Atlantic rifting and formation of Tucano–Gabon basin system. Nature, 322, 629– 632. V AN B EMMELEN , R. W. 1954. Mountain Building. Martinus Nijhoff, The Hague. W ATTS , A. B., K ARNER , G. D. & S TECKLER , M. S. 1982. Lithospheric flexure and the evolution of sedimentary basins. Philosophical Transactions of the Royal Society of London, A305, 249–281. W EGENER , A. 1929. Die Entstehung der Ozeane und Kontinente. 4th edition. Braunschwieg, Germany, 1 –246. W EGMANN , C. E. 1935. Zur deutung der migmatite. Geologische Rundschau, 26, 305– 350. W EINBERG , R. F., R EGENAUER -L IEB , K. & R OSENBAUM , G. 2007. Mantle detachment faults and the breakup of cold continental lithosphere. Geology, 35, 1035–1038. W ERNICKE , B. 1981. Low-angle normal faults in the Basin and Range province – nappe tectonics in an extending orogen. Nature, 291, 645– 648. W ERNICKE , B. 1985. Uniform-sense normal simple shear of the continental lithosphere. Canadian Journal of Earth Sciences, 22, 108– 125.

W ERNICKE , B. & S PENCER , J. 1999. Retrospective on ‘Low-angle (denudation) faults, hinterland of the Sevier orogenic belt, eastern Nevada and western Utah’ by Richard Lee Armstrong. In: M OORES , E. M., S LOAN , D. & S TOUT , D. L. (eds) Classic Cordilleran Concepts: A View from California. Geological Society of America Special Papers, 338, 357–362. W HITE , A. P. & H ODGES , K. V. 2002. Multistage extensional evolution of the central East Greenland Caledonides. Tectonics, 21, doi: 10:1029/2007TC001308. W ILLS , S. & B UCK , W. R. 1997. Stress-field rotation and rooted detachment faults: a Coulomb failure analysis. Journal of Geophysical Research–Solid Earth, 102, 20503–20514. W ONG , M. S. & G ANS , P. B. 2008. Geologic, structural, and thermochronologic constraints on the tectonic evolution of the Sierra Mazatan core complex, Sonora, Mexico: new insights into metamorphic core complex formation. Tectonics, 27, doi: 10:1029/ 2007TC002193. W RIGHT , L. A. & T ROXEL , B. W. 1973. Shallow-fault interpretation of Basin and Range structure, southwestern Great Basin. In: DE J ONG , K. A. & S CHOLTON , R. (eds) Gravity and Tectonics. Wiley-Interscience, New York, 397–407.

The frictional strength of granular fault gouge: application of theory to the mechanics of low-angle normal faults CAROLYN BOULTON1*, TIM DAVIES1 & MAURI MC SAVENEY2 1

University of Canterbury, New Zealand, Geological Sciences, Private Bag 4800, Christchurch, Canterbury 8140, New Zealand 2

GNS Science, PO Box 30368, Lower Hutt, New Zealand

*Corresponding author (e-mail: [email protected]) Abstract: There is a persistent body of literature that suggests low-angle normal faults (LANF) form and slip seismically; if true, the effective friction coefficient is much lower (,0.3) than that found in laboratory tests of rock friction (c. 0.8) and in low-displacement faults that lack well-developed fault cores. This paper summarizes and discusses the mechanisms proposed to explain the low apparent friction of crustal-scale faults with low resolved shear stresses. Emphasis is placed on differentiating static weakening mechanisms, operating at strain rates c. 10212 s21 – 10215 s21, from dynamic weakening mechanisms, operating at strain rates .1021 s21. Previous published explanations for low fault friction do not appear to meet the key requirements of (i) reducing both static and dynamic frictional strength of LANF and (ii) operating only along crustal-scale faults. Fault rock assemblages in quartzo-feldspathic continental crust reveal that grain size reduction, or comminution, plays a fundamental role in fault zone development. As a fault accrues displacement, a fault core forms that contains granular material. We postulate that dynamic rock fragmentation occurs during the shearing of confined granular material; dynamic fragmentation is a volume-dependent mechanism responsible for reducing the static and dynamic frictional strengths of faults.

Brittle normal faults accommodate crustal extension and commonly exhibit cumulative slips exceeding 10 km. If normal faulting took place with conventional dips (.308), then 10 km of extension on a single fault would necessarily be associated with .5 km of vertical structural relief. Some normal faults have extensions of 100 km or more, and typical extensions are 10–50 km, implying implausible relief generation (e.g. Ring et al. 2001; Axen 2004). Logically, large-extension normal faults should not occur. However, this problem is at least partly resolved by the fact that large displacement normal faults usually have low inclinations to the horizontal. Fault mechanics theory dictates that normal faults should slip at dips of no less than 308 to the horizontal, so a kinematic conundrum is replaced by a theoretical one (Anderson 1951; Byerlee 1978; Sibson 1983, 1994). Fundamentally, in theory and in practice, shear failures in triaxially-loaded rocks align at about 308 to the axis of maximum compressive stress in rock types with typical angles of internal friction of c. 30 –458. The vertical lithostatic load is the maximum principal stress in extensional tectonic settings, so normal faults should occur at an angle of c. 308 to 608 from vertical (Anderson 1951). Nevertheless, numerous examples of low-angle normal faults (LANF) dipping less than 308 occur

worldwide: in the Basin and Range province of North American (Lister & Davis 1989; Cowan et al. 2003); Greece (Forster & Lister 1999); the East African Rift System (Morley 1999); the Northern Appenines of Italy (Collettini & Barchi 2004; Smith et al. 2007) and West Iberia (Dean et al. 2008). Thermal, palaeomagnetic and structural constraints suggest that LANF initiated and were active at very low dips (e.g. Wernicke 1995). In the field, LANF truncate and post-date high-angle normal faults, and these structural relationships suggest a vertical principal compressive stress during the time of LANF formation (Reynolds & Lister 1987; Axen & Selverstone 1994). In order to form LANF, then, the friction coefficient of the rock must be less (and sometimes substantially less) than the laboratory (‘Byerlee’) value of about 0.6 to 0.85 (Byerlee 1978). Field investigations consistently measure Byerlee friction values in crustal rocks (e.g. McGarr & Gay 1978; Brudy et al. 1997), and historical records of seismicity have also failed to resolve the problem of LANF. Evidence gathered from the seismological record substantiates fault mechanics tenets, as no clearly resolved large magnitude (M . 5.5) earthquake has occurred on a normal-fault nodal plane dipping less than 308 (Jackson & White 1989; Colletini & Sibson 2001).

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 9–31. DOI: 10.1144/SP321.2 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Some convincing geological and seismic-reflection evidence exists showing that the 1954 Dixie Valley earthquake (Ms ¼ 6.8), Dixie Valley, Nevada, ruptured a low-angle normal fault dipping 25– 308 to a depth of 2.7 km. Fault-plane solutions for the Dixie Valley earthquake, however, are obscured by waveforms from the Fairview Peak earthquake (Ms ¼ 7.2), which occurred 4 min and 20 s earlier (Abbott et al. 2001). The lack of recorded, unambiguous, moderate to large ruptures on LANF has led to the deduction that LANF fail aseismically, microseimically or in large earthquakes with long recurrence intervals (Reitbrock et al. 1996; Collettini & Barchi 2004; Wernicke 1995). Although it appears theoretically unlikely, sufficient seismological and geological evidence exists to justify considering that LANF form at their present inclinations and remain seismically active (e.g. Wenicke 1995; Rigo et al. 1996; Abbott et al. 2001; Axen 2004). This is especially true when one considers that a number of other crustal-scale faults also appear to require low friction to explain the stress conditions under which they rupture (e.g. the Cascadia subduction fault, USA, Wang et al. 1995; the San Andreas fault, USA, Townend & Zoback 2004; the Marlborough fault zone, New Zealand, Balfour et al. 2005). This independent association of large-scale faulting with low friction lends some credibility by

association to the suggestion that LANF slip seismically under resolved stresses that necessitate low frictional strength both statically and dynamically.

Geology and occurrence of LANF A low-angle normal fault, or detachment fault, is a moderately to gently dipping fault that accommodates crustal extension. While gently dipping normal faults often appear in seismic reflection profiling of the upper few kilometres of sedimentary basins, these faults are commonly listric structures that sole into detachment horizons (for a review of the mechanisms generating fault curvature see Vendeville 1991 and White & Yielding 1991). In this paper, the term LANF applies to gently dipping (,308) planar fault surfaces of large areal extent (up to 200 km2) along which large magnitude slip (typically 10 –50 km) has taken place. These structures have been described by Coney (1980), Lister & Davis (1989), John & Foster (1993), Wernicke (1995), Axen (2004), and Dean et al. (2008) among others. Many LANF juxtapose unfaulted crystalline footwall rocks against faulted sedimentary and/or volcanic hanging-wall rocks (Lister & Davis 1989; John & Foster 1993) (Fig. 1). Ductile rocks in the footwall of LANF include mylonites and mylonitic gneisses developed under greenschist- to

Fig. 1. A diagrammatic representation of a low-angle normal fault showing a typical assemblage of ductile and brittle fault rocks. In cross section, LANF fault cores commonly containing a narrow (millimetre to centimetre) principal slip zone of highly comminuted ultracataclasite surrounded by finely comminuted cataclasite and gouge fragments (centimeters to metres wide). Fault core rocks have a fractal particle size distribution. Highly damaged and brecciated rock may extend into the hanging wall tens to hundres of meters away from the slip zone. Footwall rocks are commonly intact and comprise preserved mylonites formed in the ductile lower crust (after Handy et al. 2005).

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

amphibolite-facies metamorphic conditions. Goodwin (1999) documented intense comminution and pseudotachylyte (‘fossil earthquake’) formation on narrow shear surfaces in mylonitic LANF rocks, illustrating alternating brittle and ductile behaviour at the base of the seismogenic crust. At lower temperatures in the shallow crust, ductile fault rocks are overprinted by brittle deformation (Coney 1980). Field descriptions of brittle fault rocks commonly highlight the occurrence of finely comminuted cataclasite, fault gouge and fault breccia on narrow slip zones in the fault core (e.g. John & Forster 1993; Cladouhos 1999; Hayman 2006). The term fault core was defined by Caine et al. (1996) as ‘the structural, lithologic, and morphologic portion of the fault zone where most of the displacement is accommodated’. Low-angle normal faults form in various tectonic settings in the continental crust: in continental rift-related metamorphic core complexes (e.g. Crittenden et al. 1980; Lister & Davis 1989; Wernicke 1995; Axen 2004); in highly extended passive continental margins (e.g. Manatschal et al. 2000); in an active propagating oceanicto-continental rift (Kopf 2001); and in a postcollisional region of extension (Colletini & Holdsworth 2004). Whereas evidence of fluid involvement during faulting is scarce along some LANF (Hayman 2006), other detachment faults are conduits for hydrothermal fluids, meteoric fluids, saline seawater, mantle-derived fluids or voluminous CO2 produced from mantle degassing (Spencer & Welty 1986; Morrison 1994; Losh 1997; Kopf 2001; Collettini & Barchi 2004; Miller et al. 2004; Smith et al. 2008). LANF appear to have been active structures in a wide variety of crustal rocks with both hydrostatic and lithostatic pore-fluid pressures. The initiation of low-angle normal faults remains one of the most contentious issues in fault mechanics. Some authors contend that low-angle normal faults form at a high angle, slip and rotate to shallow angles as simple tilted fault blocks (e.g. Wong & Gans 2008) or along other normal faults (e.g. Proffett 1977). In a series of widely accepted papers, Wernicke (1981, 1985, 1992) and Wernicke & Axen (1988) detailed a ‘rolling hinge’ model to explain the formation and rotation of low-angle normal faults. In the model, a large LANF transects the upper 15– 20 km of the crust and undergoes slip, which results in asymmetrical denudation of the hanging-wall strata. Isostatic rebound then results in flexure of the LANF footwall and corollary rotation of the initially steep fault into shallow dips. Nevertheless, compelling evidence exists for the initiation of shallowly-dipping LANF (John & Foster 1993; Wernicke 1995); this requires both a static and dynamic low-friction mechanism.

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This paper addresses the question: can the application of granular flow theory, coupled with a consideration of the elastic strain energy released by grains deforming by dynamic fragmentation (Davies et al. 2007), explain the origin and significance of the brittle fault rocks exposed along lowangle normal faults? In light of ongoing debate about the apparent weakness of crustal-scale faults such as the San Andreas Fault, this paper: (1) presents a summary of frictional mechanics; (2) provides an overview of fault weakening mechanisms capable of producing low apparent friction on faults; (3) summarizes descriptions of brittle fault rocks exposed along the Chemehuevi and Black Mountain detachment faults, California, USA; and (4) proposes that the widespread occurrence of cataclastic fault rocks in fault cores warrants consideration of how granular flow theory helps to explain the low frictional strength of brittle fault rocks. In conclusion, we suggest that dynamic fragmentation is the most plausible mechanism so far proposed for explaining the low static and dynamic frictional strengths of crustal-scale faults.

Mechanics of friction At temperatures below about 350 8C, corresponding to about 12 km depth in continental crust, quartzofeldspathic rocks fail by brittle fracturing. The constitutive equations governing brittle failure are based on linear elastic fracture mechanics, and the empirical constitutive equation for the shear failure of intact rock is the Mohr– Coulomb criterion:

t  Co þ mi sn0

(1)

where t is the shear stress, Co is the cohesive strength, mi is the internal coefficient of friction and sn0 is the effective normal stress. Terzhagi & Peck (1948) illustrated that the effective normal stress is equal to:

sn0 ¼ sn  Pf

(2)

where the pore fluid pressure Pf is given by Pf ¼ rw gz

(3)

where rw is the density of the pore fluid, g is the acceleration due to gravity and z is the depth below the fluid surface. The formation of new faults occurs at about +308 to the maximum compressive stress when conditions of the Mohr– Coulomb criterion are met. Failure of pre-existing cohesionless planes is governed by Amonton’s Law:

t ¼ ms sn0

(4)

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C. BOULTON ET AL.

where ms is the coefficient of static friction. Rabinowicz (1951) demonstrated that the coefficient of static friction ms, which corresponds to the force needed to start motion, is generally greater than mk, the coefficient of kinetic friction, which is operative for moving surfaces. We use the term ‘strength’ to refer only to the static resistance to shear, and the term ‘apparent friction’ to refer to either the static or dynamic frictional shear resistance of granular material.

Requirements for reduced friction Amonton’s Law (Eq. 4) shows that apparent friction can be reduced by reducing the coefficient of friction ms, or by reducing the effective direct stress sn0 . During sliding, friction is determined by the morphology and properties of the solids in contact, and can only be altered by altering the profile of the contacts or the deformability of the materials. As with solid sliding, reducing friction in granular flow requires either intrinsically low sliding friction between the solid grains, or a reduction of the effective normal stress between the grains. Maveyraud et al. (1999) show that solid sliding friction gives consistent values of m over a wide range of rock materials and normal pressures, as long as the material remains elastic. Thus, there are three possibilities for varying m in sliding friction: (1) by altering the properties of the solid material via frictional-adiabatic heating or isothermal heating with depth; (2) by exceeding the elastic limit of the material and altering the micromorphology of the surfaces via the creation of granular material; and (3) via the chemical alteration of minerals to intrinsically low-friction materials such as graphite or some sheet silicates. The normal stress sn acting between surfaces in contact is the result of the applied stress field, which along faults includes the lithostatic and tectonic stresses; sn can be reduced by the presence of an interstitial fluid that resists some of the applied normal stress (Eq. 2) (a fluid by definition cannot resist a shear stress). If a fault core is occupied by water at hydrostatic pressure, Pf is given by Eq. (3) and the pressure gradient in the water resists normal stress. This is equivalent to the hydrostatic uplift reducing the weight of the overlying saturated rock. In the extreme case, if the pore pressure distribution is lithostatic, pore pressure resists the total weight of the overlying rock and the normal stress across the fault reduces to zero; therefore, so does the apparent friction. Any non-hydrostatic pore-fluid-pressure distribution, however, will tend to cause water to migrate by seepage (unless one of the fault blocks is completely impermeable) and the pressure distribution returns to hydrostatic. The overpressure

will dissipate with time at a rate dependent on the permeability of the fault zone and adjacent damage zones. Besides water, other potential pore fluids include molten rock, water vapour from boiling of pore water, and CO2 from frictional heating of carbonates (e.g. O’Hara et al. 2006; Han et al. 2007). During shear failure in the brittle crust, seismogenic slip occurs on a fault plane at velocities of 0.1–2 ms21 (Brune 1976). Earthquake rupture is a mixed process between frictional slip failure on a plane or thin zone of weakness (or pre-existing fault) and fracture of intact rock. In turn, the frictional strength of a fault during rupture is likely to be highly heterogeneous, incorporating failure criteria (Eq. 1) and (Eq. 4) as well as the local effects of pore water pressure (Pf) (Eq. 2; Ohnaka 2003; Shipton et al. 2006). Thus, the requirements for reducing apparent friction on faults are that: (1) the friction coefficient m is reduced, (2) the effective direct stress sn0 causing frictional resistance is reduced or (3) both of these. This applies to both static shearing (‘failure’) and to dynamic shearing (Fig. 2). Field studies indicate that mature faults with large volumes of granular fault rock commonly exhibit low frictional strength (Sibson 1994; Townend 2006; J. Townend pers. com. 2007). In addition, laboratory data indicate that low rock-onrock friction occurs only with high strain rates and high grain contact stresses (e.g. Di Toro et al. 2004). These observations suggest another criterion for assessing proposed low-friction mechanisms: any low-friction mechanisms which could operate in intact rock at low stress and/or strain rate are poorly substantiated and hypothetical. Thus, our criteria for assessing low-friction mechanisms are: they must be able to reduce apparent friction at both static and dynamic strain rates and they must be only be effective on mature faults with a fault core comprising granular material.

Mechanical explanations for low frictional strength Recent experiments with intact rocks and natural fault gouges have found that very low frictional resistance to shear can occur at high strain-rates (101 s21 to 1022 s21) under low to moderate normal stresses (0.6 to 100 MPa) (e.g. Di Toro et al. 2004; Mizoguchi et al. 2007; see also Wibberly et al. 2008, fig. 7 for a summary of experimental data). The dramatic (more than three-fold) reduction in frictional strength has been attributed to complex thermo-poro-elasto-dynamic processes. However, low-strain-rate (1024 s21 to 1027 s21)

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

13

Fig. 2. A Mohr diagram illustrating the failure criteria for frictional sliding (solid line) of a misoriented cohesionless fault with a Byerlee friction coefficient m ¼ 0.6 at hydrostatic pore fluid pressure Pf. According to Amonton’s Law, shear failure will occur with a reduction in the apparent friction coefficient (m0 ) (dashed line) or with an increase in pore fluid pressure (P0f) (dotted line). Inset figure depicts the Andersonian stress state for normal faulting (after Townend 2006).

rock-on-rock friction tests invariably show conventional friction coefficients of 0.6 to 0.85 (Byerlee 1978). In the Earth’s crust, fault rocks are generally tectonically loaded at very low strain rates (10212 s21) that are impractical to simulate in the laboratory. A number of explanations for low rock friction at very low strain rates have been proposed, mainly in the context of large faults such as the San Andreas (e.g. Townend & Zoback 2001; Scholz & Hanks 2004). Explanations for static and dynamic weak fault friction include the following: 1.

2. 3. 4.

5. 6.

weak materials such as silica gel (Goldsby & Tullis 2002), talc (Moore & Rymer 2007), clay (Tembe et al. 2006; Numelin et al. 2007) or melted rock (pseudotachylyte; e.g. Sibson 1975; Sibson & Toy 2006; Hirose & Bystricky 2007) reducing apparent friction; flash heating of asperities (Rice 1999, 2006; Hirose & Bystricky 2007); the presence of long-term pore fluid overpressures reducing effective stress (Sleep & Blanpied 1992; Sibson 2001); thermal pressurization of pore fluids during frictional sliding (Sibson 1973; Andrews 2002) or elastohydrodynamic lubrication (Brodsky & Kanamori 2001); normal interface separation during rupture along a bimaterial interface (Weertman 1980; Ben-Zion 2001); acoustic fluidization (Melosh 1996).

Here we briefly describe the mechanisms proposed for low-strength faulting, and comment on their apparent advantages and drawbacks, before evaluating them against the above criteria.

High pore fluid pressure Due to the well-known effect of pore water pressure in reducing the effective stress, this is possibly the most commonly-used mechanism (Eq. (2); Terzhagi & Peck 1948; Hubbert & Rubey 1959, Sibson 1973). We discuss the mechanism first in the context of static failure. Static failure. Fault strength might be partly controlled by the evolution of fluid pressure in the fault zone. Arrays of mineralized veins surrounding ancient and modern normal faults provide evidence for fluid cycling, and mechanical theory can account for the failure of LANF by pore fluid overpressures (Eq. (2); Parry & Bruhn 1990; Axen & Selverstone 1994; Sibson 2001). High pore fluid pressures can also be generated interseismically by compaction creep and porosity reduction (Sleep & Blanpied 1992). Scholz (1992) noted that brittle wall-rock material cannot resist pore fluid pressure greater than the least principal stress s3 without fracturing in response and allowing pore fluid to drain. Rice (1992) published a classic paper showing that fault cores can sustain high pore pressures due to locally inhomogeneous stress distributions. The Rice (1992) model requires that high pore-fluid pressures be maintained within relatively permeable fault cores by either an upwelling of overpressured fluids from depth or by impermeable seals which trap fluids at high pressures (e.g. Byerlee 1990). Axen (1992) applied this model to detachment faults in the Basin and Range, USA; he cited many examples of fluid focusing and seal formation along and near LANF fault cores. Recently, Healy (2008) expanded the model of Faulkner et al. (2006) to show that large stress rotations are possible on fault core rocks with an anisotropic

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C. BOULTON ET AL.

fabric; these stress rotations enable elevated pore pressures to develop without hydraulic fracturing in strike-slip faults. Models such as these have not yet been developed for normal faults, but it appears that pore-fluid overpressures can develop on LANF given the required low permeability. A number of researchers have investigated the permeability of fault gouge. Takahashi et al. (2007) found that dramatic slip-induced reduction of permeability occurred in quartz gouge with Na-montmorillonite at concentrations between 18 and 24% at a confining pressure of 80 MPa, suggesting that clay material concentrations of this magnitude are needed to cause sealing. At .1 km depth in the Nojima fault zone, Japan, Lin et al. (2007) report that although fracture networks are important fluid conduits, they lose permeability as pore spaces collapse and cementation proceeds. Giger et al. (2007) found that quartz gouge could seal hydrothermally by a combination of dissolution and precipitation creep processes above 700 8C, relevant to the mid-lower crustal regions. Permeability-reducing processes compete with hydrolytic weakening, stress corrosion and distributed microcracking, which act to increase the permeability of fault-zone rocks (Sornette 1999; Crampin & Chastin 2003). Zoback & Townend (2001) showed that fault-zone permeability is high and suggested that frictional failure on faults must dominate (largely aseismic) creep and compaction processes. Zoback et al. (2007) found no evidence of either cohesive materials or porewater overpressures in the 2.7 km deep San Andreas Fault Observatory (SAFOD) drill hole, California, USA. Furthermore, whether pore-water overpressures can initiate a seismic rupture on weak faults remains a topic of ongoing research; models suggest that ruptures preferentially nucleate in regions of low pore fluid pressure (e.g. Hillers & Miller 2007). Dynamic failure. Sibson (1973) suggested that frictional heating due to slip in an earthquake could raise the pressure of pore fluid in a fluid-saturated fault zone to near-lithostatic values, reducing effective pressure and frictional resistance during slip. Garagash & Rudnicki (2003a) concluded that shear heating might induce the high pore-fluid pressures that destabilize slip. However, complicated feedback mechanisms occur because high pore-fluid pressures induce fault dilatancy, which subsequently reduces shear heating. The feedback process was analysed further by Garagash & Rudnicki (2003b), who concluded that weakening effects were strongly dependent on whether fault zones are hydraulically and thermally isolated from the surrounding rock mass. Indeed, permeability is unlikely to remain constant during an

earthquake and it may become large enough that fluid pressure is limited by the permeability of the surrounding damage zone (Andrews 2002). A difficulty with thermal pressurization of aqueous fluid is that seismic slip is required to generate sufficient heat to elevate pore-fluid pressures. This topic has recently been examined closely by Rice (2006) and Rudnicki & Rice (2006). Consideration of fault-rock permeabilities leads to the conclusion that shear heating alone cannot nucleate unstable slip in the absence of independent slip weakening, but that it can have an effect following strong stress perturbations. If fault rupture is preceded by slow slip (mm s21), thermal pressurization can occur, but over the critical weakening distance, slip will be at normal friction because there is no low-friction mechanism operating. This means that the overall friction coefficient must be a combination of normal friction in the early phases and low friction during the late phases. Finally, while high pore fluid pressures may develop more readily in the thick gouge of large-displacement (.1 km) faults, this is offset by the greater difficulty of establishing and maintaining high pressures over sufficiently large areas of the fault. Brodsky & Kanomori (2001) presented an analysis elastohydrodynamic lubrication. According to this theory, grains in a fault core act like bearings that are lubricated by a thin film of slurry comprising finely comminuted gouge mixed with water. During frictional sliding, the slurry decreases the apparent friction by increasing the pore-fluid pressure. In the model, full lubrication pressure is reached after a critical slip distance of c. 0.5 m if the permeability of the surrounding rock is sufficiently low to confine the fluid to the slipping fault. A significant advantage of this model over thermal pressurization of pore fluids is that, at most, fluid pressures reach 40% of the lithostatic load; they are therefore unlikely to induce hydraulic fractures. Like thermal pressurization, however, elastohydrodynamic lubrication depends on impermeable wall rocks bounding the fault core and a critical slip distance in order to act as an effective mechanism for reducing dynamic fault friction.

Weak fault materials Minerals. Fluids in fault zones induce chemical reactions that alter the rheology of fault rocks. Most of the reactions result in the formation of weaker minerals (e.g. Janecke & Evans 1988; Gratier & Gueydan 2005). Weak materials with low coefficients of friction (m ¼ 0.2–0.4) include platy minerals such as talc, water-saturated sheet silicates, and graphite (Moore & Lockner 2004). If platelets are aligned parallel to each other, weak

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

interstitial bonds (talc, graphite) or the highpressure interstitial water (clays) allow the platelets to shear at low applied stress. A shear zone containing these materials can therefore be expected to have low static friction and also low dynamic friction if the orientation of the platelets is not altered by large strains. Recently, Moore & Rymer (2007) suggested that principal slip surfaces composed of talc, a low shear strength mineral that accommodates shear displacement by (aseismic) stable sliding, may be responsible for the observed static and dynamic low frictional strength of the San Andreas Fault. At lower greenschist-facies conditions at the base of the seismogenic zone, feldspar breakdown to fine-grained muscovite (or sericite) commonly occurs, and this reaction may be an important reaction-softening step in granitic fault zones. This mechanism has also been invoked to explain the low frictional strength of the San Andreas Fault zone (Evans & Chester 1995; Wintsch et al. 1995). At high temperatures and pressures, micas deform relatively easily by grain-boundary sliding, cleavage-plane slip and dislocation glide; the formation of an aligned mica foliation weakens highly strained fault rocks. Wibberley (1999) pointed out, however, that granitic faults may strengthen or weaken due to feldspar to mica reactions depending on the behaviour of silica released during the reactions; silica cement strengthens faults. During dynamic fault failure, Sulem et al. (2007) showed that shear heating and fluid pressurization are possible mechanisms leading to full fluidization of clay material in a shear band. Hirose & Bystricky (2007) also showed in high-velocity laboratory experiments that phyllosilicate dehydration due to frictional heating resulted in a marked decrease in frictional strength after very high strains. Boutareaud et al. (2008) tested natural clay-clast (16% clay) aggregates from fault gouge in a shear apparatus at low stresses (0.6 MPa) and found that very high strains (tens of metres) were needed to significantly reduce friction. Fault rupture zones, however, rarely contain high concentrations of these materials. Usually the low-friction materials are found at low concentrations in the fractal, cohesionless angular rock fragments comprising the gouge. Their ability to reduce the total friction of the gouge depends critically on the material being present in high concentrations or localizing onto very thin principal slip surfaces. Takahashi et al. (2007) found that the friction coefficient of quartz gouge fell from 0.7 to 0.4 as the clay (Na-montmorillonite) concentration increased from zero to 50% by volume. Brown et al. (2003) found similar behaviour with smectite, illite and chlorite. In addition, the

15

frictional resistance of many sheet silicates varies with depth: pressure and temperature conditions, as well as slip rate, determine whether the minerals exhibit stick-slip (seismic) or stable (aseismic) frictional sliding (Moore & Lockner 2004). Testing material recovered from the SAFOD drill hole, Carpenter et al. (2007) found normal friction coefficients for most materials. Serpentinite melange (which can produce talc) had a friction coefficient of 0.220.3, but was found to strengthen with increasing velocity under both dry and saturated conditions. In summary, if low-friction materials are present at high concentrations or are localized onto thin principal slip surfaces, reduced frictional strength can occur. Numelin et al. (2007) highlight the importance of clay fabric development in reducing the frictional strength of LANF in the upper brittle crust (,5 km depth), and emphasize that clays enable the faults to undergo stable (aseismic) slip. This mechanism for low apparent friction depends strongly on the spatial distribution and interconnectivity of the weak clay fabric. Dynamic (seismic) fault failure involving grain comminution down to nanometre size (Heilbronner & Kuelen 2006) at high strain rates in narrow shear bands (Rice 2006), probably disrupts the low-friction arrangement of parallel platelets. Silica gel. Frictional sliding on smooth surfaces of rock with silica contents of 70% or greater facilitates the generation of a layer of silica gel that results in a dramatic decrease in shear resistance (Yund et al. 1990; Goldsby & Tullis 2002; Di Toro et al. 2004). These studies show that the amorphous material formed as a result of grain comminution at lower speeds, 10 mm to 100 mm s21, than required for weakening caused by flash heating, but it requires displacements of about 1 m to attain steady state. Thus, the weakening distance may be related to the distance necessary to accumulate a thick layer of gouge and/or lubricant. Based on inference and microstructures, Goldsby & Tullis (2002) and Di Toro et al. (2004) suggested that shear lubricates the surface through production of a highly comminuted and amorphous material that may be thixotropic silica gel, a near-fluid that becomes less viscous when strained. When shearing ceases the material regains strength with time at a rate controlled by the bond forming chemical reaction.

Heating Under the stresses experienced by a large fault during rupture, friction inevitably generates heat. Thermal pressurization of pore fluid was discussed previously. This section will focus on melting of

16

C. BOULTON ET AL.

rock to generate pseudotachylyte (e.g. Hirose & Shimamoto 2005) and flash heating at high-stress contact asperities with local melting (e.g. Rice 1999). Clearly, these mechanisms relate only to dynamic friction as ambient temperatures sufficient to develop melt do not occur in the brittle crust. Rock melting. Pseudotachylyte is a very fine-grained fault rock that is generally thought to form coseismically by frictional melting and comminution. Spray (1995) demonstrated experimentally that comminution is a necessary precursor to frictional melting, and his results have been verified by Di Toro et al. (2004) among others. Field evidence suggests that pseudotachylytes may be divided into melt-origin types which preserve glass or glassy material (e.g. Philpotts 1964; Sibson 1975) or crush origin types, which form by mechanical stress during intense grain comminution (e.g. Ozawa Takizawa 2007). Pseudotachylyte formed by frictional melting is rarely found in the geological record, suggesting either that it rarely forms or that it is rarely preserved (Sibson & Toy 2006). Documented examples of melt-origin pseudotachylyte formed during coseismic slip indicate that pseudotachylyte first appears as isolated patches of viscous fluid, which increase the frictional resistance to shear; frictional resistance decreases significantly only when a continuous molten layer is established (Warr & Van der Pluijm 2005; Hirose & Shimamoto 2005; Fialko & Khazan 2005). Pseudotachylyte formed by the comminution of grains to amorphous materials several tens of nanometers in size can behave similarly to melt-origin pseudotachylyte (Ozawa & Takizawa 2007). Pseudotachylyte formed in this way occurs in the field as both fault veins and injection veins, indicating that it flowed under pressure gradients like a viscous fluid. We shall later show that dynamic comminution can allow a granular material to flow under relatively low applied stresses, and this process may also explain the behaviour of crush origin pseudotachylyte without invoking high temperatures. Flash heating of asperities. During rock-on-rock sliding experiments, flash heating at asperity contacts causes weakening at slip speeds above 100 mm/s in quartzite, gabbro, and granite, which all melt at room pressure, as well as soda lime glass (Goldsby & Tullis 2002; Rempel 2006). Flash heating is velocity-dependent, and above 100 mm s21, the coefficient of kinetic friction drops rapidly to ,0.2 when shear localizes along a thin principal slip zone (Rice 1999, 2006). Flash heating at asperity contacts does not occur in other rock types. In calcite marble and dolomite,

calcite undergoes a phase change and breaks down to CaO and CO2. Recent reviews reveal that the breakdown of clays in natural fault gouge via dehydration may prevent weakening from flash heating because of pore-fluid pressurization. During high velocity shearing, clayey fault gouge becomes fully fluidized, but it remains difficult to determine whether the effect is from shear heating or fluid pressurization (Brantut et al. 2007; Sulem et al. 2007). Recent experiments by Goldsby & Tullis (2007) suggest that gouge material behaves differently than intact rocks. In the presence of a thin fault gouge (,1 mm), a velocity greater than coseismic slip rates (.1 m s21) is required to initiate flash heating at grain contacts. This suggests that flash heating at asperity contacts does not occur in fault gouge. Given the ubiquity of gouge in real faults, the role of flash heating of asperities in fault weakening appears questionable.

Mechanical processes Here we deal with mechanisms that stem from aspects of contact forces between rocks or fragments experiencing shear forces, which do not require the presence of anomalous materials, pore fluids, or heat. Self-healing slip pulses. The rupture characteristics of a number of recorded earthquakes correspond to the propagation of a coherent slip pulse along the fault, with the fault healing itself again after the passage of the pulse (e.g. Heaton 1990). This makes sense: for example, suppose a fault rupture causes 1 m of horizontal ground motion at a velocity of c. 1 m s21, taking 1 second to complete. At a rupture propagation velocity of 2 km s21, the rupture will propagate c. 2 km while the ground at any one point is moving. Thus the passage of the rupture along the fault will be in the form of a 2 km long non-uniform slip pulse propagating at 2 km s21. Self-healing slip pulses propagate near the Rayleigh wave speed along bimaterial interfaces (Heaton 1990; Ben-Zion 2001) and generate spatially localized normal interface separation. Small faults, on which the majority of ongoing seismicity occurs, may be considered to be surrounded by nominally homogenous solid rock. However, plate boundaries and other major faults generally resemble bimaterial interfaces because zones of crushed rock, cataclasite and foliated gouge abruptly abut intact rock. In addition, different rock bodies are juxtaposed across large displacement faults. Low-angle normal faults, for example, commonly juxtapose high-grade metamorphic and plutonic footwall rocks against poorly consolidated sedimentary and/or volcanic hanging-wall rocks (e.g. Coney 1980).

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

Rupture along a bimaterial interface may differ significantly from a corresponding rupture in a homogeneous solid. The differences arise from the fact that ruptures propagating along a bimaterial interface induce dynamic changes in normal stress near the tip causing normal stress to vary as a function of slip (Weertman 1980; Ben-Zion 2001). The interaction between slip and normal stress along a bimaterial interface can dynamically reduce the friction strength to zero and may lead to local fault opening (Eqs 1 and 4) (Andrews & Ben-Zion 1997; Anooshehpoor & Brune 1999). This makes bimaterial interfaces mechanically favoured surfaces for rupture propagation and may help to help the reactivation of misoriented faults with lower resolved Coulomb stresses rather than the formation of new faults optimally oriented for failure (Fig. 2). However, the degree to which interfaces separate during slip remains poorly understood (Rubinstein et al. 2004). In addition, this mechanism requires slip to be present in order to initiate the pulse, so applies only to dynamic shear resistance. Acoustic fluidization. This mechanism was first proposed by Melosh (1976) in the context of impact crater formation, and applied by Melosh (1996) to faulting. It involves the effect of vibrations passing through granular materials, in such a way that the interparticle stresses vary significantly with time; in particular, at times the direct stresses are low so that frictional resistance to sliding is low. The mechanism is thus specifically applicable to granular gouge, but it presumably requires some initial source of vibration so is applicable only to dynamic friction. Sornette (1999) has closely examined acoustic fluidization in the earthquake context, noting (a) that the mechanism predicts slip velocities far smaller than those observed during earthquakes, suggesting that if acoustic fluidization is present it is not significant; and (b) that modifying the analysis to address this point by using realistically-varying parameters creates a high degree of non-linearity. Melosh (1996) claims that the theory has been confirmed by experiment, but one of the quoted sources (Zik et al. 1992) for example, reports reduced friction on a sphere moving through other spheres that are ‘vibrating at accelerations greater than that of gravity’; whereas Collins & Melosh (2003) are at pains to point out that in acoustic fluidization, individual grains are not vibrating energetically. Another significant difficulty with the concept of acoustic fluidization is that, as presented, it takes no account of grain comminution. As we discuss below, comminution drastically alters the behaviour of shearing grains, so the applicability of acoustic fluidization to faulting is difficult to assess.

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Evaluation of mechanisms We have five main contenders for causing the low apparent friction exhibited by large-scale faults: (1) pore fluid overpressure and thermal pressurization of pore water; (2) weak material (talc, silica gel, clay); (3) heating (rock melting and flash heating at asperity contact); (4) self-healing slip pulse; (5) acoustic fluidization. We also have two major criteria on which to evaluate them: (a) does this mechanism function equally well in both static and dynamic conditions and (b) does this mechanism only function for large-scale faults? The result of evaluation on this basis is shown in Table 1. None of the existing mechanisms has the potential to explain the static and dynamic weakness of large-scale faults, together with the high strength of small faults. The high strength of small faults requires only that static or dynamic apparent friction be high, not both. High pore fluid pressure fails because it also predicts both low static (due to long-term overpressure) and low dynamic (due to heating) apparent friction in small faults. The other mechanisms, being applicable equally to both large and small faults, predict that large and small faults should behave identically, which is not the case. In the light of this conclusion, we now review the effects of rock fragmentation on the dynamics of confined granular flow; this mechanism has been outlined in increasing detail in the context of land- and blockslides (Davies et al. 1999, 2006, 2007; Davies & McSaveney 2002, 2005, 2008a, b; McSaveney & Davies 2005, 2006, 2007; McSaveney et al. 2000; Smith et al. 2006). It appears to meet both the above criteria.

Granular friction The Mohr– Coulomb criterion (Eq. 1) and Amonton’s Law (Eq. 4) work very well to describe the macroscopic friction of both solids and granular materials sliding tangentially with respect to another. Detailed experiments conducted by

Table 1. Evaluation of weak fault mechanisms Mechanism Fluid overpressure Weak material Heating Self-healing slip pulse Acoustic fluidization

Static Dynamic 3 3   

  3 3 3

Largescale only     

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Bowden & Tabor (1986), however, showed that frictional strength correlates strongly with the real area of contact between two sliding surfaces, where real area of contact is the sum of the contacting spots or asperities on a surface. Very smooth surfaces and very rough surfaces have higher contact areas than intermediate surfaces, and have higher friction coefficients (see Rabinowicz 1995 for a detailed discussion). In granular materials, the relationship between the frictional strength of the material and contact area is further complicated by the fact that multiple variably-sized and irregularly-shaped particles carry the shear and normal stresses at point contacts. Hence the stresses at the individual contacts are indeterminate (but much higher than the overall mean stress), and the need for large-scale averaging is much more apparent than with solid friction (e.g. Majmudar & Behringer 2005). Moreover, the spatial arrangement of platy grains such as sheet silicates may be initially non-random, causing unusually high or low initial (static) frictional strength. During compressive shear deformation, granular materials form an inhomogeneous contact network which carries most of the external load along force chains (e.g. Matsuoka 1974; Cundall & Strack 1979; Cates et al. 1999; Morgan & Boettcher 1999; Anthony & Marone 2005). Force chains are quasi-linear structures of highly stressed grains surrounded by less-stressed grains in a weak network. Most of the stress applied to the shearing granular material is transmitted from one shear boundary to the other along force chains. In simple shear these

initially assemble parallel to the direction of maximum compressive stress and strain (about 458 to the shear zone walls) and subsequently rotate in the shear strain field to cause both shear zone dilatation and increase in force-chain shear stress (Fig. 3). Granular flow requires force chains to fail so that strain can continue, and failure can be by intergranular slip, grain rotation, buckling or grain crushing (Biegel et al. 1989; Hooke & Iverson 1995; Fig. 3). Under high lithostatic stresses with angular grains, grains are most likely to fail by buckling, or alternatively by crushing, a process in which force-chain loading causes a particle to fragment (Sammis & King 2007). Following Grady & Kipp (1987), we term particle crushing during shear dynamic fragmentation. The strength of grains varies with grain size, temperature, confining pressure, and strain rate; for silicates, grain strength also varies with the presence of absence of water (e.g. Paterson 1978). The way in which force chains fail also varies with these factors. Force chains continuously form and fail to accommodate shear deformation, and this process causes sheared granular materials to develop selfsimilar particle-size distributions (PSDs) (Marone & Scholz 1989). In confined flow, force chains are generally ten grains wide and the total width depends on the distribution of particle sizes within the granular material (Francois et al. 2002). Sammis & Ben-Zion (2008) showed that during compressive failure, shock loading and subcritical crack growth can form fragments much smaller (,10 nm) than the theoretical grinding limit

Fig. 3. Force chains are quasilinear structures that support most of the internal stresses in shearing granular material. The frictional strength of granular material is related to the stresses resolved on the force chains within it. Force chains fail by sliding, spalling, fracturing, crushing and fragmentation during cataclastic deformation. In reality, grains within a fault gouge would likely have fractally distributed shapes and sizes (after Rawling & Goodwin 2003; Davies et al. 2007).

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

(1 mm). This result means that grains will continue to fragment to smaller sizes in granular flow, driving shear localization onto thinner principal slip zones as comminution proceeds. For a fault core containing granular gouge with a mean grain size of 100 mm, slip may occur along a plane 1 mm wide. Fault cores on large-displacement faults commonly contain millimetre-thick ultracataclasite along which a majority of fault slip is inferred to have taken place (e.g. Chester & Logan 1987). Results from numerical models of compressional shear fracture initiation in intact crystalline and porous rocks show that force chains comprising individual minerals or grains may form to carry the applied stress in a manner similar to that in noncohesive granular materials (Potyondi & Cundall 2000; Hazzard et al. 2000).

Dynamic rock fragmentation As outlined above, experimental, numerical and computational studies in granular physics indicate a small proportion of grains in a granular material carry the bulk of the applied shear stress and normal stress. The frictional behaviour of dense granular material depends on micro-scale interactions that occur at grain contacts in force chains, and on the interaction between grains in the strong and weak networks (e.g. Garcia & Medina 2008). The physics governing these contacts is complex, and information gathered from models indicates that the macroscopic frictional behaviour of granular material depends on: elastic grain –grain interactions; a repulsive nonlinear viscoelastic contact force between grains; grain mass; and the friction coefficient of the material (e.g. Campbell 2002; Zhu et al. 2008). Most models simulate grains as intact spheres because simulating particle fracture is very difficult. However, under high confining pressures and/or strain rates, cataclasis involving particle fragmentation may redistribute energy in grain flows through the process of dynamic rock fragmentation. Dynamic rock fragmentation (Grady & Kipp 1987) is the mechanism whereby an intact piece of rock fails in a brittle manner (i.e. rapidly or catastrophically) under load, generating fragments that move away from the original centre of their mass at high velocity. This is the same effect as the ‘. . . large rapid compressions and decompressions . . .’ associated with asperity failures by O’Hara (2005). That fragmenting grains generate high fragment energies and hence high local pressures, is widely acknowledged, and demonstrated even under very slow applied strain rates by the phenomenon of rock-bursts in mines. Moreover, it is universally accepted that comminution or fragmentation of rock is a ubiquitous process in shearing

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of fault gouge in earthquakes (e.g. Sammis & Ben-Zion 2008). Davies et al. (1999) and Davies & McSaveney (2008a) proposed that dynamic fragmentation could explain the extraordinarily long runout distance of large landslides. According to this theory, rock avalanches, shear planes at the base of large blockslides, and fault cores all comprise granular material which accommodates shear displacement by the formation of and failure of force chains (Fig. 3). The failure of force chains by particle fragmentation creates a granular material comprising both intact grains and angular grain fragments; the interaction between these particles likely affects the rheology of the flow (e.g. Guo & Morgan 2006). Fault rupture involves high strain rates between 0.1 and 2 ms21 (Brune 1976) on thin localized slip zones about 1–5 mm thick (Rice 2006). Spatially averaged strain rates are in the range 20 –2000 s21. This means that the failure of grains is virtually instantaneous; failure conditions overlap with those of shock loading (.100 s21; Melosh 1993; Sammis & Ben-Zion 2008). Grains within force chains loaded at such high strain rates likely fail by dynamic fragmentation, during which the elastic strain energy stored in a fragmenting grain is converted instantaneously to kinetic energy. In fault gouge at seismic depths, the free travel distance of the fragments is extremely small and their kinetic energy is thus immediately converted into pressure energy exerted on the surrounding grains (Fig. 4). Under these conditions there is evidence that a fragmenting grain behaves as a high-pressure fluid (e.g. Benz & Asphaug 1995). We have shown that the magnitude of this fluid pressure is of the same order as the failure stress of the grain, which at seismic depth at the high strain rates experienced by fault cores during rupture is in the gigapascal range (Davies et al. 2007; Davies & McSaveney 2008b). The thermal effect of fragmenting grains in increasing the temperature of the material on which it impacts follows the mechanical effect of exerting pressure (Yuan & Prakash 2007), so the pressure effect is the primary effect. Dynamic fault friction. During grain failure by dynamic fragmentation, the rapid and different alterations of the grain’s elastic and shear moduli during damage transform it into a short-lived highpressure fluid, able to resist compression but not shear (Fig. 4). In this way, the fragmenting grain becomes part of the fluid phase in the bi-phase fluid –rock fault core. The ability of the force chain to transmit shear across the slip zone vanishes, while the normal stress continues to be transmitted across the slip zone due to the granular fluid pressure exerted by the fragmenting grain on the

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Fig. 4. (a) Forces on a force chain just prior to fragmentation. Like Figure 3, grains within a fault gouge would likely have fractally distributed shapes and sizes. P is the failure strength of the weakest grain in the force chain and S its resistance to shear. u is the angle between the force chain at failure and the resolved normal stress sn. The value of u is generally between 0 and 458. Lockner (1995) showed, to a first approximation, that at high confining stress, the confined strength of a grain is equal to its unconfined strength plus the confining pressure. P cos u þ S sin u  Ssn , S cos u þ P sin u  St: (b) Forces on a force chain during fragmentation. As stored elastic strain energy is released as kinetic energy, the fragmenting grain generates a pressure on other grains. S ¼ 0 because a fragmenting grain has no shear strength. P cos u ¼ Pf cos u  Ssn ; P sin u  St:

grains immediately surrounding it. In effect, the total shear resistance reduces to that of the weakly stressed force chains in the surrounding gouge. We envisage that at all instants as the rupture propagates along the fault, many force-chains are failing by fragmentation in this way and are dynamically reducing the resistance to shear on the fault. Dynamic fragmentation is a well-known phenomenon of rock mechanics (e.g. Grady & Kipp 1987), but it has been studied mainly in the context of the need to crush rocks to produce aggregate for concrete. The energy input required for this process is a cost to that industry, so is considered wasted; in the context of fragmenting granular shear, however, the energy input required to cause the rock to fail is recovered instantaneously as kinetic energy of fragments, and immediately as a high pressure on the surrounding grains. The major criticism that has been raised with respect to the role of fragmentation in landsliding is that fragmentation is an energy sink, so it must increase rather than reduce friction. This is usually deduced from the common perception, originally proposed by Griffith (1920), that a certain quantity of energy is required to create a given area of new grain surface, and this energy instantly becomes

unavailable for dynamic effects because it somehow attaches itself to the new surface. Fracture energy does correlate well with new area created, but there is no physical justification for assuming that it is instantaneously lost to the system dynamics. In fact, the elastic strain energy released at fragmentation is instantaneously available as kinetic or pressure energy. Although the released elastic strain energy degrades rapidly to heat with friction and other processes its immediate effect on fragment velocities and hence intergranular dynamics is very important (McSaveney & Davies 2008). Dynamic fragmentation thus appears to be a mechanism that is capable of reducing frictional resistance to granular shear, and has been shown to be capable of explaining quantitatively a range of field and laboratory data (Davies et al. 2007; Davies & McSaveney 2008a). We suggest that dynamic fragmentation occurs during the shearing of granular materials such as fault gouge; in the following we shall show that it can also explain the low static strength of large faults. Fragmentation and the static strengths of faults. The basis of this concept is that the local stress in any fault is spatially non-uniform; and the nature of

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

this variation can be described by extrapolating to fault-scale from laboratory and theoretical data on the distribution of interparticle stresses in static granular media. Figure 5 shows the distribution of normal static contact stresses in particles within a volume of granular material; the distribution has been verified by several experimental (Mueth et al. 1998; Løvoll et al. 1999; Blair et al. 2001) and computational (Radjaie et al. 1996; Silbert et al. 2002) studies. Behringer et al. (2008) have shown that dynamic force distributions are similar in shape. Of interest to us is the distribution of forces larger than the mean, which is the well-defined exponential tail to the right of sn  log101. Ngan (2003) and Van Eerd et al. (2007) show that the relationship between probability and stress in this region is: P(sn ) ¼ exp (sn )b

(5)

with the value of b  5/3. Equation (5) implies a fractal scaling relationship between the maximum normal stress acting on a grain within a force chain and the mean normal stress acting on the granular material. Assuming that the probability of a grain experiencing a given stress increases with the number of grains and that fault thickness remains constant, (5) can be written

sn max  (ln Psn )0:6 ¼ (ln L)0:6

(6)

where L is the length of the fault. This relationship indicates that longer faults should have higher values of maximum local stress. When a grain fails due to local stress, the higher this stress, the higher the stress that is immediately transferred to

Fig. 5. Probability distribution P(sn) of normal stress (sn) normalized with respect to the mean stress (i.e. local stress divided by mean stress) (after Løvoll et al. 1999).

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adjacent grains. If the adjacent grains are sufficiently stressed, then the stress transfer will cause them to fail in turn. Here again we use the concept that a fragmenting grain generates a local pressure equal to its strength, but in static failure, the applied strain rate is very low so the rock strength is probably lower than in the dynamic case. Earthquake ruptures are unstable episodes of irreversible slip during which stored elastic strain energy is released and seismic waves are radiated. On a microscopic scale, fragmentation of highly stressed grains may release enough elastic strain energy (c. 1 GPa) to nucleate a self-propagating earthquake rupture. This suggestion follows from experimental, numerical, and geological evidence gathered by Sammis et al. (1986), Lockner et al. (1991), Reches & Lockner (1994), Reches (1999), Ben-Zion & Lyakhovsky 2002, Hamiel et al. (2004) and Reches & Dewers (2005) that the nucleation and dynamic propagation of earthquake ruptures involves interacting fractures in a small volume of intact and/or granular rock. Our analysis of the statistical distribution of grain strength evolution with time indicates that larger faults, with a greater volume of granular material in the fault core, have a greater likelihood of a strong grain fragmenting at a low mean applied stress. A small fault will require correspondingly higher mean applied stress to cause sufficiently high local stresses to cause grain fragmentation. Thus, a self-propagating rupture is more likely to initiate at lower regional stresses on a large fault than on a small one experiencing a similar stress state. He et al. (1990) show that heterogeneity of the strength distribution over the fault reduces the ability of an initial failure to propagate. We suggest that two changes will occur with time after a fault rupture that will decrease its heterogeneity: First, tectonic stress build-up will cause progressive failure of weak force chains, which increases the minimum local strength of granular material along a fault. Second, stress corrosion (Sornette 1999) will progressively reduce the strength of the most-stressed force-chains, leading to a reduction in maximum strength. As shown in Figure 6, this results in a reduction in the overall strength of granular material along the fault. It also results in a significant reduction in strength heterogeneity. We believe that, as time passes and stress builds up, a fault becomes progressively more likely to experience a large-scale rupture as its strength distribution alters. In terms of the evaluation Table 1, the fragmentation mechanism meets all the required criteria. We postulate that dynamic fragmentation is the most plausible mechanism presented to date for explaining the static and dynamic weakness of faults.

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Fig. 6. Increase of tectonic stress distribution (full lines) and evolution of granular material strength distribution (broken lines) with time in temporal sequence 1, 2, 3, where p (%) indicates the proportion of grains within a volume of granular material.

Discussion Do LANF initiate at anomalous inclinations? The initiation of LANF in intact rock at angles that depart from the Anderson– Byerlee fault mechanics remains an outstanding problem in structural geology. Wong & Gans (2008) recently reviewed the rolling-hinge model of Wernicke & Axen (1988) and presented a thorough discussion of new geological tests for the model. Another outstanding question remains the role of the ductile lower crust in initiating new faults. Echoing the original ideas of Wernicke (1981), Wernicke (1985) and Lister & Davis (1989) argued that low angle detachment faults are rooted in the ductile lower crust and propagate upward as loci of simple shear extending the entire thickness of the brittle crust. Axen (1992) developed this idea and presented a model showing that LANFs bounding metamorphic core complexes could initiate because of mechanical and permeability anisotropies that develop due to ductile shearing. Vanderhaeghe et al. (1999) suggested that models with a brittle crust overlying a ductile crust (Brun et al. 1994) are more relevant to detachment faults formed in thickened and thermally weakened crust. In these models, brittle extension in the upper crust is accommodated by pervasive ductile flow in the lower crust. The style of extension of the upper brittle crust controls fault zone

geometry, and the low angle detachment fault corresponds to a zone of mechanical decoupling between the brittle upper and ductile middle crust, which is equated with the 450 8C isotherm within the continental crust (Brace & Kohlstedt 1980). Recent advances in our understanding of the strength of the lithosphere summarized by Jackson et al. (2008) have renewed debate on the role of the ductile crust in nucleating crustal scale fault zones through transient brittle instabilities (e.g. Handy et al. 2005). This new interpretation of a strong, thick continental lithosphere may lead to a better understanding of the growth and development of crustal-scale faults like LANF through advances in the field of geodynamics.

Seismic rupture or aseismic slip? Faults are highly heterogeneous and the true static strength of an individual fault is likely to be an average of its many different asperities or frictional contacts. Asperity type changes with time as faults accrue displacement. This occurs because of a gradual evolution in fault geometry, fault-rock lithology, and fault-rock mineralogy. Three primary processes affect the nature of strain accommodation in the cores of large-displacement faults. First, faults increase in length through the linkage of irregular fault surfaces and fault segments, generally leading to a more planar fault with time (e.g.

LOW FRICTION IN LOW-ANGLE NORMAL FAULTS

Peacock & Sanderson 1991; Anders & Schlische 1994; Childs et al. 1995, 2008). Secondly, wear processes result in the progressive accumulation of granular material in the fault core (Robertson 1982; Scholz 1987). Compilations of data over several orders of magnitude of displacement show general linear trends of fault zone thickness increasing with displacement. Although fault zone thickness varies with lithology, fault zones appear to reach maximum thicknesses of about 100 m after a displacement of between 1 and 10 km. Thirdly, fault cores commonly channel chemically active fluids that promote the syntectonic metamorphic alteration of feldspar to weak phyllosilicates (e.g. Janecke & Evans 1988). These minerals may alter the frictional strength of faults by forming an interconnected network of shear planes with low frictional strength (Bos & Spiers 2002; Collettini & Holdsworth 2004; Imber et al. 2008). Because of these spatial, geometrical, mineralogical and temporal variables, mature crustalscale faults have highly heterogeneous fault cores, which appear to accommodate slip on a variety of structures, including: very thin hundreds of micron-to centimetre-thick ultracataclasite shear surfaces in a single fault core (e.g. Punchbowl Fault, Chester & Logan 1987); along multiple creeping subparallel fault cores cutting a wider, up to 1 km thick, damage zone (e.g. San Andreas Fault, Hickman et al. 2005); or along multiple strands phyllosilicate-rich fault gouge forming a fault core up to 1 km thick (e.g. Carboneras Fault, Spain, Faulkner et al. 2003 and Zuccale Fault, Italy, Collettini & Holdsworth 2004). Heterogeneity in the core of mature, large displacement faults helps to explain the various types of seismic behaviour ranging the complete spectrum between aseismic creep, slow slip events and seismic slip (see Schwartz & Rokosky 2007 and Wibberley et al. 2008 for reviews); this assumption is supported by recent numerical and experimental models (e.g. Bos & Spiers 2001; Giger et al. 2007; Hillers et al. 2007; Ben-Zion 2008). While syntectonic chemical alteration and development of an interconnected network of weak shear planes may promote aseismic creep along mature, crustal-scale faults, in order for a fault to rupture, two conditions must be satisfied: (1) the applied shear stress must be sufficient to overcome the frictional strength – that is, the static fault friction plus any cohesion present and (2) once this is satisfied, the dynamic frictional shear resistance must also be overcome. If a fault comprises weak material everywhere, the rocks surrounding the fault cannot store sufficient elastic strain energy to drive an earthquake rupture (e.g. Reid 1910). If a rupture nucleates in a strong asperity on such a fault, the propagating

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rupture tip would granulate the weak fabric and form a material with a normal friction coefficient. Microseismicity might ensue, but a propagating rupture would halt before large strains could be accommodated. Apparent friction must be low at both static and dynamic strain rates for a large fault rupture to occur at low driving stresses. To this end, the interaction of different fault weakening mechanisms discussed above may account for the time-averaged weakness of some faults with spatially heterogeneous strength and pore fluid pressure distributions. The required conditions for a far-reaching, selfsustaining fault rupture have been investigated by He et al. (1990) and Pan et al. (2006), who found that self-sustaining failure is more likely in to occur materials with less variability in component strength. These analyses also indicate that localized areas of markedly low strength are not conducive to large-scale fault rupture. We suggest that two changes will occur with time after a fault rupture that will decrease its heterogeneity. First, tectonic stress buildup will cause progressive failure of weak force chains, which increases the minimum local strength of granular material along a fault. Secondly, stress corrosion (Sornette 1999) will progressively reduce the strength of the most-stressed force-chains, leading to a reduction in maximum strength. As shown in Figure 5, this results in a reduction in the overall strength of granular material along the fault. We believe that as time passes and stress builds up, a fault becomes progressively more likely to experience a large-scale rupture because of a reduction in the overall strength and strength heterogeneity of granular material within its core.

The role of granular material Some insight into the importance of brittle faulting in accommodating large strains on LANF comes from the work of John & Foster (1993) on the Chemehuevi low-angle normal fault, southeastern California, USA. 40Ar/39Ar and fission-track thermochronology, combined with structural data constrain the initiation angle of the regionally developed normal fault to a low angle (,308). Formation of ductile mylonites occurred prior to displacement on the Chemehuevi LANF, which always truncates high-angle normal faults in the hanging wall and appears to have accommodated 15 km displacement by brittle deformation processes such as cataclastic flow and frictional sliding in fault gouge, crush breccias, and cataclasites with rare transitional brittle–ductile protomylonite and pseudotachylyte. These observations indicate that the Chemehuevi LANF: (1) is not an exhumed brittle –ductile transition that has experienced only minor displacement (see Miller et al. 1983);

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(2) initiated and slipped at a high angle to the vertical maximum compressive stress (see Anderson 1951; Wernicke & Axen 1988); (3) was seismically active. Hayman (2006) similarly concluded that fault rocks within the Black Mountain detachments, Death Valley, USA were deformed by granular flow resulting in extreme comminution of granular gouge and strain localization. These observations are consistent with our description of granular flow with dynamic fragmentation. Unfortunately, geologists still lack criteria for distinguishing cataclasites formed at high seismic strain rates from cataclasites formed at low aseismic strain rates (Sibson 1977, 1989; Cowan 1999). Fundamentally, this inhibits our understanding of how LANFs accommodate slip (e.g. Hayman 2006; Smith et al. 2007). Imber et al. (2008) outlined experimental, geological, and seismic evidence that indicates a interconnected network of alteration minerals deforms aseismically in weak fault cores by diffusion-assisted pressure solution creep; corresponding fault rocks include phyllonite and foliated cataclasite formed at depths 3 km. This paper suggests that a better understanding of the mechanics of granular flow must be obtained before the rheological significance of fault gouge, fault breccia and cataclasite can be evaluated. According to the dynamic rock fragmentation theory, the high strain rates produced during a seismic rupture should result in more, smaller fragments being generated by each fragmentation event (e.g. Zhang et al. 2004). Thus, grain-size distributions might differ between the aseismically and seismically formed brittle fault rocks. However, given that grains in cataclasites will be very small, the recently reported alteration in fractal dimension as grain diameter decreases below about 1 mm (Heilbronner & Kuelen 2006) might obscure the distinction. Chemical alteration effects and analytical difficulties are also encountered when measuring the mineralogy and particle size distribution of very fine-grained (nanometres) material in fault gouge (e.g. Reches et al. 2007). Detailed petrographic and geochemical investigations of fault gouge, protocataclasite, cataclasite, and ultracataclasite similar to studies by Chester et al. (1985), Chester & Logan (1987), Goodwin (1999), Cladouhos (1999) and Hayman (2006) will provide more information about respective micromechanical deformation process occurring during slip on LANF. In particular, grains within fault gouge and cataclasite that have deformed via pressure-solution and dislocation creep contrast texturally with brittle fracturing mechanisms. Pressuresolution creep should lead to compaction and high pore fluid pressures through time and corollary dynamic weakening mechanisms. Petrographically, microfractures in quartz containing fluid-inclusion

planes may indicate brittle instability through dynamic fragmentation (e.g. Faulkner et al. 2006).

Mechanisms of fault weakening We have focused primarily on the mechanisms that might lead to the formation and reactivation of low angle normal faults in the seismogenic crust at depths less than c. 12 km. In the brittle crust, fault zones that contain weak materials or weak fault fabric do not violate the tenets of fault mechanics and may slide stably under very low frictional resistance (e.g. Bos & Spiers 2001; Moore & Rymer 2007; Imber et al. 2008). However, evidence for this behaviour is rare, and other weakening mechanisms must operate on faults that fail both seismically and aseismically. In the field, some geological evidence for dynamic fault weakening includes: 1. Fault-parallel veins and vein networks indicative of pore fluid overpressures (e.g. Sibson 2001); 2. fluid inclusion data indicative of coseismic variations in pore fluid pressure (e.g. Parry & Bruhn 1990); 3. pseudotachylyte formed from frictional melt (e.g. Goodwin 1999); 4. intense rock pulverization and damage asymmetry indicative of dynamic normal interface separation (e.g. Ben-Zion 2001; Dor et al. 2006); 5. intensely comminuted fault rocks showing evidence of fluidization (e.g. Hayman 2006). This paper has outlined experimental and numerical evidence for each mechanism. In our view, dynamic rock fragmentation leading to rapid comminution of fault rocks offers the best explanation available at this time for both the static and dynamic weakness of large-scale crustal rocks. Although little studied in the fault context as yet, dynamic fragmentation appears in principle capable of resolving several of the issues that other mechanisms leave unresolved.

Conclusions (1) There is reason to consider seriously the possibility that LANF can both form and rupture seismically at the low inclinations they show in outcrop. These abilities require that frictional resistance to both static rupture and subsequent slip must be less than rock friction values measured in the laboratory and at small-scale in the field. (2) The mechanisms that give rise to this reduced friction in LANF are likely to be the same as those that cause reduced friction in other largescale faults. (3) To be realistic, potential fault weakening mechanisms must explain both reduced static and

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dynamic frictional resistances to shear, and must operate predominantly at large scale. Truly aseismic faults that creep continuously at low strain rates on a spatially continuous weak fabric or mineral phase are exceptions. (4) On this basis, for faults that accommodate slip both seismically and microseismically or aseismically, dynamic rock fragmentation appears to be the most applicable mechanism so far proposed. The authors would like to thank B. Wernicke and G. Axen for thoughtful reviews of the original manuscript.

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Upper and lower crustal evolution during lithospheric extension: numerical modelling and natural footprints from the European Alps ANNA MARIA MAROTTA1*, MARIA IOLE SPALLA2 & GUIDO GOSSO2 1

Universita` degli Studi di Milano, Dipartimento di Scienze della Terra ‘A. Desio’, Sezione di Geofisica, Via Cicognara 7, I-20129 Milano, Italy

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Universita` degli Studi di Milano, Dipartimento di Scienze della Terra ‘A. Desio’, Sezione di Geologia and CNR-IDPA, Via Mangiagalli 34, I-20133 Milano, Italy *Corresponding author (e-mail: [email protected]) Abstract: When continental rifting does not develop on a stable continental lithosphere, geodynamic interpretation of igneous and metamorphic records, as well as structural and sedimentary imprints of rifting-related lithospheric extension, can be highly ambiguous since different mechanisms can be responsible for regional HT–LP metamorphism. This is the case of the European Alps, where the exposure of Variscan structural and metamorphic imprints within the present-day Alpine structural domains indicates that before the Pangaea break-up, the continental lithosphere was thermally and mechanically perturbed by Variscan subduction and collision. To reduce this ambiguity, we use finite-element techniques to implement numerical geodynamic models for analysing the effects of active extension during the Permian– Triassic period (from 300 to 220 Ma), overprinting a previous history of Variscan subduction-collision up to 300 Ma. The lithosphere is compositionally stratified in crust and mantle and its rheological behaviour is that of an incompressible viscous fluid controlled by a power law. Model predictions of lithospheric thermal state and strain localization are compared with metamorphic data, time interval of plutonic and volcanic activity and coeval onset of sedimentary environments. Our analysis confirms that the integrated use of geological data and numerical modelling is a valuable key for inferring the preorogenic rifting evolution of a fossil passive margin. In the specific case of the European Alps, we show that a relative high rate of active extension is required, associated for example with a far extensional field, to achieve the fit with the maximal number of tectonic units. Furthermore, in this case only, thermal conditions allowing partial melting of the crust accompanying gabbroic intrusions and HT– LP metamorphism are generated. The concordant set of geological events that took place from Permian to Triassic times in the natural Alpine case is justified by the model and is coherent with the progression of lithospheric thinning, later evolving into the appearance of oceanic crust.

Igneous and metamorphic records, as well as structural and sedimentary imprints of rifting-related lithospheric extension, are frequently ambiguous. This is the case with high-temperature –low-pressure (HT–LP) metamorphism, which is undoubtedly related to high thermal regimes and can be consequent to the late orogenic collapse of a collisional belt or to lithospheric thinning, thereby promoting continental rifting (e.g. Thompson 1981; Wickham & Oxburgh 1985; Sandiford & Powell 1986; Platt 1998; Beardsmore & Cull 2001). The same uncertainty holds for the normal faulting that can accommodate the lateral expansion of the thickened axial part of a mountain belt during the final stages of continental collision, when gravitational forces dominate the horizontal stress field or, alternatively, accommodate lithospheric thinning in the axial zone of continental rifting (e.g. Wernicke 1985; Molnar & Lyon-Caen 1988; Keen et al. 1989; Malavieille

1993; Vissers et al. 1995). This interpretative ambiguity is enhanced when continental rifting does not develop on a stable continental lithosphere, but follows the thermal and mechanical instabilities induced by a subduction– collision process in which the rifting precursor signals overprint the markers of a late orogenic extension. This latter case corresponds to that of the European Alps, where the exposure of Variscan structures and metamorphic imprints within the present-day Alpine structural domains indicates that before the Pangaea break-up the continental lithosphere was thermally and mechanically perturbed by Variscan subduction and collision. While the metamorphic and igneous records of Variscan orogeny are widespread in the European continental crust, a diffuse igneous activity associated with HT metamorphism accounting for a Permian–Triassic high thermal regime is peculiar to the Alpine area. The overprint of HT

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 33– 72. DOI: 10.1144/SP321.3 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Permian–Triassic evolution on the HP relics of the Variscan subduction and collision has been interpreted as induced either by late-orogenic collapse or by lithospheric extension and thinning leading to continental rifting (e.g. Lardeaux & Spalla 1991; Diella et al. 1992; Gardien et al. 1994; Ledru et al. 2001). Even the interpretation of the geodynamic environment responsible for the development of intracontinental basins hosting the Permian volcanic products allows two possible alternatives, one envisaging a strike-slip dominated regime (Arthaud & Matte 1977; Cassinis & Perotti 1994), which is compatible with the evolution of a mature collisional setting (Molnar & Lyon-Caen 1988), the other a continental rifting tectonic setting (Siletto et al. 1993; Selli 1998; Staehle et al. 2001). In both cases the continental rifting promoting Mesozoic opening of the ocean in a lithosphere thermally softened and thinned by slab break-off processes is generally accomplished in the final stages of continental collision. Trying to resolve the dualistic interpretation of the geodynamic significance of the Permian– Triassic high thermal regime, Marotta & Spalla (2007) and Spalla & Marotta (2007) used numerical models to simulate ocean subduction leading to continental collision, lithospheric detachment and subsequent gravitational thermal relaxation. Comparing predictions from each model with natural data, they concluded that the Permian–Triassic metamorphic and igneous imprints in the Alps cannot simply result from the thermal relaxation consequent to lithospheric unrooting during late orogenic extension, and that a supplementary mechanism, such as hot mantle upwelling under continental plates, was necessary to satisfy the Permian–Triassic geothermal state required to account for the natural metamorphic data. Indeed, the lithospheric thermal detachment was premature (360 Ma) with respect to natural thermal signals (290– 225 Ma). Our goal here is to expand the preliminary analysis by Marotta & Spalla (2007) on the effects of a forced extension during the Permian–Triassic period (from 300 to 220 Ma) via new models characterized by extension velocities varying from no extension to 2.0 cm a21. The initial configuration corresponds to the geodynamic setting at 300 Ma after simulation of Variscan subduction–collision in Marotta & Spalla (2007). Model predictions of lithospheric thermal state and strain localization at different structural levels are compared to a broader set of metamorphic data corroborated by the emplacement conditions of mafic intrusions, time interval of plutonic and volcanic activity and coeval onset of sedimentary environments. It is worth noting that the present study, however, recovers natural data from a fragmented setting carrying reminiscences of rifted margins

after the subduction and collision of the Alpine orogeny, whose hinterland and foreland obviously preserve the rifting imprints better. Orogenic fragmentation of the rifted margins greatly complicates the heterogeneous structural patterns inferred from seismic exploration of present-day passive margins (e.g. Boillot et al. 1995) and, hence, similarly reproduced by analogue modelling (e.g. Brun & Beslier 1996). The integrated use of natural data and numerical modelling may help to infer the puzzled rifting evolution even in a fossil passive margin.

Geological framework Alpine evolution During Alpine convergence, the involvement of both continental margins in the subduction zone is responsible for the fragmentation of metamorphic and igneous imprints left after the Pangaea breakup on both passive margins. The Alps are constituted of nappes grouped in different structural domains according to their relative position in the present-day structural setting, which is usually interpreted as reflecting a commonly accepted palaeogeography. The different nappes have been contoured by taking into account lithostratigraphic setting (ocean vs. continent), tectonic style (infrasupra crustal) and metamorphic histories (Spalla et al. 1996 and references. therein). From the external to the internal part of the chain (Fig. 1), the European Foreland and the Helvetic, Penninic, Austroalpine and Southalpine domains are distinguished. The European Foreland was downbent and underthrust at lithospheric scale during late Alpine convergence. The Helvetic domain, a thick-skinned thrust system of basement and cover slices broadly preserving pre-Alpine metamorphic and stratigraphic signatures, results from Tertiary reactivation by tectonic inversion of normal faults that dissected the European passive margin. The Penninic and Austroalpine nappe system constituting the axial part of the belt is highly heterogeneous and has been deformed and metamorphosed since the Cretaceous during oceanic subduction and continental collision; the Penninic presently consists of a melange of thin continental and oceanic basement and cover slices (nappes); the Austroalpine is fully continental material. Units with oceanic affinities derive from the sutured neoTethyan oceanic domain. The Southalpine domain represents the hinterland of the early Alpine belt and has evolved since the Cretaceous as a S-verging, thick-skinned thrust system of basement and cover units only locally affected by very lowgrade Alpine metamorphism (e.g. Brack 1981; Milano et al. 1988; Polino et al. 1990; Schmid et al. 2004). Seismic profiles synthesized after

LITHOSPHERE EXTENSION: MODELLING V. DATA 35

Fig. 1. Tectonic map of the Alps with the location of Permian –Triassic metamorphic rocks occurring in the pre-Alpine continental crust (black dots) and the main Permian –Triassic gabbro bodies emplaced in the pre-Alpine continental crust (black stars). Ca, Dol, Avp, Obb and Lv labels indicate the location of structural, volcanic and sedimentary events listed in Table 3. Keys are coded in Tables 1 and 2 and are the same used in Figures 2 and 4. Legend: (1) Southalpine basement, (2) Austroalpine basement, (3) Penninic basement, (4) Helvetic basement, (5) Tertiary intrusive stocks.

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A. M. MAROTTA ET AL.

the ECORS–CROP – NRP20– TRANSALP project (Cassinis 2006 and references therein) illustrate the tectonic framework at the lithosphere scale in which the axial part of the chain is occupied by a Cretaceous –Palaeogene rootless crustal prism between the Penninic Front and the Periadriatic Lineament (PF and PL, respectively, in Fig. 1) that bounds it towards the Helvetic (European Plate) and Southalpine (Adria Plate) domains (e.g. Polino et al. 1990; Platt 1993; Dal Piaz et al. 2001). The axial crustal prism comprises units recording high- to ultra-high-pressure metamorphism of the Cretaceous to Late Eocene (Handy & Oberhaensli 2004 and references therein). This large time interval of eclogite formation in rocks of both continental and oceanic origin suggests that they may have been generated during either continental collision or subduction of the European oceanic lithosphere (lower plate) accompanied by tectonic erosion of the Adria active margin before the onset of continental collision (e.g. Platt 1986; Polino et al. 1990; Spalla et al. 1996).

Permian – Triassic evolution Although a large amount of continental lithosphere from both European and Adriatic continental margins was absorbed in the sub-lithospheric mantle during Alpine subduction, relict metamorphic and igneous imprints of the Variscan convergence and the successive Pangaea break-up survived in the pre-Alpine continental crust. They are preserved either in small volumes within the exhumed continental crust slivers of the axial belt or in the Alpine hinterland and foreland crusts that were never reworked in deep-seated conditions. Because of the common occurrence of metamorphic and igneous markers of the Variscan convergence in the preAlpine continental crust, the Permian–Triassic high thermal regime, which has left widespread metamorphic and igneous signatures, has been interpreted via two possible geodynamic scenarios: the late-orogenic collapse of the Variscan belt enhanced by lithospheric unrooting (e.g. Malavieille et al. 1990; Gardien et al. 1994; Cortesogno et al. 1998; Schaltegger & Brack 2007), or lithospheric thinning leading to continental rifting (Handy & Zingg 1991; Lardeaux & Spalla 1991; Diella et al. 1992; Schuster et al. 2001; Thoeni 2003). In addition, even Permian volcanic activity and related basin formation have been interpretatively linked to dextral strike-slip tectonics associated with pullapart basin formations (Cassinis & Perotti 1994; Muttoni et al. 2003; Cassinis et al. 2007), or to lithospheric extension leading to the Pangaea break-up and subsequent oceanization (Selli 1998; Staehle et al. 2001). If the Carboniferous/Permian transition of the Palaeozoic plate convergence to a

transtensional –extensional tectonic regime announcing the Pangaea break-up (e.g. Golonka et al. 1994) is taken into account, the latter scenario appears to be the more likely. Pull-apart basins indicate that the completion of the thinning process of thickened Variscan crust is occurring at this time and that the intense Late Palaeozoic magmatism thus represents the precursor stage of Mesozoic rifting (Wopfner 1984; Ziegler 1993) and predates the marine transgression from the east, where the Neotethys Ocean is opening (Muttoni et al. 2003). Metamorphism. Permian–Triassic HT–LP metamorphic imprints are recorded in Penninic and Austroalpine continental units and in the metamorphic basement of the Southalpine hinterland but are not known in the Helvetic Domain (Table 1 and Figs 1 & 2), where on the contrary a polyphase Variscan structural and metamorphic evolution is recorded (Fig. 3). Permian–Triassic metamorphic signatures have been widely recognized in lower, intermediate and upper continental crust of Austroalpine and Southalpine Domains, but only locally in the upper and intermediate Penninic crust of the Western Alps (Figs 1 & 2, Table 1). PT evolutions have no peculiar character in either structural domain, as it is the case for the lithostratigraphy of tectonic units recording the HT Permian–Triassic imprints. Ages inferred for peak temperature conditions are scattered between 220 and 295 Ma (Table 1), even if mineral ages up to 170 Ma (Fig. 3) are obtained where the metamorphic imprints recorded during uplift and associated with fluid circulation are dominant (Bertotti et al. 1993; Biino & Meisel 1996; Schuster et al. 2001; Pinarelli & Boriani 2007); the exhumation paths are usually characterized by a high T/P ratio. In the data review of Table 1 and Figure 2 only PT estimates related to Tmax conditions are reported to better highlight the Permian–Triassic thermal regime. In Figure 2, Tmax  PT max estimates from Penninic, Austroalpine and Southalpine domains plot above the maximally relaxed geotherm (England & Thompson 1984; Thompson & England 1984) in a field characterized by P/T ratios that are compatible with extensional tectonics associated with astenosphere upwelling. In the Penninic domain, HT assemblages occur in sillimanite-bearing metapelites and in metaintrusives; exhumation can occur following a P-retrograde path characterized by cooling (Bouffette et al. 1993) or by heating (Desmons 1992). Permian–Triassic metamorphism of the Austroalpine domain is imprinted in sillimanite and biotite-bearing metapelites, with associated orthopyroxene- and garnet-bearing mafic granulites, amphibolites and high-grade marbles.

Table 1. Assemblages, rock types and physical conditions of Permian– Triassic metamorphism recorded in the continental crust of the Alps and associated with the Permian –Triassic magmatism Key

Tectonic unit location

Penninic tectonic system 1 Briancon basement Ruitor 2 Monte Rosa 3

Dora Maira

Lithologies

T (K)

P (gpa)

Age (Ma)

And-bearing metapelites

Metapelites

723– 823

0.1–0.3

Grt þ Sil þ Bt þ Kfs + Ms Grt þ Sil þ Bt þ Pl þ Qtz

Metapelites

–

–

Metapelites

923– 1023

0.4–0.7

pre-Alpine (295– 245)

Opx þ Pl þ Grt þ Qtz þ Hbl Sil þ Bt þ Crd þ Pl þ Qtz Opx þ Pl þ Grt þ Qtz þ Hbl Sil þ Bt þ Pl þ Qtz Hbl þ Plg þ Qtz Grt þ Sil þ Bt þ Pl þ Ms Opx þ Pl þ Grt þ Qtz þ Hbl Sil þ Bt þ Crd þ Pl þ Qtz Wm þ Chl þ Ilm + Ttn

Basic granulites

973– 1073

0.7–0.9

Acidic granulites

973– 1023

0.6–0.8

Basic and acidic granulites

973– 1023

0.6–0.7

Permian? (295– 245) Permian? (295– 245) Permian? (295– 245)

Metabasics Metapelites

823– 923 783– 853

0.3–0.45 0.25–0.45

Basic and acidic granulites

1023–1073

Metagranitoids, metadiorites Metapelites

6 7

Mt Emilius Klippe Mont Mary Nappe

8

Dent Blanche Nappe (Valpelline)

9 10

Dent Blanche Nappe (Arolla) Silvretta (Pischahorn)

11a

Languard–Campo

Sil þ Opx þ Kfs þ Bt þ Qtz

11b

Languard–Campo

Crd þ Bt þ Grt þ Sp þ Sil þ Qtz

12

Languard–Campo

13

Matsch Nappe

14

Uttenheim Ahrntal

Sil þ Bt þ Grt þ Cd þ Pl þ Qtz Hbl þ Grt þ Cpx þ Pl þ Qtz Grt þ Sil/And þ Bt + Crd þ Pl þ Qtz Grt þ Bt þ Sil Pl þ Qtz þ L

Qtz þ Ms þ And

Permian (295– 245) 250– 280

Method Rb/Sr K/Ar U/Pb

Refs Desmons (1992); Bocquet et al. (1974) Engi et al. (2001); Dal Piaz (2001) Bouffette et al. (1993)

Lardeaux & Spalla (1991) Lardeaux & Spalla (1991) Lardeaux (1981); Vuichard (1987); Biagini et al. (1995) Dal Piaz et al. (1983) Pennacchioni & Cesare (1997) Nicot (1977); Hunziker et al. (1992); Gardien et al. (1994)

0.5–0.7

280? Permian? (295– 245) 180

K/Ar

619– 711

0.3–0.6

.290

U/Pb

c. 873

c. 0.2

K/Ar

Granulites

843– 1023

0.4–0.6

.280 (295– 280) c. 290

Granulites – contact metamorphism Metapelites and metabasics

1123–1223

0.4–0.6

c. 290

Sm/Nd

923– 1093

0.5

260– 280

Rb/Sr

Spalla et al. (1995); Zucali (2001)

Metapelites

843– 913

0.3–0.55

290 + 17

Rb/Sr

Metapelites

893– 953

0.5–0.7

262 + 7 253 + 7

Rb/Sr Sm/Nd

Gregnanin (1980); Haas (1985) Borsi et al. (1980); Stoeckhert (1987); Schuster et al. (2001)

SmNd

Roda & Zucali (2008); Bussy et al. (1998) Brugger (1994); Maggetti & Flish (1993) Giacomini et al. (1999); Tribuzio et al. (1999); Schuster et al. (2001) Gosso et al. (1995); Tribuzio et al. (1999)

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(Continued)

LITHOSPHERE EXTENSION: MODELLING V. DATA

Austroalpine 4a Sesia Lanzo Zone lower element 4b Sesia Lanzo Zone lower element 5 Sesia Lanzo Zone upper element

Assemblages

Table 1. Continued Tectonic unit location

15 16a

Strieden Kreuzeckgruppe Woelz Complex

16b

Woelz Complex

17

Sopron

18

Assemblages

Lithologies

T (K)

P (gpa)

Age (Ma)

Method

Refs 38

Key

Sil þ Bt þ Pl þ Qtz þ L

Metapelites

873– 1023

0.3–0.5

Gneiss

713– 793

0.2–0.4

Metapelites

788– 828

0.35–0.45

Metapelites

848– 973

0.18–0.38

Permian (295– 245) 300 + 40

Raabalpen Complex

Grt þ Chl þ Ms/Pg þ Ab þ Qtz + Bt + Mrg Grt þ Bt þ Ms þIlm/ Rt þ Pl þ Qtz Bt-And-Sil-bearing schists Zirc þ Zrn þ Dol

773– 873

0.2

284 + 4

19a

Saualpe– Koralpe

Grt-bearing micaschists

823– 833

0.52–0.71

240– 290

19b

Saualpe– Koralpe

848– 888

0.3–0.5

249 + 3

19c

Saualpe– Koralpe

Grt + Bt + Ms + Pl + Sil + Qtz Grt + Bt + Ms + Pl + Sil + Qtz

Metacarbonates – contact metamorphism Metapelites and pegmatites Metapelites Metapelites

923– 943

0.4–0.9

Permian

Tenczer et al. (2006)

Metapelites – contact metamorphism Dolomitic limestones – contact metamorphism Metapelites and metabasics

902

0.26

282

Visona` (1995); Benciolini et al. (2006)

1023–1073

0.1

225– 234

Rb/Sr

Borsi et al. (1968); Povoden et al. (2002); Gallien et al. (2007)

923– 1022

0.4–0.5

224– 227

Rb/Sr

–

288 + 99 (Permian)

Rb/Sr best-fit

Diella et al. (1992); Bertotti et al. (1993); Sanders et al. (1996); di Paola & Spalla (2000) Boriani & Burlini (1995); Pinarelli & Boriani (2007) Hunziker & Zingg (1980); Brodie et al. (1989); Quick et al. (1992); Vavra et al. (1996); Colombo & Tunesi (1999) Henk et al. (1997); Vavra et al. (1996); Colombo & Tunesi (1999)

Crd þ Sil þ Bt

21

Monzoni complex

Grt þ Cpx-bearing dolomitic limestones

22

Dervio Olgiasca Zone

Bt þ Sil þ Pl þ Qtz + Grt + Kfs Amp + Cpx þ Pl þ Qtz +Bt

23

Strona– Ceneri Zone

Sil-, And-, Crd-bearing metapelites

Metapelites

–

24a

Ivrea Zone

Sil þ Bt þ Grt þ Crd þ Pl þ Qtz

Metapelites

953– 1053

0.45–0.65

250– 290

24b

Ivrea Zone

Grt þ Opx þ Hbl þ PlþQtz

Metabasics

1023–1223

0.8–0.9

273– 296

Sm/Nd Rb/Sr

Hoke (1990); Schuster et al. (2001) Schuster & Frank (2000) Gaidies et al. (2006)

Th– U– total Pb Sm– Nd Rb/SrU/ PbSm/Nd Sm/Nd

Nagy et al. (2002) Tropper et al. (2007) Thoeni & Miller (2000) Habler & Thoeni (2001)

See Table 2 for gabbro distribution and Table 3 for volcanic records. Radiometric estimates are expressed with the error bar. When the ‘Method’ column is empty, the age indication is based on geological constraints and expressed in brackets. Keys correspond to those of Figures 1 and 2, where Tmax –PTmax values are represented. Mineral abbreviations are after Kretz (1983).

A. M. MAROTTA ET AL.

Southalpine 20 Eisacktal

261 + 3 229 + 3 220– 260

LITHOSPHERE EXTENSION: MODELLING V. DATA

39

Fig. 2. PT estimates of the thermal peak of Permian– Triassic metamorphic rocks in Penninic, Austroalpine (Western Alps and Eastern-Central Alps) and Southalpine Domains. Ages, PT estimates, lithotypes and references are reported in Table 1. Keys are coded as in Table 1 and the corresponding samples are located in Figure 1. Metamorphic facies fields are redrawn after Spear (1993): Z, zeolite facies; PP, prehnite–pumpellyite facies; Gs, greenschist facies; Bs, blueschist facies; E, eclogite facies; EA, epidote–amphibolite facies; A, amphibolite facies; G, granulite facies. Maximally relaxed geotherm (RG) is redrawn after England & Thompson (1984) and Thompson & England (1984). Depth– pressure equivalence in km is inferred on the basis of P ¼ rgz, with r ¼ 3000 kg . m23.

HT assemblages mark newly differentiated foliations in places associated with discrete shear zones, especially in metapelites (Lardeaux & Spalla 1991; Spalla et al. 1991). Locally, a very HT metamorphic aureole developed in the gabbro country rocks (#11b in Table 1) and an intermediate T assemblage characterizes the aureole of granites (#18 in Table 1). Exhumation paths of these units may be characterized by cooling (e.g. Dal Piaz et al. 1983; Stoeckhert 1987; Vuichard 1987; Lardeaux & Spalla 1991), isothermal decompression (e.g. Spalla et al. 1995; Zucali 2001) or by heating (e.g. Schuster et al. 2001). Large parts of the uplift paths occurred under high thermal regime and the exhumation of some of the deepseated continental crust units occurred up to shallow crustal levels, thus suggesting that some Austroalpine units were part of a thinned continental margin before being subducted during Alpine convergence (e.g. Rebay & Spalla 2001).

Southalpine metapelites, mafic granulites, amphibolites and high-grade marbles are the rocks recording Permian–Triassic HT metamorphic imprints that are frequently associated with the development of pervasive foliations, sometimes mylonitic (e.g. Gosso et al. 1997). Contact aureole developed in country rocks where magma emplaced at shallow levels in basement rocks during the Permian (#20 in Table 1) or in dolomic limestones during Lower Triassic (#21 in Table 1) (Povoden et al. 2002; Benciolini et al. 2006; Gallien et al. 2007). Uplift trajectories were characterized by decompressional cooling (e.g. Brodie et al. 1989), or by increasing temperature during decompression (e.g. di Paola & Spalla 2000), and always characterized by a high T/P ratio. Magmatism. Together with the Permian–Triassic HT –LP metamorphism there is a widespread igneous activity that is characterized by both intrusive

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A. M. MAROTTA ET AL.

Fig. 3. Timing of metamorphic and magmatic events from Middle Variscan to Jurassic along the Alpine belt. Radiometric ages are plotted including analytical uncertainty ranges. H, Helvetic; P, Penninic; A, Austroalpine; S, Southalpine. Variscan evolutions are synthesized by von Raumer & Neubauer (1993), Frey et al. (1999), Marotta & Spalla (2007) and Spalla & Marotta (2007); data on gabbros are listed in Table 2. Age data on pegmatites after Staehle et al. (1990), Diella et al. (1992), Sanders et al. (1996) and Thoeni (2003). Volcanic occurrences are summarized in Table 3, age data, comprising radiometric and stratigraphic dating, are from Brusca et al. (1981), Schuster et al. (2001), Staehle et al. (2001), Schaltegger & Brack (2007), Visona` et al. (2007). Age data on metagranitoids are after Boriani et al. (1985), Bonin et al. (1993), Rottura et al. (1998), Liati & Gebauer (2003) and Pinarelli & Boriani (2007). Tmax estimates for Permian– Triassic metamorphic imprints and related ages are listed in Table 1. This metamorphic evolution recorded in the pre-Alpine continental crust is polyphase and the last stages of metamorphic re-equilibrations are responsible for age rejuvenation up to 190–170 Ma (Diella et al. 1992; Biino & Meisel 1996; Schuster et al. 2001; Pinarelli & Boriani 2007). Ophiolites age data are synthesized from Costa & Caby (2001), Masson (2002) and Stucki et al. (2003). The light grey stripe locates age interval of Permian gabbro emplaced in the continental crust of the Iberian passive margin (He´bert et al. 2008); the dark grey stripe indicates the age of the earliest radiolarian cherts related to Alpine ophiolites (Cordey & Bailly 2007).

rocks, ranging in composition from gabbro to granite, and large volumes of volcanic rocks, with compositions from basalts to rhyolite (Bonin et al. 1993; Rottura et al. 1998; Staehle et al. 2001). While the metamorphic and igneous records of the Variscan cycle (425 –295 Ma) occur in the continental crust from Helvetic to Southalpine domains, the Permian–Triassic magmatism (Figs 1, 3, 4 and Table 2) and metamorphism did not affect the Helvetic domain; available age data on these magmatic rocks significantly postdate late Variscan metamorphism, structures and magmatic activity (Fig. 3). The mafic igneous products are a peculiar

character of the Alpine continental crust with respect to the rest of the European Variscan chain and mainly occur in the Austroalpine–Southalpine domain (Fig. 1). They consist of huge gabbro bodies (Figs 1, 4 and Table 2) that are frequently associated with sub-continental peridotites (e.g. Brodie et al. 1989; Bonin et al. 1993; Schuster et al. 2001; Staehle et al. 2001; Rampone 2002; Spalla & Gosso 2003). Gabbro emplacement occurred at different crustal levels as documented by PT estimates inferred from igneous, early subsolidus recrystallizations and contact metamorphic assemblages (Fig. 4). The country rocks comprise

LITHOSPHERE EXTENSION: MODELLING V. DATA

41

Fig. 4. PT conditions of emplacement of Austroalpine and Southalpine Permian– Triassic gabbros inferred from PT estimates on magmatic assemblages, on subsolidus recrystallization assemblages developed in the contry rocks during contact metamorphism or on regional geological constraints. Depth–pressure equivalence in km is inferred on the basis of P ¼ rgz, with r ¼ 3000 kg . m23. Ages and keys are listed in Table 2, the corresponding samples are located in Figure 1. Metamorphic facies fields are redrawn after Spear (1993); metamorphic facies as in Figure 2. Data sources: Austroalpine domain: 1, Corio and Monastero (Rebay & Spalla 2001); 3, Matterhorn Collon (Monjoie 2004); 4, Fedoz-Braccia (Muentener et al. 2000); 5, Sondalo (Gosso et al. 1995); 6, Baerofen and Gressenberg (Miller & Thoeni 1997). Southalpine domain: 7, Bressanone–Chiusa dioritic belt (Visona` 1995); 8, Monzoni (Povoden et al. 2002); 9, Val Biandino (De Capitani et al. 1988); 10, Ivrea Mafic Complex (Peressini et al. 2007); 11, Ivrea Finero body (Sills 1984).

a wide selection of lithologies ranging from HT–IP metamorphics (granulites: Sills 1984; Handy & Zingg 1991; Lardeaux & Spalla 1991) to consolidated metasediments (Gallien et al. 2007 and references therein), accordingly with the structural level of the emplacement, from lower to upper crust. Rocks with HT metamorphic imprints recorded at intermediate and lower crustal levels are located in the surroundings of: the Corio and Monastero gabbros (#g1 in Table 2 and Fig. 4) as acidic and basic granulites (#4 in Table 1) in the Sesia Lanzo zone; the Dent Blanche gabbros (#g3 in Table 2 and Fig. 4) as acidic and basic granulites; the Sondalo gabbro (#g5 in Table 2 and Fig. 4) as granulites (#11 in Table 1); the Baerofen gabbro (#g6 in Table 2 and Fig. 4) as HT metapelites (#19 in Table 1); and the Ivrea gabbros (#g10 in Table 2 and Fig. 4) as granulitized metabasic and metapelites (#24 in Table 1). Upper crustal level country rocks, with contact metamorphic imprints of LP, are situated in the surroundings of: Bressanone– Chiusa gabbros (#g7 in Table 2 and Fig. 4) and diorites as LP amphibolite-facies metapelites (#20 in Table 1); and gabbros from the Monzoni complex (#g8 in Table 2 and Fig. 4) as metamorphosed Lower Triassic dolomitic limestones (#21 in Table 1). Different structural levels of emplacement are not univocally related to specific time intervals. In the Penninic domain the magmatic products vary from dioritic –granodioritic to granitic compositions, and radiometric ages range between 240 and 290 Ma (Bonin et al. 1993; Romer et al. 1998; Liati & Gebauer 2003; Bertrand et al. 2005). Granitoid bodies diffusedly intruded the Austroalpine and

Southalpine crust in the same time interval (Rottura et al. 1998; Schuster et al. 2001; Schaltegger & Brack 2007). The magmatic signature varies from alkaline/sub-alkaline to calc-alkaline. The calcalkaline affinity of part of the Permian magmatism is thought to derive from crustal contamination of basaltic magmas generated within enriched lithospheric and/or asthenospheric mantle sources. This interpretation indicates an important role of the mantle in the generation of the Permian magmatic products (e.g. Rottura et al. 1998; Voshage et al. 1990) in a context where lithospheric extension and attenuation favoured simultaneous production of lithospheric and/or asthenospheric magmas (e.g. Cortesogno et al. 1998). Pegmatite emplacement spreads on a time interval between 290 and 220 Ma in the Southalpine and Austroalpine continental crust and clusters at 225 Ma in Southalpine domain (Ferrara & Innocenti 1974; Sanders et al. 1996; Schuster et al. 2001; Thoeni 2003), where some compositions in the Ivrea Zone are peculiar (alkaline syenites and carbonatites; Staehle et al. 1990). During the Triassic, the main signal is recorded in the Southalpine domain, where alkaline rocks are emplaced into the Ivrea Zone; the intrusive complex of Mt Monzoni was emplaced at shallow depth in the eastern Southalpine crust and the rest of the domain was dominated by volcanic activity associated with large-scale tuff deposition. Staehle et al. (2001) interpret this igneous activity as the result of a long-lasting process of active rifting. The general picture emerging from the emplacement of Permian– Triassic intrusives and their exhumation, together with that of the coeval metamorphic rocks, requires a thermal regime

42

Table 2. Permian –Triassic gabbros emplaced in the pre-Alpine continental crust of the Alps Key

Location

Material

Gabbro-norite

Method

Age (Ma)

Gabbro Gabbro

Zrn Phl

Geological evidence U/Pb K/Ar Rb/Sr

Gabbro

Zrn

U/Pb

284 + 0.6

Prg

Ar/Ar

c. 260

Zrn Zrn

U/Pb U/Pb

266 – 276 281 + 19 281 + 2 242 + 4 266 + 10 300 + 12 269 + 16 280 + 10 275 + 18 261 + 10 247 + 16 255 + 9

g4 g4

Fedoz-Braccia Fedoz-Braccia

Mafic dykes (alkaline lamprophyres) Gabbro Gabbro

g5 g5

Sondalo Sondalo

Gabbro (?) Troctolite

Bt Pl-Amp-Cpx-WR

g5

Sondalo

Norite

Pl-Amp-WR

g6 g6 g6

Baerofen Baerofen Baerofen and Gressenberg

Gabbro Gabbro Eclogitized gabbro

Cpx, Pl, WR Pl-Cpx

Rb/Sr Rb/Sr Sm/Nd Rb/Sr Sm/Nd Sm/Nd Sm/Nd Sm/Nd

Permian?

Rebay & Spalla (2001)

288þ2 24

Bussy et al. (1998) Dal Piaz et al. (1977)

250 + 5

Southalpine g7 Bressanone –Chiusa dioritic belt g8 Monzoni

Gabbro-norite

Bt

Rb/Sr

276 + 4

Gabbro

Bt

Rb/Sr

225 – 234

g8

Gabbro, diorite

Zrn Bt, Amp

U/Pb Ar/Ar

232 – 238

Predazzo

Refs

Monjoie et al. (2004); Monjoie et al. (2005) Monjoie et al. (2004); Monjoie et al. (2005) Muentener et al. (2000) Hansmann et al. (2001); Hermann & Rubatto (2003) Del Moro in Boriani et al. (1985) Tribuzio et al. (1999) Tribuzio et al. (1999) Thoeni & Jagoutz (1992) Thoeni & Jagoutz (1992) Miller & Thoeni (1997)

del Moro & Visona` (1982); Visona` (1995) Borsi et al. (1968); Povoden et al. (2002) Mundil et al. (1996); Visona` (1997); Ferry et al. (2002)

A. M. MAROTTA ET AL.

Austroalpine g1 Sesia Lanzo – Corio and Monastero g2 Sesia Lanzo –Sermenza g3 Dent Blanche –Matterhorn Collon g3 Dent Blanche –Matterhorn Collon g3 Dent Blanche –Mont Collon Dents de Bertol

Lithologies

g9 g10 g10

g10 g10 g10 g10 g10 g10 g11 g11

Gabbro-diorite Diorite

WR Zrn

Rb/Sr U/Pb

279 + 5 285þ7 25

Thoeni et al. (1992) Pin (1986)

Diorite

Bt

Ar/Ar

211 + 3

Buergi & Kloetzli (1990)

Gabbro, diorite

Zrn

U/Pb

Garuti et al. (2001)

Diorite and gabbro

Zrn

U/Pb

287 + 3 292 + 4 288 + 3

Peressini et al. (2007)

Gabbro

Grt-WR

Sm/Nd

271 + 22

Voshage et al. (1987)

Gabbro

Grt-Pl-WR

Sm/Nd

248 + 8

Voshage et al. (1987)

Gabbro

Cpx, Opx, Pl

Sm/Nd

274 + 11

Mayer et al. (2000)

Gabbro

Amp, WR

Sm/Nd

267 + 21

Mayer et al. (2000)

Metabasics

Zrn

U/Pb

293 + 12

Vavra et al. (1999)

Gabbro

Grt, Cpx, Pl, Amp

Sm/Nd

Lu et al. (1997)

Gabbro

Grt, Cpx, Pl, Amp

Sm/Nd

231 + 21 223 + 10 215 + 15

Lu et al. (1997)

As in Table 1, radiometric estimates are expressed with the error bar (the dating method is specified in the column Method). Age indications are based on geological grounds and in brackets and the column Method is empty. Keys correspond to those of Figures 1 and 4, where the PT conditions inferred for gabbros emplacement are represented.

LITHOSPHERE EXTENSION: MODELLING V. DATA

g10

Val Biandino Ivrea–Upper Mafic Complex–Val Mastallone Ivrea–Upper Mafic Complex–Val Sesia/Val Mastallone Ivrea–Upper Mafic Complex Ivrea–Upper Mafic Complex–Val Mastallone and Val Sesia Ivrea–Upper Mafic Complex–Valbella Ivrea–Upper Mafic Complex–Sassiglioni Ivrea- Lower Mafic Complex–Val Sesia Ivrea- Upper Mafic Complex–Val Sessera Ivrea–Upper Mafic Complex Ivrea–Finero mafic ultramafic body Ivrea–Finero mafic ultramafic body

43

44

A. M. MAROTTA ET AL.

compatible with extensional tectonics promoting asthenosphere upwelling (Fig. 2). The continental lithosphere thinning was accommodated by ductile shear zones developed in metamorphic conditions ranging from upper amphibolite– granulite to greenschist facies (Brodie et al. 1989; Diella et al. 1992; Bertotti et al. 1993). The concentration of mafic intrusions in the Austroalpine and Southalpine basements (Fig. 1 and Table 2) strongly suggests that the rifting was asymmetric, in agreement with asymmetric rifting models (e.g. Wernicke 1985; Lister et al. 1986), and that the continental crust belonging to these two domains acted as a hanging wall during an asymmetric lithospheric thinning which, according to many authors, is a precursor of rifting (e.g. Lardeaux & Spalla 1991; Diella et al. 1992; Dal Piaz 1993). The continuous age overlap between Permian–Triassic igneous and metamorphic ages and ophiolite formation (Fig. 3) suggests that a divergent regime gradually evolved to oceanization. Ocean formation had already started in the Middle Jurassic (Masson 2002 and references therein) on the base of mid-ocean-ridge basalt, and radiolarian cherts ages (Cordey & Bailly 2007) that are in agreement with Late Triassic –Liassic ages (e.g. Borsi 1995; Costa & Caby 2001) obtained from the Chenaillet and Voltri ophiolitic complexes. The Permian– Triassic lithospheric thinning here is thought to be the precursor of the opening of the western Tethys ocean; this is also supported by new dating of gabbros and tonalites underplating the thinned continental crust of the Galicia margin at 270 –240 Ma (He´bert et al. 2008), making the evolution of the fossil passive margins of the Alps, as here inferred, strikingly similar to that of a present-day passive margin. Sedimentary basins and volcanics. Geological signatures of tectonic activity imprinted in upper crustal levels during the time interval from Late Carboniferous to Permian and Middle Triassic are well known on both sides of the Alpine belt. Updated syntheses of stratigraphic data and absolute age estimates on volcanics by Wopfner (1984), Banzet et al. (1985), Doglioni (1987), Cassinis et al. (1988), Venturini (1991), Cassinis & Perotti (1994, 2007), Barfe´ty et al. (2001), Capuzzo & Bussy (2001), Strzerzynsky et al. (2005), Schaltegger & Brack (2007), Spalla et al. (2007) and Visona` et al. (2007) tried to fit in time and space sedimentary environment changes, breaks in deposition continuity or abrupt lateral thickness variations of deposits. Such records help to individuate from facies distribution the main faults at basin scale that may simultaneously account for coeval igneous advection to the surface. Approximate ages are reported in Table 3 for the Southalpine and part of

the Helvetic tectonic domains; discrepancies within published data depend on the method used (see cited papers for details and difficulties encountered in regional litho- and chronostratigraphic correlation). In the time interval examined in the present study (300 –220 Ma, Late Gzhelian to mid-Carnian), two main volcanic cycles in two age intervals are manifest beside repeated and randomly distributed volcanic pulses as dykes, pyroclastics and flows in the sedimentary record of the upper crust. The first emplaced within Lower Permian continental clastic deposits at about 280 Ma (rhyolites and minor andesites, Bergamask– Brescian Alps); recently much earlier ages (290 and 278 Ma for andesites; Visona` et al. 2007) were put forward for the northern Dolomites. The second is Early Triassic, localized in shallow water to basin deposits and spans from Anisian to Early Carnian (andesites and basalts, Dolomites, Carnia and Brescian Alps). On the external part of the Alpine belt an earlier volcanic cycle took place from Late Carboniferous to Early Permian, and trends of associated sedimentary basins seem to reproduce directions registered in Carnia. The far Eastern Alps display the most continuous Carboniferous-Permian –Triassic tectonosedimentary record of the Southalpine domain in the Carnic sector across the Italian-Austrian border. Three changes appear in the sequence across the region from the eastern Dolomites to the Carnic Alps as an effect of alternating transpressional and transtensional regime and fault-controlled groundlevel fluctuations (Venturini 1991). This evolution is compatible either with a long-lasting, large-scale dextral strike-slip faulting reactivating Hercynian orogenic trends (Matte 1986; Massari 1988), or with initiation of a rifting process (Wopfner 1984). Mid-Permian time appears to be dominated by a wide-scale horizontal dextral megashear corresponding to the transition from Pangaea B to A and the opening of the Eastern Neo-Tethys (Muttoni et al. 2003, 2005), an event that may have easily linked together such geologic processes as acrossmantle to upper-crustal pathways over a relatively broad time span. The Permian and Triassic sequences of the Dolomites region (west of Carnic Alps) are regarded as the record of rifting tectonics (Bosellini 1965, 1973; Winterer & Bosellini 1981) and, similarly, contain volcanic products (dykes and flows); Doglioni (1987) found links of Middle Triassic volcanic activity (dykes) with a dome structure and linear trends (Table 3) and episodic sinistral strike-slip motions. Further west, the Athesian Volcanic Platform (Adige/Etsch valley and Dolomites region), the largest superficial igneous complex of volcanics and subvolcanics of calc-alkaline affinity, gives earlier absolute ages, thereby adding pinpoints to the scant fossil remnants of the associated continental clastic deposits (Poli 1997; Visona` et al.

Table 3. Summary of structural, volcanic and sedimentary events recorded from Late Carboniferous to Middle Triassic in the Southalpine (from the East to the West) and Dauphinois –Helvetic domains Geographic and tectonic unit location (see Fig. 1)

Formation or group

Basin trend (present orientation)

Volcanics

Structural trend

Fluvial deltaic to shallow marine

Auernig, Bombaso

N30, N120

N120

Eastern Alps, (E-DolomitesCarnia) Eastern Alps (DolomitesCarnia)

Continental to coastal marine, rare slope dep. Alluvial-deltaic fans

Tarvisio Breccia, Sesto Conglomerate Breccia di Ugovizza

ESE

N120

Dolomites

Continental, shallow marine

Val Gardena Sdts., Bellerophon, Scythian Werfen

N– S

Carbonate platforms, basinal sequences

Contrin, Livinallongo, Wengen, S. Cassiano

N70, N –S

Ignimbrites, tuffs, dykes

N70 mid-Trias

Basalts, andesites, basaltic andesites

Trentino -W Dolomites

Carbonate platforms, basinal sequences

Contrin, Livinallongo, Wengen, S. Cassiano

N60, N120, N300, N45, N180

Basalts, andesites, basaltic andesites

E –W, N45, N70

Tectonic regime

Age range (Period or Ma)

Progressive, oscillatory extensional faulting Episodic extensional faulting Episodic extensional faulting

Sed. Anisian; volc. Anisian – Ladinian

E– W extension

Permian and Trias

Transtension, episodic sinistral transpression, flower structures (mid-Trias) Sinistral strike slip, extensional or domal, lithospheric extension and attenuation Extensional to transtensional

Middle Trias (Anisian to Ladinian and Early Carnian)

Late Carboniferous – Early Permian, 305– 299 Middle to Late Permian

Refs

Venturini (1991); Massari et al. (1991); Venturini (1983) Venturini (1991); Massari (1988) Venturini (1990, 1991); Farabegoli et al. (1985); Lucchini et al. (1980) Bosellini (1965, 1973); Winterer & Bosellini (1981); Doglioni (1987) Doglioni (1984a, b)

Late Ladinian, and Early Carnian

Doglioni (1987); Pisa et al. (1979)

c. 280 –260 Early Permian (c. 290, c. 278)

D’Amico et al. (1980); Massari (1986); Bargossi et al. (1998); Visona` et al. (2007); Poli (1997)

45

(Continued)

LITHOSPHERE EXTENSION: MODELLING V. DATA

Southalpine (SA) Eastern Alps (Carnia)

Sediments

46

Table 3. Continued Geographic and tectonic unit location (see Fig. 1)

Sediments

Formation or group

Continental

Tregiovo, Val Gardena Sdst.

Orobic-Bergamask and Brescian Alps

Continental: alluvial and lacustrine, red beds, shelf, basinal

Collio, Verrucano Lombardo,

Lugano Valganna

Continental

Helvetic (HE) Pelvoux-Belledonne, Grandes Rousses, Aiguilles Rouges

E– W, N– S

Volcanics

Structural trend

Calc-alkaline, basaltic andes., andesites, dacites, rhyodacites, rhyolites Calc-alkaline rhyolitic ignimbrites, andesites, rhyodacites

Tectonic regime

Age range (Period or Ma)

Refs

Dextral strike slip, transtensional

c. 285 –274

Schaltegger & Brack (2007) and refs; Bargossi et al. (1998)

Transtensional, pull-apart, semigraben, transpressive, calderic

Early Permian

Vai (1980); Cassinis & Gianotti (1983); Cassinis et al. (1988, 2007); Cassinis & Perotti (1994); Cadel (1986); Cadel et al. (1996); Philippe et al. (1987); Brusca et al. (1981); Gaetani & Jadoul (1979); Casati & Gnaccolini (1967) Buletti (1983); Bakos et al. (1990); Schaltegger & Brack (2007)

Rhyolites, dacites, andesites, basalts, tuffs

Early Permian

Continental, marine

NNE –SSW?

Rhyolites, dacites, andesites

E –W shortening, N–S extension, NE– SW srike slip faulting

Multiple

Carboniferous, Permian

Shallow marine

NNE

Basalts

ESE extension

Extensional, rift half graben

Early Trias

Banzet et al. (1985); Capuzzo & Bussy (2001); Barfe´ty et al. (2001); Strzerzynsky et al. (2005) Gillcrist (1987); Lemoine & Truempy (1987)

A. M. MAROTTA ET AL.

Athesian Volcanic Platform (W Dolomites-Adige valley)

Basin trend (present orientation)

LITHOSPHERE EXTENSION: MODELLING V. DATA

2007). However, the massive nature of this volcanic platform only broadly makes possible inferring links of feeder dykes with the trends of the main faults, which may be guessed by comparison with marginal areas in which the substratum is exposed (e.g. Brescian Alps, Tregiovo). Activation of the Permian continental basins west of the Athesian Volcanic Platform was controlled by a dextral transtensional regime (pull-apart) along the Giudicarie and Insubric tectonic lineaments and by related minor fault populations (Cassinis & Perotti 1994; Cassinis et al. 2007); two volcanic cycles took place during Permian. Basins of the Orobic and Brescian Alps resulted from similar transtensional to transpressive tectonic regimes that generated asymmetric graben structures, with local calderic acceptors of both volcanics and sediments. In the external Helvetic –Dauphinois domain (Pelvoux–Belledonne) of the Western Alps, Gillcrist et al. (1987) attribute a Triassic age to eastdipping listric faults for the thickening of Triassic deposits in the half-graben structures and, together with Lemoine & Truempy (1987), note the occurrence of basalts in the syn-rift sedimentary sequence. An older signal of a Late Carboniferous–Early Permian extensional regime is reported in the same areas and, more to the north, in the Aiguilles Rouges. In summary (Table 3), along both margins of the present-day Alpine belt volcanism is concordantly seen to appear, weakly or untouched by the Alpine orogeny, after the Variscan cycle at the Late Carboniferous to Permian transition, taking place in Middle to Late Triassic times. In the whole Southalpine domain, a marked volcanic activity reappears in Lower Permian continental deposits. In the Dolomites and Carnia Early–Middle Triassic sediments host dykes and pillowed basaltic lavas. The sedimentary sequences record periodic fluctuations both of Early Permian and at the Permian– Triassic boundary in the continental –marine transition. Continental Permian basins are affected by episodic localized deepening. A regime of dextral strike-slip faulting with a variable extensional component is compatible with Permian tectonic history, although minor sinistral effects are recorded in the Dolomites. Extension begins to prevail on strike-slip displacement since Triassic times.

Numerical analysis Model set up To understand the geodynamic settings at the Permian–Triassic high thermal regime, we use finite-element techniques to model the active lithospheric extension process overprinting the deformation patterns driven by a previous phase of oceanic subduction, followed by continental

47

collision and purely gravitational evolution. The continuity, momentum and energy equations r ~ v¼0 @ tij @p g ¼  r~ @xj @xi   @T þ~ v  rT ¼ r  (KrT) þ rH rc @t are integrated within a rectangular domain, 1400 km wide and 700 km deep (Fig. 5a) in which the flow is driven by velocity boundary conditions and density contrasts: ~ v is the velocity, r the density, p the pressure,~ g the gravity acceleration, tij the deviatoric stress tensor, c the thermal capacity at constant pressure, T the temperature, K the thermal conductivity and H the heat production rate per mass unit. Model resolution is 5 km in the horizontal scale and varying from 5 to 30 km in the vertical scale, from the crust downwards. The 2D finite-element code SubMar (Marotta et al. 2006) is used for the analysis. Active extension is simulated by applying a constant velocity Ue from 0 to 230 km of the right boundary of the model (Fig. 5a) corresponding to the crustal thickness at the beginning of the simulation, with zero normal stress being assumed from 230 to 280 km; at the rest of the right boundary the velocity is fixed to zero. To test the influence of the wideness of the region where extension velocity is fixed on the thermal regime and crustal thickness variations, we also used models in which Ue is constant from 0 to 280 km; since these latter models show negligible differences with respect to the former, only models of the first type will be discussed. Compositional and physical parameters are listed in Table 4. More details on the model set-up are given in Marotta et al. (2006). Four model types were developed, differing for different values of the assumed forcing extension velocity, 0.0, 0.5, 1.0 and 2.0 cm a21, simulating extension of far-field origin due, for example, to active subduction. The following common features characterize all the models. (1) The crust is compositionally differentiated from the mantle via the Lagrangian particle technique (e.g. Christensen 1992), as in Marotta & Spalla (2007) and Spalla & Marotta (2007). At each time step, the non-dimensional function C describing the elemental composition is calculated on the basis of the density of the oceanic or continental crust particles in the element, such that C ¼ 12 C o 2 C c, with C o and C c density of the oceanic or continental crust particles in the element. C, C o and C c vary from 0 for pure mantle to 1 for pure crust (either oceanic or continental).

48

A. M. MAROTTA ET AL.

Fig. 5. Model set-up and boundary conditions. (a) Sketch illustrating the compositional layering and the velocity boundary conditions. (b) Starting configuration of the system at the beginning of the numerical simulations representing the mechanical and thermal evolution reached at 300 Ma after a period of active oceanic subduction (lasting 52.5 Ma) and subsequent gravitational evolution (lasting 72 Ma); model grav2 in Marotta & Spalla (2007).

(2) An incompressible viscous fluid is assumed with temperature and composition-dominated viscosity and density and calculated for each elements as follows:

m ¼ mm [1  C o  Cc ] þ mo Co þ mc C c with

mi ¼ mi0 (C) exp

 i   Ea (C) 1 1  nR T T0

and

r ¼ ro ½1  aðT  To Þ þ Do rCo þ Dc rCc where m is the elemental viscosity, index i stands for mantle (i ¼ m), continental crust (i ¼ c) and oceanic crust (i ¼ o), mi0 is the reference viscosity at the reference temperature T0, Eia the activation energy, T the temperature, a the thermal expansion

LITHOSPHERE EXTENSION: MODELLING V. DATA

49

Table 4. Material properties used in the numerical modelling Continental crust Rock components

66% gneiss þ 33% granite

Mean densitya (kg m23) Thermal conductivityb (W mK21) Heat generationb (1026 W m23) Rheology Activation energy (kJ mol21) A (Pa2n s21) N

2640 3.03 2.5 Dry granitec 123 7.92  10229 3.2

Oceanic crust

Mantle

7% basalt þ 16% dolerite þ 77% gabbro 2961 2.1 0.4 Dry diabased 260 8.04  10225 3.4

100% dry dunite 3200 4.15 0.002 Dry dunitee 444 6.31  10217 3.41

References: a, Dubois & Diament (1997) and Best & Christiansen (2001); b, Rybach (1988); c, Ranalli & Murphy (1987); d, Kirby (1983); e, Chopra & Peterson (1981).

factor and n i the exponent of the power law creep equation. We assume a viscous behaviour for the whole system and do not account for elasticity and brittle behaviour at the shallow levels of the lithosphere. Note that elasticity becomes crucial when the bending of the top of the slab is analysed so as to optimize the fit between the natural trench profiles and when the bathymetry predicted by the geodynamic models, the interaction of subducted slabs with the 660 km discontinuities (Funiciello et al. 2003) and the instabilities driving the initiation of subduction (e.g. Hall et al. 2003) are investigated. All these processes are beyond the remit of the present study. Rather, the assumption of a viscous lithosphere proves to be appropriate in reproducing the average continuum properties of a discontinuous medium and makes analyses of gross mechanical behaviour during different geodynamic processes possible (e.g. Marotta et al. 2006 and references therein). Since only a viscous rheology is considered, the crust is expected to remain strong at much higher stresses than the brittle limit and we introduce a viscosity cut-off value to minimize this effect. Although we checked different values, we here discuss only models accounting for a viscosity cutoff of 1025 Pa s since it guaranteed the best agreement between natural data and model prediction during the previous subduction and collisional phase (Marotta & Spalla 2007). (3) Complexities, such as the phase transition at 410 km or phase transitions in subducted materials are not taken into account. Our analysis does not include viscous dissipation either. Previous works (e.g. Burg & Gerya 2005) indicate that viscous heating may be important for high deformation rates and that, at crustal levels, it decreases with time. Since our analysis is controlled by low values of deformation rate, we do not expect significant changes in our major conclusions.

(4) Thermal boundary conditions correspond to fixed temperatures at the top and the bottom of the model domain, at 300 K and 1600 K, respectively, and to zero thermal flux imposed at the vertical sidewalls. In the following discussion, the models will be cited as Ue0.0, Ue0.5, Ue1.0 and Ue2.0, according to the respective assumed velocity for the forced extension. While the base of the crust is defined compositionally, the base of the lithosphere is defined thermally by the isotherm 1500 K. The initial configuration of the four models corresponds to the shallow and deep heterogeneities generated at 300 Ma by an oceanic subduction, which consumed a 2500 km wide ocean during a 52.5 Ma convergence from 425 to 372.5 Ma (e.g. Tait et al. 1997; von Raumer et al. 2003), followed by continental collision and a purely gravitational evolution for 72.5 Ma, thereby reproducing the Variscan subduction and collisional phases affecting the preAlpine continental crust (Marotta & Spalla 2007; Spalla & Marotta 2007). In order to account for this previous history, we use as initial distributions of temperature and crustal markers those predicted by model grav2 in Marotta & Spalla (2007) and represented in Figure 5b. The white vertical line in Figure 5a indicates the contact between the two continents at the beginning of the simulated forced extensional phase and is located approximately at 270 km in the horizontal scale of Figure 5b. Since model grav 2 predicted that the largest amount of subducted crustal material rises below the overriding plate, making it the favoured location for Permian–Triassic extension and metamorphism, we implemented our active extension models by limiting the forcing to the overriding plate and focused the discussion there on model predictions and the comparison with natural data. For further details about model setup and assumptions, the reader is referred to Marotta & Spalla (2007).

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Thermo-mechanical evolution Figure 6 shows the velocity and temperature fields predicted after 5.5 Ma (panels a, d, g and l), 40.5 Ma (panels b, e, h and m) and 70.5 Ma (panels c, f, i and n) of forced extension at a rate of 0.0 cm a21: purely gravitational evolution (panels a, b and c), 0.5 cm a21 (panels g, h and i), 1.0 cm a21 (panels d, e and f ) and 2.0 cm a21 (panels l, m and n). Model Ue0.0 is exclusively controlled by gravitational forces and, hence, the thermal and compositional re-equilibration of the system is very slow and only at the very late stages of the evolution are significant variations in the thermal field recognized, mainly in the deep portion of the system (Fig. 6c). At crustal levels the system remains rather stable, from both a thermal and a physical point of view (Fig. 6b, c). When the extension is forced, the thermal re-equilibration of the system accelerates at a rate proportional to the rate of extension. From the base of the crust of the overriding plate, a progressive heating occurs at deep levels since the beginning of the forced extension and an almost homogeneous 80-km thick lithosphere is predicted after 70.5 Ma of active extension at a rate of 2.0 cm a21 of fixed extension, which is comparable to an initial thickness for a young continental lithosphere (Fig. 6l). Striking changes occur in the markers’ distribution in the area where forced extension is concentrated. Subducted crustal markers belonging to the subducted slab and coming from the bottom of the model domain reach shallow depths, thereby inducing crustal material underplating (Fig. 6f, i & l). While no significant variation in continental crust thickness is predicted by model Ue0.0 (Fig. 6a, b & c), models Ue0.5, Ue1.0 and Ue2.0 predict a progressive thinning of the crust (Fig. 7a, b & c). However, the physical thinning is not accompanied by an equivalent thermal thinning of the crustmantle system, thus leading to a progressive cooling of the crust –mantle interface. This effect may be ascribed either to the boundary conditions used to simulate the forced extension – the location of extensional velocity boundary conditions may be too close to the deformation area – or to the slow forced extension that makes the lithosphere cool faster than it is being mechanically thinned. Figures 8 and 9 show the variation in time of the strain regime (panels i1) and maximum strain rate (panels i2) at lithosphere (Fig. 8) and crustal (Fig. 9) levels for the purely gravitational model Ue0.0. Segments indicate the deviatoric strain rate eigenvectors closer to the horizontal, with length proportional to the magnitude. Light grey is indicative of extension; dark grey of compression. Areas where the maximum values of the maximum strain rate localize identify locations of

potential compressional or extensional failure (brittle or ductile faulting; Figs 8 & 9, panels i2). At lithospheric level, four major areas of alternate faulting localize in the regions where the major lateral heterogeneities developed at depth during the previous subduction–collision cycle. Deformation is compressive between 250 and 300 km and between 350 and 400 km and extensional between 150 and 200 km and between 500 and 600 km. Extensional faulting localizes above the two areas of upwelling of previously subducted oceanic crustal material (compare with Fig. 6a, b & c). The magnitude of maximum shear strain in these areas decreases with both distance from the ancient suture zone (x ¼ 0) and with time as a consequence of the gravitational relaxation of the system. Deep deformation propagates upward, through the crust (Fig. 9). However, while the lower crust undergoes the same type of deformation as the lithospheric mantle at the very early stages of the examined evolution, the upper crust is under compression throughout its horizontal extension (Fig. 9, panels a, b, c, d & e) and only at later stages (at least after 70 Ma from the beginning of the simulation) does normal faulting propagate at shallow depths (Fig. 9g, h & i). It may be noted that the system underwent a previous convergence phase that induced a shallow compressive stress regime that could delay the propagation of normal faulting from mantle to crustal levels. The decoupling between deep and shallow crustal deformation field recognizable in Figure 9 is better illustrated in Figure 10, where at different ages crustal regime is plotted at surface (continuum line) or averaged through different thicknesses: the whole crust (grey area), the upper crust (dashed line) and the lower crust (dotted line). During the first 70 Ma two major deformed areas are distinguishable in the lower crust, one under compression between 250 km and 500 km from the suture zone, and the other under extension from 500 km to 600 km. The magnitude of the deformation rates in these regions decreases with time, keeping the same style and wavelength throughout the considered time span. The lower crust shows a very different deformation pattern. The average upper crustal regime and surface regime are almost coincident through the simulation time span, showing very high frequency changing in magnitude, but always compressive from 300 Ma until about 230 Ma. Only in the very late stages and only far away from the ancient suture zone does deep extension succeed to propagate to shallow depths (Fig. 10f, g & h). The deformation style at lithospheric and crustal levels changes drastically when a forced extension is assumed (Figs 11 & 12). For all the values of forced extensive velocity, extension is widespread

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Fig. 6. Large-scale (black arrows) and temperature (grey colour and dashed thin lines) fields predicted after 5.5 Ma (panels a, d, g and j); 40.5 Ma (panels b, e, h and k) and 70.5 Ma (panels c, f, i and l) of forced extension at a rate of 0.0 cm a21 (purely gravitational evolution; panels a, b and c); 0.5 cm a21 (panels g, h and i); 1.0 cm a21 (panels d, e and f) and 2.0 cm a21 (panels j, k and l). Black points represent oceanic- and continental-type crust.

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Fig. 7. Mechanical thinning of crust thickness in time predicted by models Ue0.5 (panel a), Ue1.0 (panel b) and Ue2.0 (panel c). Both absolute age (with respect to present day) and relative age of simulation are indicated.

through the lithosphere mantle (Fig. 11) and compression localizes only at depth in the lithosphere in relation to the asthenospheric circulation. The width of these compressive areas increases with time and with the increase of the extensive forcing since the progressive thinning of the crust allows upwelling of deep mantle circulation. Unlike the lithospheric mantle, the crustal regime is very heterogeneous. In model Ue0.5 (Fig. 12) nucleation of normal faulting zones appears periodically at shallow crustal depths, with a periodicity of about 10 Ma, at the beginning of forcing phase, then decreasing to about 5 Ma as evolution progresses. It is worth noting that localization of the intense shear zone migrates toward the limit of the study region, decreasing progressively in intensity; small pulsations persist in the late phase of the system evolution. When forcing velocity increases to 1.0 cm a21, the magnitude of maximum shear strain rate amplifies by a factor of 2 (Fig. 13). Furthermore, nucleation of maximum shear zones occurs with higher frequency, the periodicity being about 5 Ma. Nucleation and relaxation cycles persist until about 230 Ma, when a relatively stable phase starts with a deformation rate that is invariant in the vertical direction and increases towards the limit of the study area. The beginning of this stable deformation phase seems to be related to the magnitude

of the forcing velocity. In fact, model Ue2.0 reaches this stage at about 260 Ma (Fig. 14), earlier than in model Ue1.0, although periodicity of nucleation of maximum shear zones is the same, about 5 Ma. A peculiarity of high velocity active rifting models Ue1.0 and Ue2.0 is the occurrence of a horizontal low strain rate zone between 210 km and 220 km. The shape of these zones is similar for both models. If we focus on the temperature field and on the variation of crustal thickness for both models, we note that the low strain rate regions appear when crustal thickness decreases below 10 km and the mantle reaches low depths and temperature between 600 K and 800 K, thus driving local mantle hardening. The vertical extensional gradient between the crust and the base of the lithosphere may be responsible for the delay between crustal thinning and thermal erosion of the lithosphere.

Discussion of modelling predictions in the framework of natural data It seems realistic to compare the natural tectonic signals of the upper and lower continental crust with the extensional components of the deformation predicted in the lithosphere by numerical simulation and the natural PT values, which are listed in

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Fig. 8. Variation in time of strain regime (panels a1 – f1) and maximum strain rate (panels a2 –f2) predicted by the purely gravitational model Ue0.0 at lithospheric level. Segments indicate the deviatoric strain rate eigenvectors closer to the horizontal (colour: white for compression and black for extension). Length is proportional to strain rate magnitude. Black continuum lines indicate isotherms.

Table 1 and recorded in different tectonic units of all first-order structural domains, with the predicted PT conditions of each crustal marker. Since the natural PT estimates were obtained on continental and oceanic rocks, the following main conditions are checked for a complete fitting between model predictions and natural data: (1) corresponding ages; (2) coincident continental or oceanic lithological affinity; and (3) coincident PT values. In order to visualize the fitting of natural PT data with the predicted PT conditions of the markers, a colour is attributed to the markers that satisfy the fit, a different one for each natural datum being listed in Table 1. This graphic representation makes it possible to follow throughout the simulated time span the localization of successive rocks positions belonging to the present-day structural domains (Figs 15, 16, 17 & 18).

Metamorphism Model Ue0.0. The fit between PT natural data and model predictions occurs only for rocks from the Austroalpine domain (Fig. 15), namely data 9, 11a, 19a, 19c of Table 1 and Figure 2, respectively corresponding to the Arolla Series of Dent Blanche nappe, to the Languard-Campo nappe granulites and Saualpe– Koralpe metapelites and pegmatites. Rocks recording these Tmax values are metaintrusives and metapelites belonging to tectonic units in which some Permian gabbros are emplaced (Mont Collon-Dents de Bertol in the Dent Blanche nappe, Sondalo in Languard – Campo nappe and Baerofen and Gressenberg in Saualpe –Koralpe eclogitic complex). The data fit occurs all along the upper plate, with the exception of 11a localized above the mantle plume at 300 km

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Fig. 9. Variation in time of strain regime (panels a1 –i1) and maximum strain rate (panels a2 –i2) predicted by the purely gravitational model Ue0.0 at crustal level. Segments indicate the deviatoric strain rate eigenvectors closer to the horizontal (colour: white for compression and black for extension). Length is proportional to strain rate magnitude. Black continuum lines indicate isotherms.

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Fig. 10. Variation in time of crustal regime at surface (solid black line) and averaged through the whole crust (grey area); the upper crust (dashed black line) and the lower crust (dotted black line).

accompanying the rise of the previously subducted oceanic and continental crust. Duration of the fit is controlled by estimated ages in all the tectonic units (Fig. 19). Note that the P/T ratio characterizing these data is 1.1 . 1023 GPa . K21. Model Ue0.5. Here the fit occurs with PT conditions of additional units from both Penninic and Austroalpine domains (Fig. 16). In the Penninic domain the fit is realised by Dora-Maira metapelites (datum 3 in Table 1 and Fig. 2); it is localized above the mantle plume at 300 km and persists from the beginning until c. 275 Ma in spite of the larger time interval suggested by the inferred geological age (Fig. 19). In the Austroalpine domain, units fitting PT data correspond to Sesia Lanzo zone basic and acidic granulites, Mont Mary nappe metapelites, Dent Blanche nappe metaintrusives from Arolla Series, Languard-Campo granulites, Matsch nappe metapelites, Uttenheim Ahrntal metapelites, Woelz complex metapelites, Saualpe –Koralpe metapelites and pegmatites (respectively 4a, 4b, 7, 9, 11a, 13, 14, 16a, 16b, 19a, 19c in Table 1 and Fig. 2). While Austroalpine data fit for units 9, 14, 19a and 19c occurs all along the upper plate, the rest of the data concentrate the fit above the mantle plume at 300 km. Among the new data are those from Sesia Lanzo granulites that crop out in the surroundings of the Corio and Monastero gabbros. The P/T ratio characterizing data from both Penninic and Austroalpine domains is

0.6 . 1023 GPa . K21. The fit in both tectonic domains is generally decreasing in time and is only incipient for some of the data (4a, 4b and 14) (Fig. 19). Model Ue1.0. The fit occurs for the same units as in model Ue0.5 (Fig. 17) with a general increase in the amount of fitting markers and, for some of them (4a, 4b, 7 and 16b), with a longer duration (Fig. 19). In the Austroalpine domain, a new datum (6 in Table 1 and Fig. 2, corresponding to Mt Emilius metabasics) matches with markers mainly during the last period of its estimated age. By contrast, the incipient fit of datum 14 in model Ue0.5 disappears here. Data-fit areas are more heterogeneously distributed along the upper plate with respect to the previous models, and at c. 250 Ma fitting also initiates with underplated crustal markers. Model Ue2.0. In this higher velocity extensional model, marker fitting with data from the Southalpine domain (24a and 24b of Table 1 and Fig. 2, corresponding to Ivrea zone metapelites and metabasics, respectively) is one of the main differences with respect to the previous fitting pattern. These PT estimates were performed on metapelites and metabasics from the Ivrea zone, which comprises the gabbro intrusions constituting the Upper and Lower Mafic Complex and the Finero mafic body (Table 2 and Fig. 4), both associated with sub-continental ultramafics. Permian Tmax recorded in Brianc¸onnais metapelites of the Ruitor basement

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Fig. 11. Variation in time of strain regime (panels a1 –f1) and maximum strain rate (panels a2 –f2) predicted by models Ue0.5 (panels a1 and b1), Ue1.0 (panels c1 and d1) and Ue2.0 (panels e1 and f1) at lithospheric level. Segments indicate the deviatoric strain rate eigenvectors closer to the horizontal (colour: white for compression and black for extension). Length is proportional to strain rate magnitude. Solid black lines indicate isotherms.

(datum 1 of Table 1 and Fig. 2) are finally reproduced during model simulation, with a maximum at c. 290 Ma. In addition, as we are dealing with Austroalpine domain, the PT values recorded in central Sesia Lanzo zone granulites, Valpelline Series granulites of Dent Blanche nappe, metapelites and metabasics from southern Languard Campo nappe, Uttenheim Ahrntal metapelites, Strieden Kreuzechgruppe metapelites, Sopron metapelites, Saualpe –Koralpe metapelites (data 5, 8, 14, 15, 17 and 19b of Table 1 and Fig. 2, respectively) are reproduced by model predictions. In this case, too, while zones of data fit are heterogeneously distributed along the upper plate, the fitting with crustal underplated markers becomes massive and starts between 280 and 270 Ma. At the same time, the fit on the thinned continental crust belonging

to the upper plate is noticeably reduced, up to the end at c. 250 Ma. The occurrence of this misfit is consequent to the high thinning of the continental crust of the upper plate, which is responsible for the upward migration of crustal markers to a depth 5 km, unaccompanied by an equivalent thermal erosion of the system. The differences in correspondence between model predictions and natural PT-age data for the four models are well synthesized in Figure 19, where the different weights of natural rock age estimates are shown using light grey for radiometric ages and dark grey for geologically determined ages (Table 1). Where the age is geologically determined, model fit may suggest a reliable constraint to the time interval for the corresponding metamorphic imprints. With respect to model Ue0.0, simulating a

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Fig. 12. Variation in time of maximum strain rate predicted by model Ue0.5 at crustal level. Solid black lines indicate isotherms. Thick black dashed lines indicate the base of the continental crust.

purely gravitational evolution, forced extension models improve the fit proportionally to the extensional velocity increase and, for model Ue2.0 only, the fit is across all structural domains except for

the andalusite-bearing Silvretta metapelites (datum 10 of Table 1). The only exception occurs for data 19a and 19c, for which the fit along the upper plate decreases in time until the end at c. 285 Ma.

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Fig. 14. Variation in time of maximum strain rate predicted by model Ue2.0 at crustal level. Solid black lines indicate isotherms. Thick black dashed lines indicate the base of the continental crust.

Fig. 13. Variation in time of maximum strain rate predicted by model Ue1.0 at crustal level. Solid black lines indicate isotherms. Thick black dashed lines indicate the base of the continental crust.

After a gap of c. 20 Ma, the data fit again, but now with underplated crustal markers. This fit discontinuity also characterizes other data and is always consequent to the end of fit on the upper-plate continental crust and to the successive fit on the underplated crust. The maximum shear strain rate pattern predicted by the purely gravitational model (Ue0.0, Figs 9 & 10) shows significant strain localization, mainly

under a compressional regime, only at the crust – mantle interface. Only active extensional models (Ue0.5, Ue1.0 and Ue2.0, Figs 12, 13 & 14) show strain localization throughout the crust, making the development of shear zones both in lower and upper crust possible: this prediction is in rather good agreement with the occurrence of discrete shear zones and mylonitic textures at granulite to amphibolite facies conditions (e.g. Gosso et al. 1997; Lardeaux & Spalla 1991).

Magmatism As noted above, the most peculiar character of the Permian– Triassic magmatic activity is the diffuse emplacement at the base of the crust of gabbro

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Fig. 15. Comparative analysis between PT natural data and modelling prediction for model Ue0.0. Black points represent oceanic- and continental-type crust. Different colours are attributed to the markers satisfying the fit; a different one for each natural datum listed in Table 1. For greater clarity; key numbers of fitted natural data are also indicated in each panel. Thin black dashed lines indicate isotherms.

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Fig. 16. Comparative analysis of PT natural data and modelling prediction for model Ue0.5. Meaning of black and coloured points and black dashed lines as in Figure 15.

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Fig. 17. Comparative analysis of PT natural data and modelling prediction for model Ue1.0. Meaning of black and coloured points and black dashed lines as in Figure 15.

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Fig. 18. Comparative analysis of PT natural data and modelling prediction for model Ue2.0. Meaning of black and coloured points and black dashed lines as in Figure 15.

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Fig. 19. Fit duration for natural data listed in Table 1. (a) model Ue0.0; (b) model Ue0.5; (c) model Ue1.0; (d) model Ue2.0. Light grey and dark grey colours are used to differentiate the radiometric age estimates from the geologically determined ages of the natural data; respectively (Table 1).

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stocks and the occurrence of basaltic products in the volcanics. To satisfy this kind of magmatic activity, PT conditions predicted for lithospheric and astenospheric mantle by different models must allow transgression of the peridotite solidus (e.g. Watson et al. 1990). Predictions from all considered models satisfy the thermal state necessary for partial mantle melting, although to different extent and at different depths. In model Ue0.0, PT conditions favourable for producing mafic melts are reached in the subcontinental mantle at 75 km depth but localized above the two thermal highs between 200 and 400 km and between 500 and 600 km, where previously subducted crust upwells at shallow depths (Figs 5b, 6 & 15). With the increase in the assumed extensional velocity, PT conditions favourable to partial mantle melting are reached at shallower depths and affect a wider region, which is expanded all along the upper plate in model Ue2.0 at 50 km depth. Basaltic melt production is thus compatible with all the simulated tectonic settings, although the huge amount of acidic magmas associated with gabbro emplacement and basaltic volcanic products should also be taken into account in trying to discriminate among them. For this purpose we may recall that the bimodal character of the Permian igneous activity was justified with crustal contamination of basaltic magmas generated within enriched lithospheric and/or asthenospheric mantle sources in lithospheric extension (e.g. Cortesogno et al. 1998; Rottura et al. 1998). To allow the partial melting of continental crust, the thermal state must be similar to that suggested by the lower P/T ratios characterizing metamorphic imprints recorded in the pre-Alpine continental crust (Table 1 and Fig. 2). Only the high extensional models Ue1.0 and Ue2.0 satisfy all the above conditions.

Tectonic history from sedimentary and volcanic evolution Examining the natural time-space periodicity of faults, sedimentary basins and volcanic trends and the modelled sequential imaging of strain states at various levels of the lithosphere provides an insight into the puzzling question of the modes of migration of tectonic signals from mantle to surface (Figs 8–14). As noted at the end of section (a) above, the purely gravitational model does not allow the development of main fault seams, driving both basin opening and magma conduits (Figs 8, 9 & 10). On the other hand, the main tectonic regime persisting in the upper and intermediate crust is compressive up to 230 Ma, when the extensional focus located in the mantle between 500 and 600 km from the ancient suture zone is

transferred up to the surface, although at very low magnitude. Consequently, only a negligible crustal thinning is accommodated in the continental lithosphere. This is in contrast to the sedimentary and volcanic records indicating successive tectonic pulses accounting for phases of basin opening and short-lived volcanism. In active extensional models Ue0.5, Ue1.0 and Ue2.0 (Figs 12, 13 & 14), periodic strain localizations are predicted throughout the crust up to the surface. Periodicity of these strain localizations ranges between 10 and 5 Ma, comparable as order of magnitude with volcanic pulses, episodic basin deepening, periodic facies fluctuations and related episodic faulting. Fault propagation between upper and lower crust makes identification of magma ascent pathways possible from deep-seated mafic intrusions and partially molten sites of the crust.

Conclusive remarks Our analysis demonstrates that the integrated use of geological data (metamorphic PT estimates, magmatic and sedimentary histories) and predictions of numerical modelling, which take into account the thermal instabilities induced by mechanic perturbations, sheds light on the tectonic evolution of a fossil passive margin dispersed in separate fragments within a subduction –collisional belt, as is the case for the Permian–Triassic thinned continental lithosphere recycled in the Alpine orogen. On a large scale, uprising of previously subducted crust occurs mainly underneath the upper continental plate, reaching shallower depth for the higher extensional model, and underplating takes place over a wider area below the thinned crust. Crustal thinning is proportional to the assumed extensional velocity and ranges from almost 0, for a purely gravitational evolution, to c. 80%, when the maximum rate of extension is 2.0 cm a21. In the purely gravitational model, decoupling of the deformation regime occurs between the upper crust and the lower crust/mantle coherent system, whereas in the active extensional models the extensional regime is widespread through the crust, allowing strain rate localization that in the natural systems may correspond to faulting that generates basin opening and provides magma ascent pathways from deep seated mafic and acidic intrusions. Strain localization shows a periodicity comparable as per order of magnitude with that of episodic faulting and related volcanic pulses, periodic basin deepening and facies fluctuations. In the active extensional models, a low extension velocity of 0.5 cm a21 reduces the minimal P/T ratio to a half with respect to that of the purely gravitational model. However, crustal thinning is not accompanied by a concurrent thermal erosion and,

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hence, the thermal state of the crust remains rather low to satisfy completely the PT values frozen in Permian–Triassic mineral assemblages. Only at a rate of forced extension of 2.0 cm a21 does the data set from the Southalpine domain begin to fit and the maximal number of tectonic units from all structural domains are matched by the predicted thermal regime. Only this last model generates thermal conditions suited to a partial melting of the crust accompanying mafic magma emplacement and HT– LP metamorphism. Our model is physically compatible with the natural case in which the distribution of gabbro complexes in the Austroalpine and Southalpine crusts suggests that the rifting was asymmetrical and that the continental crust belonging to these two domains acted as a hanging wall during an asymmetric lithospheric thinning that paved the way to the opening of the western Tethys ocean. This interpretation is supported in the natural Alpine case by the gradual age transition between the Permian–Triassic magmatic and metamorphic histories that ends with ocean crust formation; this picture is reinforced by a comparison with the new Permian ages obtained from gabbros and tonalites underplating the thinned continental crust of the present-day passive Galicia margin. The authors thank the editor Uwe Ring and the two reviewers, Claire Currie and Klaus Gessner, for their constructive criticisms and helpful suggestions. Cesare Perotti is thanked for useful discussions. M.I.S and G.G. were funded by FIRST07 and CNR-IDPA. A.M.M. was funded by FIRST07.

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Reconciling short- and long-term measures of extension in continental back arcs: heat flux, crustal structure and rotations within central North Island, New Zealand T. A. STERN Institute of Geophysics, Victoria University of Wellington, P.O. Box 600, Wellington, New Zealand (e-mail: [email protected]) Abstract: Geodetic, palaeomagnetic and andesite-age data all point to a fan-like opening of central North Island, New Zealand, in the past 5 Ma. Palaeomagnetic rotation rates for the eastern North Island are c. 68 Ma21 and this is accompanied by back-arc extension of up to 19 mm a21 at the Bay of Plenty coast and lesser values further south. Although the geodetic observations only span a decade, they show a remarkable consistency with the palaeomagnetic and volcanic arc migration data that span 5 Ma. A fan-like pattern of extension in central North Island is implied, similar to that seen in other continental back-arcs. When rapid fan-like openings do occur, they are likely to be accompanied by extreme thermal events. Heat flux from the central North Island occurs at one of the highest continental rates recorded: c. 26 MW per km of strike of volcanic zone, or 4.3 GW in total. Prior to the Pliocene extension and rotation, central North Island had a c. 20 Ma history of compression and overthrusting. It is proposed that thickening, then subsequent detachment of mantle lithosphere during this phase had a role to play in the Pliocene back-arc opening and the consequent extreme heat flux.

Continental rifts zones are highly varied in their kinematics and structure. Particularly enigmatic are continental back-arc rifts. Some are relatively benign with little evidence of extension while others show evidence of high extensions rates, high heat-flow and volcanism. Two observations underscore this variability (Fig. 1a): Firstly some continental back-arc basins appear to be hot regardless of whether extension is low or high (Hyndman et al. 2005). Secondly, extension and/or rotation in back-arc regions is shown from palaeomagnetic evidence in some continental back-arcs to have occurred rapidly for relatively short bursts of geological time (Otofuji & Matsuda 1987) (Fig. 1b). In general, continental back-arcs are characterized by high elevations for their crustal thickness, high heat flow, thin mantle lithosphere, in some cases active extension, and evidence for older fold and thrust belts that predate the back-arc extension (Fig. 1a) (Hyndman et al. 2005). In this study, we focus on a continental back-arc system that is smaller, younger and kinematically more simple than the western Cordillera of the USA (Humpherys 1995). This system is within the central North Island of New Zealand and is considered here to include not only the areas of Quaternary volcanism, but also the Miocene North Wanganui and Taranaki basins further to the west (Fig. 2). A previous study has compared geological and geodetic observations within the central North Island (Nicol & Wallace 2007) and finds a broad

compatibility of GPS and geological data that span 1.5 Ma. We extend this comparison by examining estimates for five key data sets that span up to 5 Ma: the integrated heat flux; geodetically and GPS-determined strain and extension rates; palaeomagnetic rotations; volcanic migration; and up-todate crust and upper mantle images. The intent of this paper is to explain all five sets of observations within the context of a simple kinematic and dynamic model. In particular, we address the following questions: what condition leads to fan-like opening of continental back-arcs, why are these rotations rapid for short periods, and why is the heat output in the central North Island so high compared to regular continental rifts and other continental back-arcs?

Tectonic setting Continental New Zealand has evolved and deformed along the boundary zone between the Australian and Pacific plates (Fig. 2). East of the North Island, oblique subduction of the Pacific plate takes place and this obliquity increases southward to the northern South Island. Westward of the Hikurangi Margin (Fig. 2), subduction occurs at a low angle (158) until the East Coast of the North Island where the dip of the Pacific plate steepens to around 508 (Reyners 1980; Henrys et al. 2006). Fore-arc marine sediments ranging in age from Miocene to Pliocene make up the East Coast

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 73– 87. DOI: 10.1144/SP321.4 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. (a) Conceptual model of a continental back arc system based on the southern Cordillera back arc basin (western North America) showing elevated heat flow, small scale convection in the mantle, and a highly thinned mantle lid (Currie & Hyndman 2006a). Further landward from the back-arc basin remains of an older fold and thrust belt may be found. Scale varies from 1000 km across for the western North American case to 100 km for the central North Island discussed here. (b) Reconstruction of the Japan Sea for the Miocene showing the rapid rotation for SW Japan based on palaeomagnetic data and K –Ar dating (Otofuji & Matsuda 1987). Note the high rates of rotation (c. 208 Ma21) and of extension (21 cm a21), both many times higher than the rates in central North Island, New Zealand. M, Moho; dark shade, mantle lithosphere; light shade, crust.

Basin (Fig. 2) (Ballance 1993). Palaeomagnetic measurements within these sediments record a range of late Cenozoic rotations and deformation that are key to an understanding of the style of back arc processes to be discussed shortly (Lamb 1988; Mumme et al. 1989; Rowan et al. 2005). A narrow spine of greywacke-cored mountains represents the Axial Ranges of the North Island (Fig. 2), and further west again a linear and regular spaced line of active or recently active andesite/ dacite volcanoes defines an active volcanic front (Fig. 2) (Cole et al. 1995). West of the active volcanic front is the Taupo Volcanic Zone (TVZ): an elongate region about 40 km wide containing most of the active geothermal fields and mainly rhyolitic rocks younger than 500,000 years (Wilson et al. 1995). The TVZ forms just the eastern portion of the wedge-shaped Central Volcanic Region (CVR), which is largely defined geophysically as a region of subsided basement that is isostatically compensated by thinned crust

(Stratford & Stern 2006). Based on the level of the residual gravity field, volcanic migration patterns and seismic crustal structure, it is argued that the CVR had opened by some form of asymmetric fan-like process (Davey et al. 1995; Stern 1987).

Volcanic arc migration Patterns of volcanic migration in the central North Island were initially difficult to decipher (Macpherson 1946; Kear 1964). Two factors in conjunction with the plate tectonic paradigm helped to clarify the situation: K –Ar dates (Fig. 3) (Stipp & Thompson 1971), and the recognition that low and high-potash andesites have different origins (Dickinson & Hatherton 1967). In New Zealand and elsewhere it was noted that lowpotash andesites (,1.5% K2O at 60% SiO2) erupt above the c. 100 km isobath of the subducted Pacific plate (Hatherton & Dickinson 1969;

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Fig. 2. Location and generalized geological structure of North Island, New Zealand. Inset shows how the Havre Trough (the oceanic back-arc basin associated with the Tonga–Kermadec system) has a natural extension into the wedge-shaped Central Volcanic Region (CVR) of continental New Zealand. The Taupo Volcanic Zone (TVZ) is the eastern active portion of the CVR. The active volcanic front is the line of active andesitic/dacitic, low-potassium (K) volcanoes running from the offshore White Island to Mt Ruapehu in the centre of the North Island. Mt Taranaki, 100 km due west of Mt Ruapehu, is a recently active basaltic-andesite with a high-K affinity (Price et al. 1999). NIDFB, North Island Dextral Fault Belt (Beanland & Haines 1998). Grey diagonal pattern, outcrop of greywacke basement rock.

England et al. 2004). Higher potash andesites (2–2.5% K2O at 60% SiO2) and basalts as found in western North Island (Fig. 3), appear to be related to the approximate 200 km isobath of the

subduction zone (Hatherton 1969). More contemporary explanations for high-potash volcanism do, however, exist (McKenzie 1989; Farmer et al. 2002; Stern et al. 2006).

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Fig. 3. (a) K –Ar ages of low-K andesites in central North Island. Sources of data are listed in Stern et al. (2006). Superimposed rotation and translation scheme (Calhaem 1973a; Stern et al. 2006) is shown and the curved arrow above ‘East Coast Basin’ represents the location of palaeomagnetic data giving rotations (with respect to fixed Australian plate) of c. 68 Ma21 for the past 5 Ma (Wright & Walcott 1986). Note the close accord with volcanic rotation rate of about 78 Ma21 for the same period. A –A0 represents the profile discussed in (b). Locus for pole of rotation from Calhaem (1973b). (b) A projection of K–Ar dates onto profile A– A0 of (a). The gradient of the least squares fit to the data gives a value of 23+2 mm a21, which is the apparent southeasterly rate of migration of the active volcanic front. Note there are three new data points compared to the similar plot in Stern (1987): the unpublished age for Rolles Peak of 0.71 Ma (Wilson et al. 1995); the Motiti Island data point (Briggs et al. 2006), which appears to be an outlier and is circled (if this data point is excluded the fit is still 23+2 mm a21); and the andesite cone of Mt Hauhungatahi (just west of Mt Ruapehu), which has a published age of c. 0.8 Ma (Cameron et al. 2003).

A series of rotating and translating arcs, like an opening fan, was the first kinematic interpretation of K –Ar ages for low-potash andesites (Calhaem 1973a, b). Implicit in this interpretation are two assumptions: First, low-potash andesites erupt above a Wadati –Benioff zone where the depth to the zone is of the order of 100 km (Hatherton 1969; Syracuse & Abers 2006); second, the

delineation of older, coeval, linear arcs of andesites is tantamount to tracing previous positions of the plate boundary back in time. With this model, rotation rates for the Wadati –Benioff zone, with respect to a fixed Australian plate, have been calculated to be about 28 Ma21 for the period 16 –4 Ma, and 6+18 Ma21 from 4 Ma to the present (Calhaem 1973a; Stern et al. 2006). The apparent pole positions for the rotating arcs have also migrated (Fig. 3a). Some interpretations, however, rejected the concept of migrating volcanic arcs (Wilson et al. 1995) or a rotating subduction zone (Brothers 1984; Kamp 1984). With the advent of geodetic (Sissons 1979; Walcott 1984), GPS (Wallace et al. 2004), and palaeomagnetic data; (see following sections) the concept of a rotating, and internally deforming, eastern North Island, accompanied by asymmetric extension in the central North Island has, however, become clearer (Davey et al. 1995; Lamb 1988; Wallace et al. 2004). For much of the past 5 Ma the interpreted motion of the andesite arcs is both a rotation and a southward translation, with respect to a fixed Australian plate (Fig. 3a). A rotation rate of c. 78 Ma21 over the past 4 Ma is implied by the 308 angle subtended by the 4 and 0 Ma axes. The hinge point of the rotation appears to near Mt Ruapehu, and the apparent southward migration of the hinge point is 15–20 mm a21 (Fig. 3a). If the K –Ar ages are plotted against perpendicular distance from the active volcanic front, the average migration of the arcs in a southeasterly direction, is 23 +2 mm a21 (Fig. 3b). This result is an update with three new ages on that published previously when the rate was estimated at 20+2 mm a21 (Stern 1987). Most uncertainty in this analysis of rotation arcs concerns the position of the 4 Ma arc, as its southern extent and azimuth are difficult to assess because volcanic eruptives are progressively buried by ash and ignimbrites south of the interpreted southern extent of the 4 Ma arc. A 108 uncertainty in the azimuth of the 4 Ma arc would imply a migration rate over 4 Ma of 7+28 Ma21 (Stern et al. 2006).

Geodesy Strain estimates for central North Island were originally made from repeated geodetic observations that spanned just a few areas of the country (Walcott 1984). These data could be expressed as pure and simple shear, or as velocities with respect to either a fixed Pacific or Australian plate (Fig. 4a– c). Early estimates for the extension rates of 7–12 mm a21 across the CVR came from strain as derived from re-observing geodetic networks over a 50-year time interval (Sissons 1979;

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Fig. 4. (a & b) Shear strain components g1 and g2 based on triangulation data over the grids shown (Walcott 1984). Data based on re-observation of the grids after 70 years. (c) Velocities derived from shear strains for 70-year period and extrapolated over the whole country using numerical methods (Walcott 1984). Velocities with respect to a fixed Australian plate. (d) Velocities from GPS assumed the pt 2404 on the Australian plate is fixed. These data span 10 years of data from 1990 to 2000 (Wallace et al. 2004). The extension rates shown of 17+2 and 9+2 mm a21 are calculated for the eastern margin the CVR at the latitudes shown.

Walcott 1984). More recent GPS data that span a decade show much the same pattern (Fig. 4d) as the geodetic data, but indicate a higher extension rate of 17+2 mm a21 at the Bay of Plenty Coast with rates decreasing to 9+1 mm a21 at Lake

Taupo (Fig. 4d) (Wallace et al. 2004; Nicol & Wallace 2007). Geological data also show extension rates at the Bay of Plenty coast of 19+4 mm a21 over the past 1.5 Ma (Nicol & Wallace 2007), and 13+6 mm a21 (Lamarche et al. 2006). Both

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estimates overlap within the uncertainty range of the GPS estimate (17+2 mm a21). Some onshore extension rates from fault mapping are less than half that cited above from geodesy (Villamor & Berryman 2001). As suggested previously (Wallace et al. 2004), this difference could be because of the presence of hidden faults or dyke intrusion. The latter is particularly relevant given the dominant heat-flux from the CVR (see later section).

Relationship between extension and volcanic arc migration Migration of the low-K, andesite arc across central North Island is seen as a migration of the subjacent subducted pacific plate, and not necessarily a proxy for differential extension in the overlying Australian plate (Stern 1987). Nevertheless, the rate of extension from geological and GPS data (19+4 and 17+2 mm a21 respectively) (Nicol & Wallace 2007) (Fig. 4b) are only just significantly different from the 23+2 mm a21 arc-normal volcanic migration rate. It would appear from available data, therefore, that most of the arc-normal migration of low-potash andesites in the North Island is accompanied by back-arc extension in the overlying plate at a comparable rate.

Palaeomagnetic rotations Palaeomagnetic observations from the East Coast were made in the 1980s (Mumme et al. 1989; Wright & Walcott 1986) (Fig. 4a). Motivation for this work was largely driven by structural geology studies that describe a NW– SE orientated, Cretaceous –Palaeogene passive margin trend in the rocks of the northern portion of eastern North Island (the Raukumara Domain) (Fig. 5), yet the same rocks show a SW –NE trend in the mideastern North Island (the Wairoa Domain) (Lamb 1988; Stoneley 1968). This contrast in geological trend is consistent with the inferred differential rotations from palaeomagnetic data (Fig. 5). To the north the Raukumara Domain has experienced minor rotation only in the past 10 Ma, while in the south the Wairoa Domain has rotated up to 408, implying a shear zone between the two domains (Lamb 1988; Rowan et al. 2005). Because of inherent uncertainties several interpretations of the palaeomagnetic data have been given (Rowan et al. 2005): (1) the Wairoa Domain (Fig. 5) underwent a rotation rate of 18 Ma21 from 25–5 Ma, with respect to the Australian plate, then 6– 78 Ma21 from 5 Ma to present day; (2) the Wairoa Domain underwent no rotation until 5 Ma then rotated at 58 Ma21 until the present; (3) the Wairoa

Domain rotated at 48 Ma21 from mid-Miocene to present day. Options 1 and 2 are preferred here as they are consistent with rotations interpreted from andesite migration data (Fig. 3a) and with the observation that at c. 5 Ma central North Island was subjected to a rapid regional dome-like uplift (Pulford & Stern 2004). Such a surface uplift would have rapidly changed the potential energy of the lithosphere (Platt & England 1993) and thus provide a trigger for the onset of more rapid extension, and subsequent rotation. Rotations inferred from GPS data and a preferred geological model show values for eastern North Island almost half that cited above (Wallace et al. 2004). However, the choice of geological block model will affect the inferred rotation, whereas palaeomagnetic rotations are not model dependent.

Relationship between extension and rotation In order to explain the 5–78 Ma21 rotation rates in the Wairoa Domain and the early estimate of extension in the CVR of c. 10 mm a21 it was proposed that some of the rotation was effectively explained by differential strike-slip rates in the North Island Dextral Fault Belt (Figs 2 & 5) (Lamb 1988; Walcott 1987). Subsequent geological investigations found that strike-slip rates are only of the order of 4 mm a21 in the northern portion of the fault belt (Beanland & Haines 1998; Nicol & Wallace 2007). Therefore the observed rotation of the Wairoa Domain is most easily described as the rotation of a rigid body. For a rotation rate of, say, 68 Ma21 and the distance to the effective pole of rotation of c. 200 km (Fig. 3a), then the inferred displacement at the Bay of Plenty coast over 1 Ma is ru  21 km, where r is the 200 km distance to the pole of rotation (Fig. 3a) and u ¼ 6(2p/360) is the angle subtended (in radians) for 1 Ma. A 21 km displacement over 1 Ma at the Bay of Plenty coast is equivalent to a rate of 21 mm a21, which is in close accord with the GPS determined rates, the geological rate from faulting, and the volcanic migration rate. Hence, there is a close accord between short- (GPS  10 years) and long-term (4 Ma) indicators of kinematics in the central North Island, as has been noted previously (Nicol & Wallace 2007).

Crustal structure More than 20 years of active and passive source seismic exploration show that the crustal and upper mantle structure of the central North Island

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Fig. 5. A reconstruction for the Early Miocene based on palaeomagentic rotations and other geological observations (after King 2000; Lamb 1988; Walcott 1984). NIDFB, North Island Fault Belt (Fig. 2). Raukumara and Wairoa Domains with contrasting rotation properties are shown. At this time the CVR did not exist and the NIDFB and the Taranaki fault would have been absorbing some of the relative convergence and strike-slip motion between the Pacific (PAC) and Australian (AUS) plates. The shaded areas represent sedimentary depocentres with their age ranges given in Ma (Stern et al. 2006). A, A’ refers to interpretation cross-section in Figure 8. Inset shows present day situation with Wairoa Domain rotating as a unit in concert with the fan-like opening of the CVR. This is consistent with only small amounts of strike–slip motion in the NIDFB as discussed in text. Raukumara Domain is translating without rotation based on paleomagnetic data (modified from Lamb 1988).

varies radically (Fig. 6a). To the west of the CVR the crust is 25– 27 km thick but is underlain by a region of anomalously low upper-mantle P and S wave speeds (Stern et al. 1987; Horspool et al. 2006; Stratford & Stern 2006; Salmon 2008; Seward et al. 2008) and high seismic attenuation (Salmon et al. 2005). East of the CVR more normal crustal thicknesses of c. 35 km are present (Reyners et al. 2006; Stratford & Stern 2006). Within the CVR the crust and upper-mantle structure displays the most anomalous character compared to regular continental crust and upper mantle (Fig. 6a). Upper crust P-wave speeds are slow (6 km s21) down to a depth of c. 15 km (Stratford & Stern 2006). Below 15 km P-wave

speeds rapidly increase until values of 7.3– 7.5 km s21 are observed starting at a depth of about 20+2 km in the centre of the CVR (Stratford & Stern 2006). Studies of the Pn wave speed using earthquake sources also show these same wave speeds and that they must extend to depths of at least 50 –70 km into the upper mantle (Haines 1979; Reyners et al. 2006; Seward et al. 2008). Highly anomalous Vp/Vs ratios of up to 2.05 are proposed for the depth range 15 –30 km beneath the CVR, which implies c. 2% partial melt (Harrison & White 2004). One model is, therefore, that the CVR has a thin crust that is directly underlain by the asthenosphere, which has anomalously low Pn wave speeds due to 1–2% of partial melt.

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Fig. 6. (a) Crustal and upper mantle structure of the CVR along the profile A– A0 of Figure 3. Structural model and P-wave speeds based on wide-angle reflection and refraction work (Stratford & Stern 2006), Pn wave speeds from earthquake travel time data (Seward et al. 2008) and Vp/Vs ratio from earthquake data (Harrison & White 2004). Solid black triangle represents position of active volcanic front. (b) Crust and mantle seismic P-wave solution from earthquake tomography (Reyners et al. 2006). The deep blue colours represent the high P-wave seismic velocitites of the subducted Pacific plate. Black crosses represent locations for subduction zone earthquakes. Solid black triangle represents position of active volcanic front.

Between depths of 15 km and 20– 22 km P-wave speeds are 6.8 –7.3 km s21 and these are considered to be underplated new crust (Stratford & Stern 2006). Harrison & White (2004) proposed the underplated layer is thicker, extending from 15 to 30 km.

Earthquake tomography studies provide a broader and deeper view of crustal and mantle structure, although with less resolution, than controlled source seismology solutions (Fig. 6b) (Reyners et al. 2006). A comparison of the controlled source and tomographic solutions reveals

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broad similarities of an apparently thinned crust and low Pn velocities beneath the central North Island (Fig. 6a & b). In detail there are differences, however. For example, the tomographic model does not reproduce the extreme low P-wave velocities (1–3.2 km s21) in the top 2 km of volcanoclastic deposits of the CVR, or the higher Vp of .7 km s21 at depths 20 –35 km that the controlled source solution shows (Fig. 3a). This difference can be explained by the necessity for tomography to smooth through strong velocity variations and to give a smooth velocity variation with no discrete boundaries (Lay & Wallace 1995; Rawlinson & Sambridge 2003). Interpretations of the rocks with seismic P-wave velocities of 7.3 –7.5 km s21 beneath the central North Island include underplated crust (Harrison & White 2004) or anomalous upper mantle in a state of upper melt (Stratford & Stern 2006). There is, nevertheless, more agreement on the Moho being a broad (.5 km) transition zone rather than a discrete interface, and pervasive low P-wave velocities in the upper mantle (Fig. 6a & b). Additional support for partial melt in the upper mantle, and missing mantle lithosphere, beneath central North Island comes from seismic attenuation (Q21 p ) studies (Eberhart-Phillips et al. 2008; Salmon et al. 2005) and regional exhumation and uplift analysis (Pulford & Stern 2004). There is also geophysical evidence for partial melt in the crust of the CVR. A seismological receiver function study shows anomalously low S-wave speeds in the depth range of about 10 km, beneath the CVR, which can be interpreted to be due to the crust being in a state of partial melt (Bannister et al. 2004). Corroboration of the existence for crustal melts comes from recent magnetotelluric studies (Heise et al. 2007). Key points from the crustal structure results for the CVR are as follows. (1) There are no signs of rocks with P-wave velocities characteristic of greywacke-schists deeper than 15 km. Given a New Zealand crust, originally 30 km thick, where the top 3 km of crust is largely comprised of erupted volcanoclastic rocks, then the effective stretching factor (b) is at least 30/12  2.5 (McKenzie 1978). The b-factor could be higher than this if an unknown portion of the rocks above 15 km depth are made up of intrusions. (2) From a depth of c. 20 km rocks with anomalously low P-wave speeds of 7.3–7.5 km s21 are observed and these extend down into most of the mantle wedge. Where the boundary between crustal underplating and anomalous upper mantle sits is to some extent semantic and unclear. However, what is clear is that mantle asthenosphere sits a shallow level beneath the CVR, possibly at a depth of only

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20 km. This is likely to be a critical factor in explaining the high heat output from the CVR.

Heat output One of the most distinctive geophysical parameters from the central North Island is heat output from the Central Volcanic Region. Heat output of the region is almost totally expressed in the convective discharge of hot water and steam from a series of geothermal systems along its eastern side (Fig. 7a). Natural heat output was measured before the area was exploited for power production with the initial assessment being based on terrestrial measurements with calorimeters (Studt & Thompson 1969). These data were later complemented with lake and river measurements to give 4.3+0.5 109 W (Bibby et al. 1995). Because more heat could be escaping undetected via the water table and small river systems, the 4.3 GW estimate must still be considered a minimum. This estimate is similar to that for the Yellowstone caldera in the western US (Lowenstern & Hurwitz 2008). A useful means of comparison for regions of different length scales, which discharge heat via quasi-linear volcanic rift zones, is to normalize heat output by the length of the volcanic region. For example, dividing 4.3 GW by the 160 km length of the CVR one gets a normalized heat output of 0.026 GW (26 MW) per km of strike of volcanic zone. This is similar to the normalized heat output of Iceland (Palmason & Saemundsson 1974), which is an entirely different setting as it straddles an active oceanic spreading centre. The length of Iceland’s rift zone is c. 300 km, so the normalized MW/km comparison is useful as it suggests a heat sources of similar magnitudes. Heatflow is a concept appropriate to conductive regimes (Fowler 1990), whereas virtually all the heat discharged from the CVR is by advection. Nevertheless we can estimate an ‘effective’ heat flow by dividing the heat output by the area of down-flow for the active geothermal systems. The estimate for the down-flow region is somewhat subjective, but one measure is that area that includes all boreholes where zero geothermal gradient is observed (Studt & Thompson 1969; Stern 1987). Effective heat flow ¼ heat output/area of downflow ¼ 4.3 10 9 W/5000 km 2  860+100 mW m22. This is about 14 times greater than regular continental heat flow and 8 times greater than heat flow of typical continental back arc basins (Currie & Hyndman 2006b; Schellart 2007). Are we viewing a steady state heat-output from the CVR, or is some transient occurring?

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Fig. 7. (a) Map of CVR showing location and heat output for active geothermal fields and areas through which heat output is 4.3 GW. Older and now extinct fields are also shown. Dashed line encompasses active geothermal fields and bore holes where zero geothermal gradient are found (Studt & Thompson 1969; Stern 1987). Note that if the heat output is divided by the area encompassed between the dashed line and the active volcanic front the effective heat–flow is 860 mW m22. (b) Relationship between heat output and an intrusion rate as given by equation 1 in the text. Projected thicknesses of injected cooling layer required to explain heat output based on 10 mm a21 and 20 mm a21 rates of intrusion in the CVR.

A partial explanation for high heat flow from some continental back-arcs with low extension rates (Hyndman et al. 2005) is that they undergo rapid rotations and extension, for short periods,

followed by relative quiescence. Because heat conduction is slow, heat flow will remain high for many millions of years after the rapid extension has ceased. Examples of fan-like, rapid extension have

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been seen in the Miocene records of the Japan sea (Otofuji & Matsuda 1987), the Ligurian Sea between southern France and Sardina –Corsica (Faccenna et al. 2001) and the Kuril Basin (Schellart et al. 2003). Rapid fan-like extension appears to be the case at present for the CVR with 17 –20 mm a21 extension at its north end. So the fact that most continental back arcs show heat flows of about eight times less than the effective heat flow of the CVR could be because the CVR is in a particularly active stage. Attempts to explain the heat output of the Central Volcanic Region have included heat from plastic deformation (Hochstein 1995), cooling intrusives that accompany extension (Stern 1987), and conduction with a non-linear geothermal gradient (Weir 1998). Of these explanations the cooling intrusives model is easy to test with a simple numerical analysis (Beardsmore & Cull 2001) constrained by spreading rate, crustal structure and assumed thermal constants (thermal capacity, latent heat of melting); i.e. H ¼ r D S Vp ðDT Cp þ LÞ

(1)

where: H ¼ heat output (Watts), r ¼ density of intruded rock (2750 kg m23), D ¼ thickness of intruded layer (m), S ¼ strike length of spreading axis (160 km), Vp ¼ intrusion rate, DT ¼ temperature drop (800 8C), Cp ¼ specific heat (1000 J kg21 8C21 ) and L ¼ latent heat of intruded rock (3.2 105 J kg21) (McKenzie & Sclater 1969; Stuwe 2002). Equation 1 describes the steady state heat output from a rectangular volume of rock that grows in one direction at a rate of V m s21. H (in Watts) is graphically shown (Fig. 7b) as a function of intrusion rate and intrusion thickness (D). If the maximum extension rate of 17–20 mm a21 is effectively an intrusion rate, then a layer about 11 km thick is required to satisfy the observed heat output. A layer 11 km thick could be accommodated within the upper crust of the CVR between depths of say 3–15 km. Geological evidence shows that surface water can circulate to these depths (Wickham & Oxburgh 1985). Moreover, between 3–15 km deep in the CVR the geophysical data are consistent with the presence of partially molten rocks (Bannister et al. 2004; Heise et al. 2007). If a lower intrusion rate is used in equation 1 then the thickness of intruded rock gets larger and more problematic. For example, an intrusion rate of 10 mm a21 requires a layer 22 km thick (Fig. 7b). This is a troublesome result as it is not clear how circulating water could reach to depths of more than 20 km and efficiently extract the heat. It may, however, be that the process described by equation 1 is over simplified and more complicated

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heat transfer mechanisms are required (e.g. Weir 1998). Or meteoric fluid can reach to a depth greater than 20 km and efficiently extract heat from a circulating asthenosphere. Nevertheless, the plot of Figure 4b at least gives us a guide to visualizing the magnitude of the problem for explaining the present day heat output of the CVR as a steady state process.

Discussion Two important, and possibly generic, features of the central North Island back-arc system are seen in other continental back-arc regions (e.g. Fig. 1b); the fan like opening at plate tectonic rates, and the high heat output. No clear explanation has been given for the fan-like openings, and particularly, for why they are so rapid and episodic. Attempts to explain the high heat-flux of continental backarc regions include 3D mantle flow effects combined with non-uniform extension in time (Schellart 2007), or small-scale mantle convection (Currie & Hyndman 2006b). Mantle processes, including convection, can also be invoked to explain the high heat output and vertical movements of the central North Island. Like western North America (DeCelles 2004), the western North Island had a long (c. 25 Ma) period of compression, thrust faulting and shortening (Holt & Stern 1994; King 2000), prior to the rapid rotation, uplift (Pulford & Stern 2004) and extensional phase of the last 4– 5 Ma. Slow thickening over tens of millions of years then rapid release, in perhaps less than 1 Ma, of a mantle lithospheric instability beneath orogenic belts is recognized as viable mechanism for rapid uplift, followed by a switch from compressional to extensional tectonics (Platt & England 1993; Gemmer & Houseman 2007). Several tens of kilometres of shortening in 10 –20 Ma is sufficient for an instability of the mantle lithosphere to develop and eventually detach (Neil & Houseman 1999). Cross-sectional cartoons of this proposed process for central North Island are given in Figure 8. About 1 –10% of relative plate movement is proposed to have been taken up in the Taranaki fault zone during the late Oligocene–Miocene (Fig. 8a & b), where an old Permian suture (Mortimer 2004) through western North Island was exploited as a zone of weakness. By mid-Miocene the mantle had thickened sufficiently to create downward displacement on the Earth’s surface creating a series of quasi-circular basins (Figs 5 & 8c). After detachment (Fig. 8d) a rapid and regional uplift is predicted, which is documented to have started at 5 Ma from geological evidence (Pulford & Stern 2004). Thereafter extensional

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Fig. 8. Conceptual cross-sectional cartoons for the development of mantle thickening, instability removal, surface uplift and then back-arc extension in the central North Island of New Zealand. Cross section is related to A–A0 of Figure 5.

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and rotational rates increased. Deep (600 km) earthquakes beneath the Taranaki fault region (Fig. 2) may represent the last vestige of the detached mantle instability (Stern et al. 2006). If thickening is progressive and detachment of thickened and unstable mantle lid is episodic along an arc, as argued for western North Island in the Miocene (Stern et al. 2006) (Fig. 8), then the rapidly extending lithosphere will be pinned at the end where there is still thickened and attached mantle lithosphere: the pinning point will act like a rotation pole. The apparent rotation pole for the North Island volcanic arcs is just south of the line between Mts Taranaki and Ruapehu (Fig. 3a), where there is a sudden jump in lithospheric thickness (Stern et al. 2006; Eberhart-Phillips et al. 2008; Salmon 2008). Hence the rapid openings in the form of a fan, and the extreme heat pulse, as seen today in the North Island of New Zealand.

Conclusions (1) A combination of geodetic, palaeomagnetic, crustal structure and heat flux data from central North Island, New Zealand provides an opportunity to examine both kinematics and dynamics of a continental back-arc extensional zone. (2) Continental back-arc extension in North Island is characterized by two distinctive, and possibly generic, features: rapid fan-like opening for a period of a few million years, and extremely high heat output. (3) We suggest that both the fan-like opening and the high flux are related to the geological history of active continental margins where compression and mantle thickening tend to dominate for the initial 10 –20 Ma of any given system. The rapid release of a mantle instability provides the requisite potential energy increase to initiate and sustain back-arc spreading for a period of a few million years. High heat flux results from the mantle melts being emplaced right to the base of a now thinned continental crust. I thank G. Lamarche, S. Lamb and W. Stratford for constructive reviews. M. Reyners is thanked for permission to use Fig. 6b.

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EXTENSION, NORTH ISLAND, NEW ZEALAND magmas: evidence from Egmont Volcano, North Island, New Zealand. Journal of Petrology, 40, 167–197. P ULFORD , A. & S TERN , T. A. 2004. Pliocene exhumation and landscape evolution of central North Island, New Zealand: the role of the upper mantle. Journal of Geophysical Research, 109, doi10.1029/2003jf000046. R AWLINSON , N. & S AMBRIDGE , M. 2003. Seismic traveltime tomography of the crust and lithosphere. Advances in Geophysics, 46, 81– 197. R EYNERS , M. 1980. A microearthquake study of the plate boundary, North Island, New Zealand. Geophysical Journal of the Royal Astronomical Society, 63, 1–22. R EYNERS , M., E BERHART -P HILLIPS , D., S TUART , G. & N ISHIMURA , Y. 2006. Imaging subduction from the trench to 300 km depth beneath the central North Island, New Zealand, with Vp and Vp/Vs. Geophysical Journal International, 165, 565– 583. R OWAN , C. J., R OBERTS , A. P. & R AIT , G. J. 2005, Relocation of the tectonic boundary between the Raukumara and Wairoa Domains East Coast, North Island, New Zealand: implications for the rotatio history of the Hikurangi margin. New Zealand Journal of Geology and Geophysics, 48, 185 –196. S ALMON , M. L. 2008. Crust and upper mantle inhomgenities beneath western North Island, New Zealand: evidenece from seismological and electromagnetic data. PhD thesis, Victoria University of Wellington, Wellington. S ALMON , M., S AVAGE , M. K. & S TERN , T. A. 2005. Seismic attenuation, temperature, H20, mantle melting and rock uplift, Central North Island New Zealand. Eos Transactions of AGU, 86, Abstract T13B-0478. S CHELLART , W. P. 2007. Comment on “the thermal structure of subduction zone back arcs” by Claire A. Currie and Roy D. Hyndman. Journal of Geophysical Research, doi:10.1029/2007JB005287. S CHELLART , W. P., J ESSELL , M. W. & L ISTER , G. S. 2003. Asymmetric deformation in the backarc region of the Kuril arc, northwest Pacific: new insights from analogue modelling. Tectonics, doi:10.1029/ 2002TC001473. S EWARD , A. M., H ENDERSON , C. M. & S MITH , E. G. C. 2008. Models of upper mantle beneath the Central North Island, New Zealand, from speeds and anisotropy of sub-horizontal P-waves (Pn). Journal of Geophysical Research, 114, B01301; doi: 01310.01029/ 02008JB005805. S ISSONS , B. A. 1979. The horizontal kinematics of the North Island of New Zealand. Victoria University of Wellington, Wellington. S TERN , T. A. 1987. Asymmetric back-arc spreading, heat flux and structure of the Central Volcanic Region of New Zealand. Earth and Planetary Science Letters, 85, 265–276. S TERN , T. A., S MITH , E. G. C., D AVEY , F. J. & M UIRHEAD , K. J. 1987. Crustal and upper mantle structure of northwestern North Island, New Zealand,

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Cretaceous felsic volcanism in New Zealand and Lord Howe Rise (Zealandia) as a precursor to final Gondwana break-up A. J. TULLOCH1*, J. RAMEZANI2, N. MORTIMER1, J. MORTENSEN3, P. VAN DEN BOGAARD4 & R. MAAS5 1

GNS Science, Private Bag 1930, Dunedin, New Zealand

2

Department of Earth, Atmospheric & Planetary Sciences, Massachusetts Institute of Technology, Cambridge MA 02139, USA

3

Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, BC, V6T 1Z4 Canada 4

5

IFM-GEOMAR, Leibniz Institute for Marine Sciences, Wischhofstr. 1-3, D-24148 Kiel, Germany

Department of Earth Sciences, University of Melbourne, Melbourne VIC 3010, Australia *Corresponding author (e-mail: [email protected]) Abstract: We report new, highly precise, U –Pb and Ar/Ar ages for seven Cretaceous rhyolites, tuffs and granites from across Zealandia spanning a 30 Ma period from arc magmatism to continental break-up. Combined with previously published data, these reveal a strong episodicity in Cretaceous silicic magmatism outside the Median Batholith. 112 Ma tuffs are known only from the Eastern Province in association with a Cretaceous normal fault system. Both 101 and 97 Ma groups of rhyolites and tuffs occur across the entire width and half the length of Zealandia from near the palaeotrench to the continental interior, indicating widespread and effectively instantaneous extension. We attribute an increase in A-type character with time (112–101– 97– 88–82 Ma) to the progressive thinning of the Zealandia continental crust whereby, with time, there is less opportunity for crustal contamination. Extension directions associated with 101, 97 and 82 Ma magmatism and associated core complex exhumation across Zealandia are all oriented c. 308 oblique to the margin. These observations suggest Zealandia rifting was controlled by either .83 Ma capture of Zealandia by the Pacific Plate and/or ,83 Ma Zealandia– West Antarctica spreading, rather than by laterally migrating triple junctions, slab windows or plume heads.

The continent of Zealandia split from Gondwanaland in the Late Cretaceous at about 83 Ma (Sutherland 1999; Mortimer 2004a; Laird & Bradshaw 2004; Davy 2006). Zealandia, about two-thirds the size of the conterminous United States, remains largely submerged as a result of crustal thinning and extension associated with pre-break-up Cretaceous tectonics. Only the convergence and crustal thickening associated with development of the Neogene Australia – Pacific plate boundary through Zealandia allowed a small part of this continent to rise above sea level to form the main islands of New Zealand (Fig. 1). The pre-Late Cretaceous rocks of Zealandia can be divided into an Eastern Province and Western Province, separated by the Median Batholith. New Zealand’s Eastern Province represents a forearc – accretionary wedge assemblage in which sedimentation and deformation was active from the Permian to at least the late Early Cretaceous

(Coombs et al. 1976; Bradshaw 1989; Mortimer 2004a). Major units in the Eastern Province include the greywacke-dominated Torlesse Terrane and its metamorphosed equivalent, the Otago Schist (Forsyth 2001). The Western Province represents a much earlier, Early Palaeozoic, phase of accretion; by the Permian the Western Province was stabilized Gondwana crust. The Median Batholith is a composite, long-lived Devonian –Early Cretaceous Cordilleran batholith that contains the record of most of Zealandia’s subduction-related magmatism (Mortimer et al. 1999). Tulloch & Kimbrough (2003) describe the 230–105 Ma phase of the batholith as comprising two parallel plutonic belts: a 230–130 Ma eastern (oceanward) belt and a 130–105 Ma western (continentward) belt with distinctive high Sr/Y chemistry. The exact time in the Cretaceous at which there ceased to be a subducting slab beneath the Eastern Province and Median Batholith remains uncertain.

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 89– 118. DOI: 10.1144/SP321.5 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Sample location map. Black dots are sites for which we report new U– Pb or Ar/Ar dates in this paper. Median Batholith (230–105 Ma) separates Permian–Jurassic accreted terranes (EPr, Eastern Province) from Early Palaeozoic Gondwanan terranes (WPr, Western Province). HPl is Hikurangi Plateau.

If there is a consensus age it would be 100 Ma, a time that divides the youngest Median Batholith plutons (Tulloch & Kimbrough 2003) and youngest Eastern Province sediments seemingly deposited in a trench environment (Cawood et al. 1999), from the oldest intraplate basaltic magmatism in Zealandia (Weaver & Pankhurst 1991; Tulloch 1991; Baker et al. 1994). However ages as old as 128 Ma, the youngest age of ‘normal’ subductionrelated, I-type, low Sr/Y magmatism in the Median Batholith (Tulloch & Kimbrough 2003) or as young as 82 Ma, the age of youngest documented thrusting in the Eastern Province accretionary prism (Mazengarb & Harris 1994), could also be considered for cessation of subduction. The recognition of minor but widespread 95– 100 Ma intraplate mafic magmatism in Zealandia has been an important peg in deciphering the tectonomagmatic development of Cretaceous Zealandia. In this paper we add to this discussion by presenting new, high-accuracy and precision, U –Pb zircon and Ar/Ar biotite ages, along with geochemical and tracer isotopic data, for eight felsic igneous rocks, mainly rhyolites and tuffs, from Zealandia. All these samples are from outside the main region of Cretaceous silicic magmatism, the Median

Batholith. We discuss the significance of our new results in the context of earlier data and models of Cretaceous Zealandia siliceous magmatism in the interval 130-83 Ma (Muir et al. 1998; Waight et al. 1997, 1998a, b; Tulloch & Kimbrough 2003). Our results confirm the widespread eruption of intraplate (A-type) magmas at c. 100 Ma across Zealandia, and thus suggest absence of a subducting slab beneath Zealandia at this time. Our results also reveal distinctively older and younger episodes of magmatism that require different tectonic explanations.

Samples All of our dated samples come from outside the Median Batholith (Fig. 1), from both the trenchward and continent-ward side of the batholith. Some rocks were collected specifically for this study, and some were obtained from existing collections. Samples prefixed P, R and OU are archived in the GNS P, GNS R and Otago University rock collections. Sample and location details for these rocks are available on the PETLAB database (http://pet. gns.cri.nz/).

CRETACEOUS FELSIC VOLCANISM

Methods Zircon and biotite were separated from the rock samples by standard crushing, heavy liquid and magnetic separation techniques, and were subsequently hand-picked under a binocular microscope based on purity and crystal morphology. U– Pb ID-TIMS (isotope dilution thermal-ionization mass spectrometry) analyses were carried out at the Massachusetts Institute of Technology (MIT), U –Pb LA-ICPMS (laser ablation inductively coupled plasma mass spectrometry) analyses at University of British Columbia (UBC) and Ar/Ar analyses at IFM-GEOMAR (Kiel). Analytical and data reduction methods for U –Pb ID-TIMS, U– Pb LA-ICP-MS and Ar/Ar dating are described in the Appendix. It should be noted that Ar/Ar analyses result in ‘cooling ages’ that mark the time when the rock cooled below the closure temperature of Ar in the analysed mineral (e.g. biotite). This is in contrast to U –Pb zircon ages that represent the crystallization event. The use of Ar/Ar technique in this study, however, is limited to a discrete volcanic eruption (Shag – Eweburn) and does not engender a detailed comparison with the bulk of the U –Pb zircon ages.

Eastern Province The Houhora Complex. The Houhora Complex of the Early Cretaceous Mount Camel Terrane of northernmost New Zealand comprises basalt and basaltic andesite lava, pillow lava, rhyolitic tuff, tuff breccia, conglomerate, sandstone and mudstone (Isaac et al. 1994; Isaac 1996). The terrane is structurally and compositionally distinct from the adjacent Waipapa Terrane and has no direct counterpart anywhere in New Zealand. The boundary with the Waipapa Terrane is not exposed. Fold axes trend WSW–ENE, normal to the inferred Gondwana margin, suggesting that the Houhora Complex was not simply accreted at the convergent margin. But correlation with similar rocks on Three Kings Islands and Reinga Ridge to the NW (Mortimer et al. 1998) does suggest a Gondwana marginparallel distribution. The dated ignimbrite sample (P52296) is a greenish heterogeneous crystal lithic vitric tuff, variable in texture and lithic content. A lithic clast dominating one thin section has minor plagioclase, up to 1 mm in size, set in a trachytic groundmass. A second thin section is dominated by banding, some of which may have recrystallized shard texture. The third and most homogeneous thin section exhibits fine-grained granophyric texture with minor K-feldspar phenocrysts as large as 1 mm. Zircon is characterized by short prisms (length to breadth ratio approximately 2:1). Secondary quartz and chlorite are abundant.

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Whatuwhiwhi Formation unconformably overlies Houhora Complex and contains a basal conglomerate of Houhora Complex-derived granitoids (Isaac 1996). We dated a granite clast (P52300) from the Whatuwhiwhi Formation (Isaac 1996) which overlies the Houhora Complex. The sample consists of homogeneous granophyre, similar in texture to parts of the P52296 tuff. Both Houhora Complex and Whatuwhiwhi Formation are believed, on the basis of poorly dated fossils and regional correlations, to be Cretaceous in age (Isaac 1996). The Motu Tuff. It occurs in the lower part of the Matawai Group in Raukumara Peninsula of the North Island. Motu River is the type locality for the Motuan Stage (100 –103 Ma) of the New Zealand Cretaceous (Crampton et al. 2004). The unit includes three 1.5–5 m thick pale weathering tuff beds. The dated bed (T2; X16/1353 1779; P63724) is a pale yellowish weathering, wavy laminated, redeposited crystal vitric tuff of rhyolitic composition. In thin section it is homogeneous, contains well-preserved glass shards and little terrigenous material is apparent. Crystal fragments (c. 3% volume, 0.05 –0.15 mm in size) are quartz, minor fresh plagioclase and rare, mostly altered, biotite. A U –Pb zircon age of 101.6 + 0.2 Ma was reported in Crampton et al. (2004). The Eweburn Tuff. The Eweburn Tuff (and Shag Valley Ignimbrite, see below) occur approximately 50 km apart along the NW– SE-trending Waihemo Fault system (Forsyth 2001). The Eweburn Tuff forms a .8 m thick stratum towards the base of the 3000 m thick Kyeburn Formation of the basal cover sequence (Matakea Group). Kyeburn Formation consists of c. 80 km2 of non-marine schist derived fanglomerates, conglomeratic fluviatile sediments and minor lacustrine mudstones (Bishop & Laird 1976). It unconformably overlies Torlesse greywacke and low grade Otago Schist, and is downfaulted into the schist in a graben-like structure controlled by the aforementioned Waihemo Fault, and by the NE– SW striking Danseys Pass Fault (Forsyth 2001). The dated sample (P77501 ¼ R7583) consists of coarse grained lithic crystal vitric tuff with vitroclastic flow texture. Crystals (0.1–1.0 mm in size, 20% by volume) are plagioclase, quartz, greenish-brown biotite, minor opaque oxide, green hornblende, and possible pale green diopsidic clinopyroxene. Vitric shards are c. 0.2 mm in size. Biotite from the same sample (KY4; R7583) was originally dated by K – Ar at 104 + 2 Ma) and two other samples from this locality yielded K –Ar ages of 107 + 2 Ma and 108 + 3 Ma (Adams & Raine 1988). The Shag Valley Ignimbrite. It forms a .20 m thick flow within the Horse Range Formation of the

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Matakea Group (Steiner et al. 1959; Forsyth 2001). The c. 400 m thick non-marine formation occupies a series of elongate strips in half-graben associated within strands of the Waihemo Fault system. The ignimbrite consists of lithic crystal vitric tuff. Rare pumice and metasedimentary lithic clasts and crystal fragments (0.1–1.0 mm, 15%) of plagioclase, quartz and brown biotite are set in a devitrified vitroclastic groundmass (P22174). Biotite flakes are commonly flow-aligned. The dated ignimbrite sample (P77502) consist of biotite-phyric moderately to densely welded tuff with rare lithic fragments up to 1 cm in size. In thin section the rock consists of whole and broken plagioclase and variably resorbed quartz crystals along with tabular grains of fresh brownish biotite, all in a densely welded matrix. The vitric shards are considerably smaller than those in the sample of Eweburn Tuff. The Shag Valley Ignimbrite matrix is weakly clay-altered, however feldspar and biotite phenocrysts remain fresh.

Western Province The Stitts Tuff (Nathan et al. 2002). It comprises the basal 60 m of the non-marine Pororari Group (.5100 m, Nathan 1978). Above the Stitts Tuff, most of the group consists of coarse, rapidly deposited detritus which records the surface breaching of the Paparoa Metamorphic Core Complex (Tulloch & Kimbrough 1989; Tulloch & Palmer 1990). The Stitts Tuff is a fine-grained white or grey rhyolitic vitric tuff with occasional beds of dark carbonaceous shale, deposited in a lake. The dated sample (P49125) is a finely laminated, well sorted, crystal vitric tuff. Angular clasts (c. 5 –10%; 50– 100 um), dominated by quartz, and lesser feldspar, with traces of altered biotite and zircon are set in a devitrified groundmass, characterized by abundant glass shad pseudomorphs. Muir et al. (1997) reported AS3-calibrated SIMS U –Pb zircon ages of 101 + 2 Ma (MSWD 1.9) and 102 + 3 Ma (MSWD 1.3). The DSDP 207A rhyolite. It represents one of only three pre-Cenozoic basement samples ever recovered from the Lord Howe Rise (Mortimer 2004b; Mortimer et al. 2008). This Deep Sea Drilling Program hole penetrated 513 m below sea level and bottomed in 156 m of rhyolitic volcanic rocks (van der Lingen 1973). Our own examination of thinsections of core reveal glassy rhyolite that varies considerably in its texture, degree of devitrification and (cognate volcanic) lithic content. No mafic minerals were observed although a pseudomorph may have replaced amphibole in P63857. No foreign lithic clasts were observed that might represent pre-rhyolite basement. The dated sample (P63855;

369.13–369.39 m) consists of relatively fresh perlitic rhyolite. Phenocrysts of quartz and sanidine (5% volume, up to 3 mm in size) are set in a light brown essentially isotropic glass showing flow banding, but little sign of devitrification. The phenocrysts are often partly euhedral, and resorbed embayments are common. Opaque oxide and zircon are common accessory minerals. The K –Ar dates on sanidine from DSDP 207A rhyolite of 95-97 Ma (calculated with new decay constants) were reported by McDougall & van der Lingen (1974). The Whataroa Granite. It forms a c. 4  1 km sliver immediately west of the Alpine Fault in southern Westland (Cox & Barrell 2007). The dated sample (P52280) is a coarse-grained K-feldspar megacrystic granite exhibiting ductile deformation, overprinted with brittle shearing and epidote alteration. Primary biotite and ?amphibole has been replaced with olive-green fine-grained biotite in shear bands. There is abundant but accessory titanite, allanite and zircon. The shearing is likely related to the nearby Alpine Fault; no country rocks are exposed in contact with the Whataroa Granite. The Canavans Quartz Monzonite. It crops out over several hundred square metres on the northern side of Canavans Knob (Wooding 1984; Cox & Barrell 2007). The dated sample (P45553) consists of fine –medium-grained two-pyroxene– hornblende– biotite quartz monzogranite –monzodiorite. Biotite forms conspicuously strongly red-brown large plates; hornblende and pyroxene are often smaller, anhedral equant and rounded, sometimes forming numerous inclusions in feldspar. The rock is conspicuously undeformed, but some mafic banding and dykes are present. More mafic variants have similarities with the appinitic rocks of the Victoria Range. Minor ne-normative mafic dykes (Wooding 1984) suggest these rocks post-date convergent margin magmatism of the Median Batholith, but it is possible that the dykes are significantly younger than the granitic rocks.

New geochronology results Eastern Province Houhora Complex. Nine air-abraded single zircon grains were analysed from the Houhora ignimbrite sample P52296 by the ID-TIMS method at MIT (Table 1; Fig. 2). Two distinct zircon populations can be recognized in this sample. The older population consists of variably discordant zircons with 207Pb/206Pb dates ranging from c. 136 Ma to c. 991 Ma. The younger group is an Early Cretaceous cluster of five mainly concordant analyses with 206 Pb/238U dates in the 101.71– 101.92 Ma range.

Table 1. U –Pb TIMS isotopic analyses of zircon from felsic igneous rocks Ratios Sample Fractions Footnote (a)

c

Pb (pg) (b)

Pb* Pbc (b)

Th U –

206

208

206

204Pb

206Pb

238U

(c)

(d)

(e)

err (2s%) –

Pb

Pb

Pb

207

Pb

Age (Ma)

(e)

err (2s%) –

235U

207

Pb

(e)

err (2s%) –

206Pb

206

207

207

238U

235U

206Pb

–

–

–

Corr. coef. –

Pb

Pb

Pb

0.76 0.06 0.49 0.34 0.27 0.81 0.64 0.61 0.62

853.0 19985.8 1011.8 420.0 639.9 429.5 547.2 1058.6 1473.4

0.244 0.024 0.158 0.115 0.095 0.261 0.203 0.196 0.198

0.015890 0.114193 0.018743 0.056445 0.059017 0.015944 0.015842 0.015929 0.015905

(0.10) (0.09) (0.22) (0.15) (0.22) (0.19) (0.30) (0.10) (0.09)

0.10539 1.13626 0.12561 0.42219 0.44619 0.10595 0.10480 0.10584 0.10573

(0.44) (0.14) (0.51) (0.66) (0.65) (1.14) (0.73) (0.34) (0.27)

0.04811 0.07217 0.04860 0.05425 0.05483 0.04819 0.04798 0.04819 0.04821

(0.41) (0.10) (0.45) (0.61) (0.58) (1.07) (0.64) (0.31) (0.25)

101.63 697.05 119.71 353.97 369.65 101.97 101.32 101.88 101.72

101.74 770.73 120.14 357.62 374.61 102.25 101.20 102.16 102.05

104.5 990.6 128.8 381.4 405.4 108.8 98.3 108.7 109.7

0.395 0.680 0.493 0.434 0.471 0.453 0.496 0.420 0.415

P52300 Whatawhiwhi granitic clast z1 aa 0.8 7.1

0.64

431.9

0.207

0.015837

(0.15)

0.10595

(0.83)

0.04852

(0.79)

101.29

102.26

124.8

0.406

P49125 Stitts Tuff z3 aa z4 aa z2 aa z7 aa z6 aa z5 aa z9 ca z10 ca z12 ca z13 ca

17.0 39.3 21.2 3.6 23.1 10.1 3.7 9.4 13.7 9.1

0.31 0.23 0.22 0.59 0.23 0.76 0.72 0.62 0.83 0.72

1087.3 2536.5 1373.5 227.6 1510.7 590.6 231.6 569.9 775.8 535.0

0.103 0.083 0.086 0.206 0.075 0.243 0.235 0.197 0.268 0.233

0.072172 0.071657 0.061972 0.059976 0.057864 0.015687 0.015668 0.015663 0.015680 0.015667

(0.11) (0.10) (0.10) (0.16) (0.09) (0.14) (0.23) (0.11) (0.10) (0.10)

0.56270 0.56545 0.48477 0.45541 0.43157 0.10417 0.10529 0.10362 0.10424 0.10448

(0.32) (0.17) (0.25) (0.91) (0.25) (0.71) (1.82) (0.77) (0.67) (0.79)

0.05655 0.05723 0.05673 0.05507 0.05409 0.04816 0.04874 0.04798 0.04821 0.04837

(0.29) (0.13) (0.22) (0.86) (0.22) (0.66) (1.73) (0.73) (0.63) (0.75)

449.23 446.13 387.61 375.48 362.63 100.34 100.22 100.19 100.30 100.21

453.28 455.07 401.34 381.06 364.29 100.62 101.64 100.11 100.68 100.90

473.9 500.4 481.2 415.1 374.9 107.3 135.1 98.3 109.8 117.3

0.440 0.616 0.486 0.378 0.467 0.414 0.472 0.431 0.398 0.450

P63855 DSDP 207A rhyolite z1 aa 1.0 7.8 z4 aa 1.2 18.2 z3 aa 0.6 24.3 z2 aa 0.9 17.6

0.54 0.57 0.60 0.58

485.3 1101.9 1444.4 1061.5

0.174 0.184 0.193 0.185

0.015160 0.015173 0.015168 0.015167

(0.11) (0.08) (0.11) (0.09)

0.10148 0.10053 0.10085 0.10041

(0.89) (0.31) (0.48) (0.40)

0.04855 0.04805 0.04822 0.04801

(0.84) (0.29) (0.45) (0.38)

97.00 97.07 97.05 97.04

98.14 97.27 97.56 97.15

126.0 102.0 110.0 99.8

0.472 0.379 0.316 0.337

0.7 1.3 0.4 1.5 0.9 0.5 1.4 0.9 0.3 0.8

93

(Continued)

CRETACEOUS FELSIC VOLCANISM

P52296 Houhora ignimbrite z1 aa 0.7 14.7 z2 aa 0.9 298.4 z3 aa 1.0 16.4 z4 aa 1.3 6.4 z5 aa 1.7 9.7 z6 aa 1.1 7.3 z7 aa 0.8 9.0 z8 aa 0.5 17.7 z10 aa 0.7 24.7

94

Table 1. Continued Ratios c

208

206

204Pb

206Pb

238U

(c)

(d)

(e)

err (2s%) –

Th U –

5.4 5.1 146.0 50.3 30.0

0.83 0.39 0.68 0.72 0.99

320.3 337.0 8444.7 2886.1 1621.5

0.265 0.126 0.217 0.231 0.317

0.013784 0.013793 0.013783 0.013763 0.013748

OU50106 Canavans quartz monzonite z1 ca 0.5 25.8 1.45 z2 ca 0.5 26.3 1.68 z3 ca 0.5 45.7 1.37

1262.3 1224.9 2256.4

0.465 0.538 0.440

0.015791 0.015771 0.015781

P52280 Whataroa granite z3 ca 0.7 z5 ca 3.5 z1 ca 0.2 z4 ca 0.8 z2 ca 0.2

Pb

Pb

Pb

207

Pb

Age (Ma)

(e)

err (2s%) –

(0.15) (0.09) (0.08) (0.09) (0.10)

0.09112 0.09178 0.09090 0.09101 0.09054

(0.09) (0.09) (0.09)

0.10468 0.10460 0.10471

235U

207

Pb

206

207

207

238U

235U

206Pb

–

–

–

Corr. coef. –

Pb

Pb

Pb

(e)

err (2s%) –

(1.37) (0.70) (0.13) (0.23) (0.36)

0.04795 0.04826 0.04783 0.04796 0.04776

(1.31) (0.66) (0.10) (0.21) (0.33)

88.25 88.31 88.24 88.12 88.02

88.55 89.16 88.34 88.44 88.01

96.6 112.0 91.0 97.3 87.5

0.506 0.443 0.628 0.443 0.406

(0.33) (0.34) (0.30)

0.04808 0.04810 0.04812

(0.30) (0.31) (0.28)

101.00 100.87 100.94

101.09 101.02 101.12

103.3 104.4 105.3

0.405 0.379 0.354

Notes: Corr. coef., correlation coefficient. Age calculations are based on the decay constants of Steiger & Ja¨ger (1977). (a) All analyses are single zircon grains. aa, air-abraded zircon; ca, thermally annealed and chemically treated zircon. (b) Pbc is total common Pb in analysis. Pb* is radiogenic Pb concentration. (c) Measured ratio corrected for spike and fractionation only. (d) Radiogenic Pb ratio. (e) Corrected for fractionation, spike, blank, and initial common Pb. Mass fractionation correction of 0.25%/amu + 0.04%/amu (atomic mass unit) was applied to single-collector Daily analyses. Total procedural blank less than 1.0 pg for Pb and less than 0.1 pg for U. Blank isotopic composition: 206Pb/204Pb ¼ 18.31 + 0.53; 207Pb/204Pb ¼ 15.38 + 0.35; 208Pb/204Pb ¼ 37.45 + 1.1.

206Pb

A. J. TULLOCH ET AL.

Pb* Pbc (b)

Sample Fractions Footnote (a)

Pb (pg) (b)

206

CRETACEOUS FELSIC VOLCANISM

95

Fig. 2. TIMS U– Pb condordia plots generated using Isoplot 3.0 (Ludwig 2003). All data point error ellipses are 2s and represent single crystal zircon analyses (MIT). Dashed lines represent the U-decay constant error envelope about the concordia. See text for the age error inclusive of tracer calibration and decay constant errors.

The latter are interpreted to represent the crystallization age of zircon in this rock. Because of excess age scatter among the younger zircon population, most probably due to Pb loss, no reliable weighted mean age can be calculated from the available

data. The two oldest analyses of this cluster (least affected by Pb loss) gave 206Pb/238U dates of 101.88 + 0.10 Ma and 101.97 + 0.19 Ma. Thus, an approximate crystallization age of 101.9 Ma is assigned to the Houhora ignimbrite. A single

96

A. J. TULLOCH ET AL.

zircon grain from a granite clast (P52300) in the overlying Whatuwhiwhi Formation yielded a 206Pb/238U date of 101.29 + 0.15 Ma in the same range as those of the younger zircon cluster of P52296. Eweburn Tuff. 40Ar/39Ar biotite analytical data obtained at Kiel University are compiled in Table 2. Biotite mineral separates KY4bts (1.762 mg), KY4bt2 (1.623 mg), and KY4bt3 (1.789 mg) were incrementally heated by increasing the laser power from 0.05 to 15 W. Gas fractions from each of the 20 heating steps were analysed and yielded integrated ages of 114.0 Ma, 113.3 Ma, and 114.9 Ma, respectively. Their age spectra (Fig. 3) show plateaus in the mid-hightemperature range that comprise 45.6% (n ¼ 5 steps), 64% (n ¼ 6 steps) and 66.5% (n ¼ 8 steps) of the total 39Ar released. Plateau ages are calculated as 112.4 + 0.4 Ma (KY4bts), 112.4 + 0.4 Ma (KY4bt2), and 112.9 + 0.5 Ma (KY4bt3) (2s uncertainties). Multiple disturbances with excess 40 Ar/39Ar ratios are observed in the lowtemperature heating steps (reflecting alteration and/or 39Ar recoil loss during the irradiation) and high temperature heating steps (reflecting the degassing of mineral inclusions and/or excess 40Ar). In addition, all biotite age spectra display sharp negative peaks in the low-temperature range with apparent ages up to 30 Ma less than the plateau age, possibly indicating post-eruptive thermal overprinting and partial resetting of the Eweburn Tuff biotites at 84 + 9 Ma (weighted average of minimum apparent age steps). Age results from biotite stepheating plateaus are coeval within two-sigma uncertainties, therefore a composite inverse-variance weighted mean age can be calculated. This average is 112.5 + 0.2 Ma (2s, MSWD 1.5) and provides our best age estimate for the eruption of the Eweburn Tuff. This is c. 5 Ma older than the oldest K –Ar age reported by Adams & Raine (1988). Shag Valley Ignimbrite. Zircons separated from a 2 kg sample of the ignimbrite (P77502) comprised relatively coarse, elongate, euhedral, pale-yellow crystals. A total of 16 individual analyses were done using line scans by LA-ICP-MS at UBC (Table 3). The data yield a slightly skewed Gaussian distribution of 206Pb/238U ages on a relative probability plot which we presume is due to Pb-loss (Fig. 3). After removal of two outlying ages, the data yield a weighted average 206Pb/238U age of 112.3 + 0.4 Ma (2s, MSWD ¼ 0.83). We consider this age to be the age of eruption of the tuff. It is c. 11 Ma older than the K –Ar ages reported in Adams & Raine (1988).

Western Province Stitts Tuff. Ten single zircon grains from the tuff sample P49125 were analysed by the ID-TIMS method at MIT, four of which had been pre-treated by the chemical abrasion (CA-TIMS) technique. The latter, together with one air-abraded analysis, comprise a coherent cluster that yields a weighted mean 206Pb/238U age of 100.250 + 0.054 [0.18] Ma (MSWD ¼ 1.1), which we report as the age of crystallization of the magma forming the tuff. This age agrees within uncertainties with AS3-calibrated, U –Pb SHRIMP (Sensitive High Resolution Ion Microprobe) ages of 101 + 2 Ma and 102 + 3 Ma (Muir et al. 1997). The remaining five analyses are distinguishably older and variably discordant with 207 Pb/206Pb dates of 500 Ma, 481 Ma, 474 Ma, 415 Ma and 375 Ma. These xenocrystic zircon age groups are consistent with those from both the Cambro-Ordovician Ross Orogen, which form a significant detrital population in the Western Province greywacke basement (Ireland 1992; Ireland & Gibson 1998) and the Devonian Karamea Suite plutons (Muir et al. 1996; Tulloch et al. 2009) that intrude the greywacke basement. DSDP 207A. Four air-abraded single zircon ID-TIMS analyses (MIT) overlap within 2s error (Table 1, Fig. 2) and yield a weighted mean 206 Pb/238U age of 97.044 + 0.045 [0.17] Ma (MSWD ¼ 0.48), which we interpret to best represent the age of crystallization of zircon in this rhyolite unit. Our U – Pb zircon age overlaps the 96.0 + 1.1 Ma sanidine K –Ar age reported by McDougall & van der Lingen (1974; recalculated with decay constants of Steiger & Ja¨ger 1977). Whataroa Granite. Five single zircon CA-TIMS analyses were carried out at MIT on Whataroa Granite sample P52280 (Table 1, Fig. 2). The data exhibit some scatter in age probably due to small but persistent Pb loss. The three oldest analyses (least affected by Pb loss) comprise a coherent cluster with a weighted mean 206Pb/238U age of 88.268 + 0.050 [0.16] Ma (MSWD ¼ 0.73) interpreted as the crystallization age of the pluton. Canavans Quartz Monzonite. Three single zircon crystals analysed by the CA-TIMS method at MIT form a coherent age cluster with a weighted mean 206Pb/238U age of 100.94 + 0.15 [0.28] Ma (MSWD ¼ 2.0). This age best represents the timing of zircon crystallization in the rock.

Dating summary Combined with previously published ages in the arc –break-up gap age range the age results,

Table 2.

40

Ar/39Ar analysis data for Eweburn biotite (P77501)

Heating step

Laser power (W)

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar

Mol 39ArK

Ca/K

% 40ArA

Cum 39ArK

KY4bt2 biotite step 2 heating Mass ¼ 1.623 mg 1 5.00E 2 02 2 1.00E 2 01 3 1.50E 2 01 4 2.00E 2 01 5 2.50E 2 01 3.00E 2 01 6 7 4.00E 2 01 8 5.00E 2 01 9 6.00E 2 01 10 7.00E 2 01 11 8.00E 2 01

1.83E þ 02 3.78E þ 01 3.26E þ 01 3.46E þ 01 3.38E þ 01 3.44E þ 01 3.63E þ 01 3.42E þ 01 3.32E 1 01 3.32E 1 01 3.29E 1 01

J ¼ 3.68E 2 03 + 5.61E 2 06 (0.152%; 2s) 3.67E 2 01 5.06E 2 01 9.54E 2 16 1.92E 2 01 6.44E 2 02 1.48E 2 14 6.63E þ 00 7.04E 2 02 2.94E 2 16 2.66E þ 00 6.77E 2 02 6.14E 2 16 1.55E þ 00 6.30E 2 02 1.01E 2 15 21.19E 2 01 6.01E 2 02 1.58E 2 15 29.76E 2 02 6.56E 2 02 5.29E 2 15 22.22E 2 02 5.72E 2 02 1.64E 2 14 8.56E 2 03 5.34E 2 02 3.33E 2 14 1.70E 2 02 5.37E 2 02 3.77E 2 14 4.93E 2 02 5.22E 2 02 2.67E 2 14

7.21E 2 01 3.77E 2 01 1.31E þ 01 5.22E þ 00 3.04E þ 00 22.33E 2 01 21.91E 2 01 24.35E 2 02 1.68E 2 02 3.33E 2 02 9.67E 2 02

8.17E þ 01 5.03E þ 01 6.12E þ 01 5.67E þ 01 5.45E þ 01 5.17E þ 01 5.35E þ 01 4.95E þ 01 4.75E 1 01 4.78E 1 01 4.69E 1 01

4.39E 2 03 7.25E 2 02 7.38E 2 02 7.67E 2 02 8.13E 2 02 8.86E 2 02 1.13E 2 01 1.88E 2 01 3.42E 2 01 5.15E 2 01 6.38E 2 01

2s

2.03E þ 02 1.16E þ 02 7.97E þ 01 9.73E þ 01 1.01E þ 02 1.11E þ 02 1.12E þ 02 1.12E þ 02 1.21E þ 02 1.22E þ 02 1.12E 1 02 1.13E 1 02 1.13E 1 02 1.13E 1 02 1.12E 1 02 1.15E þ 02 1.15E þ 02 1.11E þ 02 1.13E þ 02 1.27E þ 02

7.91E þ 00 1.45E þ 00 1.30E þ 01 7.92E þ 00 5.96E þ 00 3.47E þ 00 1.32E þ 00 9.78E 2 01 2.48E þ 00 1.74E þ 00 5.24E 2 01 9.25E 2 01 1.14E 1 00 2.42E 1 00 1.82E 1 00 1.33E þ 00 1.15E þ 00 3.30E þ 00 3.16E þ 00 9.21E þ 00

2.10E þ 02 1.21E þ 02 8.25E þ 01 9.71E þ 01 9.93E þ 01 1.07E þ 02 1.09E þ 02 1.11E þ 02 1.12E 1 02 1.12E 1 02 1.12E 1 02

1.39E þ 01 1.98E þ 00 2.06E þ 01 9.19E þ 00 6.10E þ 00 2.62E þ 00 2.50E þ 00 1.09E þ 00 6.47E 2 01 8.37E 2 01 7.30E 2 01 97

(Continued)

CRETACEOUS FELSIC VOLCANISM

KY4bts biotite step-heating Mass ¼ 1.762 mg J ¼ 0.0036787814855 + 0.0000056054010209 (0.152%; 2s) 1 5.00E-02 1.69E þ 02 1.46E-01 4.63E-01 1.66E-15 2.87E-01 8.09E þ 01 7.70E-03 2 1.00E 2 01 3.49E þ 01 7.89E 2 02 5.66E 2 02 2.41E 2 14 1.55E 2 01 4.80E þ 01 1.19E 2 01 3 1.50E 2 01 2.90E þ 01 7.50E 2 01 5.70E 2 02 1.19E 2 15 1.47E þ 00 5.77E þ 01 1.25E 2 01 4 2.00E 2 01 3.23E þ 01 3.13E 2 01 5.86E 2 02 1.82E 2 15 6.13E 2 01 5.34E þ 01 1.33E 2 01 5 2.50E 2 01 3.66E þ 01 1.59E 2 01 7.07E 2 02 2.62E 2 15 3.13E 2 01 5.71E þ 01 1.46E 2 01 6 3.00E 2 01 4.27E þ 01 9.83E 2 03 8.62E 2 02 3.45E 2 15 1.93E 2 02 5.97E þ 01 1.61E 2 01 7 4.00E 2 01 3.42E þ 01 5.87E 2 03 5.72E 2 02 9.85E 2 15 1.15E 2 02 4.94E þ 01 2.07E 2 01 8 5.00E 2 01 3.34E þ 01 1.24E 2 02 5.41E 2 02 2.59E 2 14 2.43E 2 02 4.79E þ 01 3.27E 2 01 9 6.00E 2 01 3.38E þ 01 24.39E 2 03 5.06E 2 02 5.72E 2 15 28.61E 2 03 4.43E þ 01 3.54E 2 01 10 7.00E 2 01 3.38E þ 01 2.86E 2 03 5.01E 2 02 6.06E 2 15 5.60E 2 03 4.38E þ 01 3.82E 2 01 11 8.00E 2 01 3.28E 1 01 4.36E 2 02 5.19E 2 02 3.91E 2 14 8.54E 2 02 4.68E 1 01 5.63E 2 01 12 9.00E 2 01 3.27E 1 01 6.53E 2 02 5.14E 2 02 2.60E 2 14 1.28E 2 01 4.64E 1 01 6.83E 2 01 13 1.00E 1 00 3.26E 1 01 1.10E 2 01 5.12E 2 02 1.46E 2 14 2.16E 2 01 4.64E 1 01 7.51E 2 01 14 1.10E 1 00 3.30E 1 01 1.14E 2 01 5.25E 2 02 1.11E 2 14 2.23E 2 01 4.69E 1 01 8.02E 2 01 15 1.20E 1 00 3.32E 1 01 1.20E 2 01 5.31E 2 02 7.80E 2 15 2.35E 2 01 4.73E 1 01 8.38E 2 01 16 1.50E þ 00 3.31E þ 01 1.26E 2 01 5.13E 2 02 1.34E 2 14 2.46E 2 01 4.57E þ 01 9.00E 2 01 17 3.00E þ 00 3.31E þ 01 9.29E 2 02 5.12E 2 02 1.41E 2 14 1.82E 2 01 4.57E þ 01 9.65E 2 01 18 5.00E þ 00 3.30E þ 01 8.02E 2 02 5.34E 2 02 3.53E 2 15 1.57E 2 01 4.77E þ 01 9.82E 2 01 19 1.00E þ 01 3.41E þ 01 5.95E 2 02 5.56E 2 02 3.07E 2 15 1.17E 2 01 4.82E þ 01 9.96E 2 01 20 1.50E þ 01 3.91E þ 01 3.32E 2 02 6.52E 2 02 9.07E 2 16 6.50E 2 02 4.93E þ 01 1.00E þ 00 Integrated age ¼ 113.97 + 0.40 Ma Plateau age ¼ 112.42 + 0.44 Ma (2s, including J-error of 0.152%) Plateau includes 45.6% of the 39Ar (steps 11 through 15) MSWD ¼ 0.44, probability ¼ 0.78

Age (Ma)

Table 2. Continued Laser power (W)

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar

Mol 39ArK

Ca/K

12 9.00E 2 01 3.27E 1 01 4.74E 2 02 5.11E 2 02 1.94E 2 14 9.29E 2 02 13 1.00E 1 00 3.28E 1 01 8.78E 2 02 5.14E 2 02 1.37E 2 14 1.72E 2 01 14 1.10E 1 00 3.28E 1 01 1.05E 2 01 5.16E 2 02 8.09E 2 15 2.06E 2 01 15 1.20E þ 00 3.27E þ 01 1.08E 2 01 5.07E 2 02 6.50E 2 15 2.12E 2 01 16 1.50E þ 00 3.20E þ 01 7.86E 2 02 4.80E 2 02 9.44E 2 15 1.54E 2 01 17 3.00E þ 00 3.23E þ 01 9.26E 2 02 4.89E 2 02 1.36E 2 14 1.82E 2 01 18 5.00E þ 00 3.21E þ 01 23.70E 2 02 4.84E 2 02 5.19E 2 15 27.26E 2 02 19 1.00E þ 01 3.27E þ 01 23.74E 2 02 5.05E 2 02 2.40E 2 15 27.32E 2 02 20 1.50E þ 01 3.49E þ 01 6.11E 2 03 4.87E 2 02 1.84E 2 16 1.20E 2 02 Integrated age ¼ 113.28 + 0.35 Ma, 2s Plateau age ¼ 112.37 + 0.37 Ma, 2s Plateau includes 64% of the 39Ar (steps 9 through 14) MSWD ¼ 2.0, probability ¼ 0.081

Cum 39ArK

Age (Ma)

2s

4.62E 1 01 4.64E 1 01 4.65E 1 01 4.58E þ 01 4.43E þ 01 4.47E þ 01 4.46E þ 01 4.56E þ 01 4.12E þ 01

7.28E 2 01 7.91E 2 01 8.28E 2 01 8.58E 2 01 9.02E 2 01 9.64E 2 01 9.88E 2 01 9.99E 2 01 1.00E þ 00

1.13E 1 02 1.13E 1 02 1.13E 1 02 1.14E þ 02 1.14E þ 02 1.15E þ 02 1.14E þ 02 1.14E þ 02 1.31E þ 02

7.43E 2 01 1.03E 1 00 1.28E 1 00 1.57E þ 00 1.15E þ 00 1.24E þ 00 1.34E þ 00 2.98E þ 00 3.47E þ 01

2.59E þ 02 1.21E þ 02 9.24E þ 01 9.85E þ 01 1.01E þ 02 1.09E þ 02 1.11E þ 02 1.12E þ 02 1.20E þ 02 1.19E þ 02 1.12E 1 02 1.13E 1 02 1.12E 1 02 1.13E 1 02 1.13E 1 02 1.13E 1 02 1.13E 1 02 1.14E 1 02 1.16E þ 02 1.11E þ 02

9.07E þ 00 2.46E þ 00 1.59E þ 01 7.19E þ 00 5.30E þ 00 3.91E þ 00 1.28E þ 00 9.63E 2 01 2.19E þ 00 1.80E þ 00 1.00E 1 00 1.10E 1 00 1.31E 1 00 1.59E 1 00 2.10E 1 00 1.12E 1 00 1.30E 1 00 3.11E 1 00 2.46E þ 00 1.61E þ 01

KY4bt3 biotite step-heating Mass ¼ 1.789 mg J ¼ 3.68E-03 + 5.61E-06 (0.152%; 2s) 1 5.00E 2 02 2.15E þ 02 9.50E 2 01 5.85E 2 01 2.07E 2 15 1.86E þ 00 8.05E þ 01 9.47E 2 03 2 1.00E 2 01 3.55E þ 01 3.69E 2 01 5.63E 2 02 1.85E 2 14 7.23E 2 01 4.68E þ 01 9.42E 2 02 3 1.50E 2 01 5.11E þ 01 5.40E þ 00 1.27E 2 01 8.84E 2 16 1.06E þ 01 7.22E þ 01 9.82E 2 02 4 2.00E 2 01 3.34E þ 01 3.99E þ 00 6.32E 2 02 1.12E 2 15 7.84E þ 00 5.44E þ 01 1.03E 2 01 5 2.50E 2 01 3.23E þ 01 2.43E þ 00 5.75E 2 02 1.59E 2 15 4.78E þ 00 5.16E þ 01 1.11E 2 01 3.00E 2 01 3.31E þ 01 27.66E 2 02 5.45E 2 02 2.45E 2 15 21.50E 2 01 4.87E þ 01 1.22E 2 01 6 7 4.00E 2 01 3.38E þ 01 21.15E 2 01 5.62E 2 02 8.15E 2 15 22.24E 2 01 4.92E þ 01 1.59E 2 01 8 5.00E 2 01 3.32E þ 01 29.50E 2 03 5.32E 2 02 2.38E 2 14 21.86E 2 02 4.73E þ 01 2.68E 2 01 9 6.00E 2 01 3.38E þ 01 21.84E 2 01 5.10E 2 02 4.43E 2 15 23.61E 2 01 4.47E þ 01 2.88E 2 01 10 7.00E 2 01 3.37E þ 01 22.98E 2 02 5.15E 2 02 5.29E 2 15 25.85E 2 02 4.52E þ 01 3.12E 2 01 11 8.00E 2 01 3.28E 1 01 3.66E 2 02 5.20E 2 02 3.95E 2 14 7.17E 2 02 4.68E 1 01 4.92E 2 01 12 9.00E 2 01 3.27E 1 01 9.85E 2 03 5.12E 2 02 2.97E 2 14 1.93E 2 02 4.63E 1 01 6.28E 2 01 13 1.00E 1 00 3.28E 1 01 3.13E 2 02 5.19E 2 02 1.93E 2 14 6.13E 2 02 4.67E 1 01 7.16E 2 01 14 1.10E 1 00 3.31E 1 01 21.31E 2 02 5.23E 2 02 1.28E 2 14 22.58E 2 02 4.68E 1 01 7.75E 2 01 15 1.20E 1 00 3.30E 1 01 21.47E 2 02 5.20E 2 02 9.28E 2 15 22.89E 2 02 4.66E 1 01 8.17E 2 01 16 1.50E 1 00 3.28E 1 01 5.22E 2 02 5.12E 2 02 1.40E 2 14 1.02E 2 01 4.62E 1 01 8.81E 2 01 17 3.00E 1 00 3.26E 1 01 8.83E 2 02 5.07E 2 02 1.84E 2 14 1.73E 2 01 4.60E 1 01 9.65E 2 01 18 5.00E 1 00 3.39E 1 01 21.68E 2 01 5.49E 2 02 2.71E 2 15 23.29E 2 01 4.79E 1 01 9.77E 2 01 19 1.00E þ 01 3.51E þ 01 21.96E 2 01 5.77E 2 02 4.42E 2 15 23.84E 2 01 4.87E þ 01 9.97E 2 01 20 1.50E þ 01 3.72E þ 01 27.56E 2 01 6.71E 2 02 5.70E 2 16 21.48E þ 00 5.36E þ 01 1.00E þ 00 Integrated age ¼ 114.91 + 0.46 Ma Plateau includes 66.5% of the 39Ar (steps 11 through 18) Plateau age ¼ 112.88 + 0.50 Ma (2s, including J-error of .152%) MSWD ¼ 0.45, probability ¼ 0.87 Measured ratios are corrected for blanks, backgrounds, and radioactive decay. Reactor-produced corrections are 36Ar/39ArCa ¼ 0.445 + 0.005, 37Ar/39ArCa ¼ 1005.8 + 7.2, and 40Ar/39ArK ¼ 0.004 + 0.002. Ages of individual steps are calculated assuming an initial 40Ar/36Ar ¼ 295.5 and do not include error in the irradiation parameter J.

A. J. TULLOCH ET AL.

% 40ArA

98

Heating step

CRETACEOUS FELSIC VOLCANISM

99

Fig. 3. Isotopic ages of Shag Valley Ignimbrite and Eweburn Tuff, North Otago. (a) LA-ICP-MS U –Pb age plots of Shag Valley Ignimbrite 06NZ-11, showing all data. Removal of the two youngest ages produces a Gaussian peak with an age of 112.3 + 0.4 Ma (2s, MSWD 0.83). (b) 40Ar/39Ar age spectra of Eweburn Tuff biotite step-heating analyses. Apparent age boxes are +2s. Solid symbols are heating steps used to calculate plateau age. Blank symbols are analyses excluded from plateau ages. Graphics generated using Isoplot 3.0 (Ludwig 2003).

summarized in Table 4 and Figure 4, define four distinct groups in time and space.

post-127 Ma rhyolites (Allibone & Tulloch 2004) on NE Stewart Island are possible correlatives.

112 Ma group. The Eweburn Tuff and Shag Valley Ignimbrite have essentially the same age (weighted average ¼ 112.5 + 0.2 Ma, MSWD ¼ 0.8) suggesting the possibility that they are products of the same eruption. This is the only rhyolitic volcanism of this age known in Zealandia. It occurs on the Otago Schist accretionary wedge core of the Eastern Province, although plutons of this age are common in the Median Batholith (Tulloch & Kimbrough 2003).

97 Ma group. The DSDP 207A rhyolite can be grouped with rhyolites from the Mount Somers Volcanic Group (which range from 97.0 + 1.5 Ma to 98.0 + 1.2 Ma; Tappenden 2003) and c. 97 Ma Takahe seamount granite (Mortimer et al. 2006) to define a tight age group with mean age calculated at 97.04 + 0.05 Ma (MSWD ¼ 1.2). Coeval alkaline mafic magmatism includes that of the Tapuaenuku (96.2 + 0.8 Ma; Baker & Seward 1995) and Mandamus Igneous (96.2 + 1.7 Ma; Tulloch 1991; Weaver & Pankhurst 1991; Tappenden 2003) complexes. The 97 Ma group has a similar Zealandiawide spatial distribution to the 101 Ma group.

101 Ma group. The Houhora ignimbrite, Motu Tuff, Stitts Tuff and Canavans quartz monzonite form a distinct group with a mean age of 100.5 Ma. This c. 101 Ma group erupted across all of Early Cretaceous Zealandia from the tip of the accretionary prism (Houhora, Motu) to the west of the Median Batholith (Stitts) and deep in the Gondwana interior (Canavans). In detail, however, ages vary from 100.264 + 0.063 Ma in the west (Stitts Tuff) to 101.6 + 0.2 Ma (Motu Tuff), and c. 101.9 Ma (Houhora ignimbrite), in the east. No examples are known from the Median Batholith, although

85 Ma group. The 88 Ma Whataroa Granite can be grouped with a generally slightly younger set of siliceous igneous rocks such as French Creek Granite (81 Ma) and Hohonu peralkaline rhyolites and intense Buller– Hohonu lamprophyre dyke swarms (83 + 3 Ma; Adams & Nathan 1978) in a c. 82 Ma group (Fig. 4). As these rocks have potential correlatives in 85– 82 Ma alkalic basaltic

100

Table 3. U –Pb analyses (laser ablation ICP-MS) of zircons from the Shag River ignimbrite (P77502) Radiogenic ratios Analysis

Pb/

206

Pb

0.0498 0.04895 0.04754 0.04899 0.04971 0.05086 0.04875 0.04931 0.04881 0.04578 0.04831 0.0484 0.04797 0.0489 0.04858 0.04809

Error 0.00181 0.00154 0.00097 0.0008 0.00162 0.00078 0.0007 0.00132 0.00091 0.00103 0.00094 0.00134 0.00082 0.00074 0.00063 0.00139

207

235

Pb/

U

0.12092 0.11838 0.11496 0.11873 0.11953 0.12374 0.11872 0.11978 0.11853 0.11278 0.11779 0.11631 0.11803 0.12096 0.12071 0.11915

Error

Ages (Ma) 206

0.00436 0.00368 0.00233 0.00192 0.00386 0.00188 0.00169 0.00318 0.00219 0.00252 0.00227 0.0032 0.002 0.00182 0.00157 0.00342

Pb/

238

U

0.01706 0.01722 0.01743 0.01745 0.01744 0.01748 0.01751 0.01754 0.01756 0.01756 0.01758 0.01761 0.01767 0.01768 0.01772 0.01775

Error 0.00021 0.00017 0.00012 0.0001 0.0002 0.0001 0.00009 0.00016 0.00012 0.00013 0.00012 0.00017 0.00011 0.0001 0.00008 0.00018

207

Pb/

206

Pb

185.8 145.4 75.5 147.2 181.2 234.7 135.9 162.5 138.9 0.1 114.5 118.8 96.5 143.1 127.6 103.6

Error 82.6 72.0 48.5 37.9 74.3 35.0 33.3 61.5 43.2 39.3 45.3 64.1 41.1 35.2 30.5 67.0

207

Pb/235U 115.9 113.6 110.5 113.9 114.6 118.5 113.9 114.9 113.7 108.5 113.1 111.7 113.3 115.9 115.7 114.3

Error

206

Pb/238U

4.0 3.3 2.1 1.8 3.5 1.7 1.5 2.9 2.0 2.3 2.1 2.9 1.8 1.7 1.4 3.1

Error

109.1 110.1 111.4 111.5 111.5 111.7 111.9 112.1 112.2 112.2 112.3 112.5 112.9 113 113.2 113.4

1.4 1.1 0.8 0.6 1.2 0.6 0.6 1.0 0.7 0.8 0.8 1.1 0.7 0.6 0.5 1.1

Individual errors are 1s; calibration standard is Plesovic zircon (Slama et al. 2007).

Table 4. Summary of new age results, and sample locations Unit Eweburn Tuff Shag Valley Ignimbrite Houhora ignimbrite Whatuwhiwhi granite clast Motu Tuff‡ Stitts Tuff DSDP207A rhyolite Whataroa Granite Canavans qtz monz.

Setting

Sample

Grid ref*

Method

Age

Non-marine half angle graben, Tr – J basement Non-marine half angle graben, Tr – J basement Felsic-basaltic andesite association Felsic-basaltic andesite association Immed above Motuan unconform. Non-marine half angle graben in Pz –Mz basement Lord Howe Rise Intrudes Palaeozoic basement Intrudes Palaeozoic basement

P77501 P77502 P52296 P52300 P63724 P49125 P63855 P52280 P45553

H43/833 767 J43/336 267 N03/269 107 O03/486 021 X16/135 178 K29/098 280 36.9625S, 165.4343E I35/962 662 H35/799 550

Ar –Ar bt U –Pb ICPMS zrn U –Pb TIMS zrn U –Pb TIMS zrn U –Pb TIMS zrn U –Pb TIMS zrn U –Pb TIMS zrn U –Pb TIMS zrn U –Pb TIMS zrn

112.5 112.3 c. 101.9 c. 101.9 101.6 100.264 97.044 88.268 100.94

*All NZMS 260 map references, except for P63855. More details on samples at http://pet.gns.cri.nz/ † TIMS errors do not include decay constant errors, which are reported in the text (Results). ‡ Age published in Crampton et al. (2004).

Age err (2s)† 0.2 0.4 0.2 0.063 0.045 0.050 0.15

A. J. TULLOCH ET AL.

P i m b o d a j l n e c f h g k

207

CRETACEOUS FELSIC VOLCANISM

101

volcanism on the Chatham Islands they too have a Zealandia-wide spatial distribution. At the time of eruption and intrusion, the Median Batholith arc to the east would have been extinct for some 20 Ma, and Tasman Sea spreading would just have been commencing to the west.

Geochemistry Whole rock chemical analyses (X-ray fluorescence majors, Ga and transition metal trace elements, ICP-MS rare-earth elements and other trace elements) are presented in Table 5, and tracer isotopes in Table 6. The rocks are a mix of intrusions, lavas, ignimbrites and airfall tuffs. The bulk compositions of the latter are particularly prone to depart from primary igneous values and all rocks, because of their antiquity, show some secondary alteration.

Compositional variation of age groups

Fig. 4. Age relationships of Cretaceous mainly felsic volcanic and plutonic rocks. Error bars are two sigma. Previously dated units are presented as grey bars. (a) Ages plotted normal to Cretaceous orogenic margin; towards palaeotrench on right side of diagram, towards continental interior on left side of diagram. (b) Relative probability plot for data from (a). Note that peak heights on probability density curve are driven by age precision more than by number of samples. Histogram bins are 2 Ma. Abbreviations: DSDP, DSDP 207A; Ewe, Eweburn Tuff; Hou, Houhora Complex; Sti, Stitts Tuff; Wha, Whataroa Granite; Can, Canavans Knob; Shg, Shag. Previously dated rocks: Fcg, French Creek Granite (Waight et al. 1997); Bla, Hohonu lamprophyres (Waight et al. 1997); BPt, Buttress Point (Phillips et al. 2005); Man, Mandamus (Tulloch 1991; Weaver & Pankhurst 1991); Mso, Mount Somers (Tappenden et al. 2002); Mot, Motu Tuff (Crampton et al. 2004); Stitts Tuff (Muir et al. 1997); Tak, Takahe (Mortimer et al. 2006); Tap, Tapuaenuku (Baker & Seward 1996); Wib, Wishbone (Mortimer et al. 2006). Age of 81.2 + 1.3 Ma for French Creek Granite is weighted average of SIMS age (81.7 +1.8 Ma; Waight et al. 1997) and K– Ar hornblende age (P52260, 80.5 + 2 Ma; Nathan et al. 2000). Age for Buller lamprophyres is average of Adams & Nathan (1978) K– Ar hornblende ages, recalculated to current decay constants.

112 Ma group. Eweburn (Na . K) and Shag (K . Na) analyses indicate subalkaline dacitic and rhyolitic compositions respectively (Fig. 5), although anomalously low alkalis and highly aluminous compositions suggest either incorporation of sedimentary material, or alteration. However, the extreme high Sr and Ba content of some samples is likely to be primary as microprobe analyses of fresh feldspar phenocrysts from Shag Valley Ignimbrite (Hill 1999) also show high SrO contents (e.g. OU65482; SrO plagioclase ¼ 0.35% and K –feldspar ¼ 0.58 wt% SrO). Alkali-lime indices are anomalously calcic, suggesting that the igneous composition of these samples have been modified by loss of alkalis. A multi-element plot (Fig. 6a) for the Shag Valley Ignimbrite shows a pattern similar to the subduction-related I-type pre-130 Ma Darran Suite of the Median Batholith. However, the Sr-Nd isotopic composition (see below) for this rock overlaps that of the underlying Otago Schist– Torlesse accretionary prism of the Eastern Province through which the Eweburn and Shag magmas must have passed. This suggests that its subduction characteristics have been inherited from these metasedimentary rocks, themselves derived from subduction related igneous rocks, elsewhere on the Gondwana margin (e.g. Wandres et al. 2004; Adams et al. 2007). The Eweburn Tuff is characterized by extremely high Sr and high Ba. The Sr, Ba anomalies might suggest partial melting of an overthickened accretionary prism at depths at which feldspars were not stable. However, such conditions would also favour garnet growth and consequent Y-depletion in partial melts, which is not observed. The Shag-Eweburn rocks differ from coeval igneous rocks in the Median Batholith in having

102

A. J. TULLOCH ET AL.

Table 5. XRF and ICPMS analyses of mid Cretaceous felsic igneous rocks from Zealandia Sample OU65482 Pluton loc Eweburn Rock type Ignimbrite Age (Ma) SiO2 TiO2 Al2O3 Fe2O3T MnO MgO CaO Na2O K2O P2O5 LOI Total Ba Rb Sr Y Zr Nb Th Pb Ga Zn Cu Ni V Cr Hf Cs Sc Ta U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu T, Zrn-Sat

R7583 P22174 P77502 OU50106 P50286 Eweburn Shag Shag River Canavans Camel Tuff Ignimbrite Ignimbrite Qtz monz Andesite 112

68.89 0.4 13.33 1.86 0.02 0.63 3.58 2.59 1.15 0.13 7.6 100.18 2187 149 4571 29 162 16 21 26 16 48 5 6 33 4 – – – – 3 27.2 57.8 – 29.2 5.21 0.94 – 0.84 – 1.14 – – 2.81 0.37 800

112

101 60.36 0.85 14.76 6.02 0.08 4.82 5.33 2.81 3.23 0.23 0.78 99.19

64.57 0.68 16.98 3.47 0.05 0.89 3.87 4.11 1.86 0.17 3.03 99.69

70.17 0.33 13.53 1.34 0.01 0.66 1.09 1.45 4.84 0.02 6.17 99.61

69.1 0.34 14.3 1.34 0 0.66 1.5 1.7 3.87 0.1 6.06 99.59

1575 100.3 2197 30 272 11.09 13.98 17.31 26 73 0 3 65 4 7.24 8.02 9.6 0.78 5.43

3729 207.6 401 16 148 9.95 24.16 10.33 15 53 5 6 20 1 4.59 10.42 3.6 1.12 2.66

739 163 318 13.7 218 10.3 23.9 8 18.9 38 0 0 18 0 5.9 – – 1.2 3.44

42.48 84.32 10.28 37.76 7.43 1.79 5.99 0.96 5.68 1.12 2.89 0.4 2.37 0.36

40.52 77.66 8.78 30.3 5.23 1 3.92 0.56 3.26 0.6 1.58 0.23 1.43 0.23

34.5 71.7 7.83 28.6 4.51 0.82 3.63 0.5 2.59 0.5 1.41 – 1.26 0.18

831

814

853

variable, but generally much higher, Sr contents, especially at such high SiO2 levels (in itself rare in Median Batholith), and much more enriched Sr– Nd isotopic compositions. 101 Ma group. Houhora Complex samples range from basaltic andesite and trachyandesite

P50288 P52296 P52298 P52297 Camel Houhora Camel Camel Basaltic Ignimbrite Ignimbrite Grt clast Andesite 102

56.19 2.1 14.22 11.34 0.18 2.55 3.37 5.28 2.01 0.5 2.34 100.08

58.78 1.96 13.83 10.18 0.17 2.41 5.19 3.34 1.63 0.48 2.41 100.38

702 138 558 29 182 12 15 22 20 69 26 92 121 211 – – 14 – 3

397 51 191 53 258 17 9 13 23 93 23 4 209 4 – – 27 – 2

276 37 83 45 277 22 8 16 20 89 21 2 204 9 – – 19 – 4

43 80 – – – – – – – – – – – –

27 67 – – – – – – – – – – – –

10 81 – – – – – – – – – – – –

749

772

785

77.05 0.2 11.85 1.54 0.01 0.48 0.11 3.62 3.84 0.02 1.24 99.96

77.72 0.17 11.82 1.8 0.01 0.25 0.36 4.73 2.44 0.02 0.89 100.21

76.47 0.26 13.26 0.29 0.01 0.07 0.09 5.01 3.82 0.03 0.73 100.04

319 160 50 64 487 20.28 19.18 34.2 20 85 11 2 7 0 12.98 0.99 3 1.67 4.42

277 105 288 62 539 23 19 28 20 114 8 0 6 0 – – 0 – 5

492 96 58 50 461 25 16 17 20 56 5 0 7 0 – – 2 – 4

49.09 102.47 12.39 46.9 10.19 0.7 10.05 1.8 11.59 2.43 6.71 1 6.19 0.95

45 92 – – – – – – – – – – – –

40 84 – – – – – – – – – – – –

919

898

917

(metaluminous) to high silica rhyolite (peraluminous), with marginally alkaline compositions (data of Law 1983; Palmer et al. 1995; Mortimer et al. 1998; Nicholson & Black 2004; this study). The suite is medium K with Na . K. Alkali-lime indices scatter widely from mildly calcic to alkalic. The suite is bimodal with a c. 66– 76% SiO2 gap

CRETACEOUS FELSIC VOLCANISM

103

Table 5. Continued P52300 Camel Grt clast

P63723 Motu Tuff

102 76.93 0.25 11.7 1.79 0.01 0.25 0.37 4.68 2.41 0.02 0.89 99.3

P63724 Motu Tuff

OU65479 Stitts Tuff

102 68.00 0.27 14.2 1.28 0.03 0.58 2.93 0.45 1.97 0.03 7.17 96.91

71.14 0.26 12.98 0.9 0.01 0.54 2.73 0.37 1.89 0.04 6.49 97.35

74.26 0.23 13.5 1.98 0.07 0.43 1.29 0.69 3.98 0.04 3.73 100.20

P49125 Stitts Tuff

P63855 DSDP207 Rhyolite

P52280 Whataroa A type grt

P52260 Hohonu Ra A type grt

100

97

88

81

77.96 0.17 12.58 0.88 0 0.41 0.15 1.19 3.56 0.06 2.86 99.82

73.86 0.15 11.8 1.83 0.02 0.07 0.56 3.2 4.79 0 3.65 99.93

68.04 0.56 15.03 4.33 0.07 0.74 2.3 3.59 4.59 0.19 0.88 100.32

73.79 0.19 13.33 2.07 0.03 0.16 0.49 3.77 5.36 0.03 0.59 99.81

Median Batholith Stewart Island Average .65% SiO2 HiSY

LoSY

72.00 0.21 15.16 1.58 0.02 0.46 1.94 4.31 3.46 0.06 0.53 99.72

71.65 0.30 14.59 2.13 0.03 0.59 1.57 4.17 3.89 0.08 0.78 99.78

629 87.6 66 47 511 30.96 23.79 15 23 24 5 0 7 1 15.08 1.67 2.6 2.31 3.6

8739 33 8756 19 7 15 18 24 15 34 11 0 15 0 – – 5 – 10

8206 50.6 9888 63 263 15.06 17.91 19.1 12 21 5 2 0 5 7.56 2.38 4 1.14 4.19

461 239 71 43 – 17 20 29 19 43 5 8 25 19 – – – – 7

286 215 22 28 136 15 18 15 17 28 2 3 12 4 – – – – 6

237 216 29 68 299 28.92 20.91 32.03 23 149 0 0 0 0 9.76 10.78 2.9 2.14 5.4

923 99 263 41 391 32 12 15 21 70 5 3 28 2 – – 6 – 2

201 210 31 57 320 79 22 21 25 56 1 1 5 0 – – 2 – 5

1169 75 599 6 116 4.61 5.99 24.48 15 27 9 4 22 8 – – 3.07 – 1.57

750 103 294 13 157 8.35 14.87 19.56 16 31 7 5 33 6 – – 5.27 – 2.2

43.77 93.72 11.25 42.58 9.45 1.19 8.3 1.43 8.72 1.75 5.04 0.77 4.99 0.77

15 5 – – – – – – – – – – – –

42.99 79.45 9.68 33.71 6.08 1.03 4.41 0.69 4.03 0.78 2.1 0.33 2.17 0.34

34 64 – – – – – – – – – – – –

31 57 – – – – – – – – – – – –

64.27 132.56 16.12 62.47 13.77 0.78 12.57 2.15 13.06 2.61 7.14 1.04 6.38 0.96

47 90 – – – – – – – – – – – –

72 133 – – – – – – – – – – – –

18.18 37.61 3.24 12.07 2.17 0.63 1.48 0.19 1 0.19 0.5 0.07 0.52 0.09

21.42 60.1 6.12 24.04 5.18 0.94 3.99 0.59 3.33 0.64 1.69 0.24 1.55 0.24

860

855

882

888

835

852

762

787

XRF analyses by Spectrachem (Lower Hutt), except P22174, P49125, P52296, 52297, 52298, 52300 (VUW) P77502 (ALS Chemex l, Vancouver), OU65482, OU65479 (Otago University). REE and (for those samples which have REE) and, Ba, Th, Nb, Y, Hf, Ta, U, Pb, Rb, Cs, Sr, Sc and Zr by ICPMS (GeoAnalytical Lab, Washington State University). Zrn–Sat, zircon saturation temperature (8C) —, not analysed; 0, below detection.

between trachydacite and high silica rhyolite. Primitive mantle-normalized multielement plots show small negative Nb anomalies, and strong positive Zr, Y anomalies compared to average low Sr/Y arc magma of the Median Batholith. Other members of this group (except for the Motu Tuff) show similar, but weaker patterns. Canavans Quartz

Monzonite is distinctly high-K, perhaps expected due to its inboard Gondwana position intruding relatively potassic Palaeozoic greywackes. 97 Ma group. Subalkaline felsic igneous rocks from the 97 Ma group include DSDP 207A, Mount Somers rhyolites and the Takahe granite. Both

104

Table 6. Sr, Nd isotopic data for mid Cretaceous igneous rocks Unit

P22174 R7583* P49125 P63724* P52296 P52300 P63855

Shag ignimbrite Eweburn tuff Stitts Tuff Motu Tuff Camel ignimbrite Camel grt clast DSDP 207 rhyolite

Sr Rb (ppm) (ppm) 201.0 100.3 129.8 50.6 159.1 87.3 214.3

377.9 2197 300.7 9888 47.71 82.0 18.06

87

Rb/86Sr

1.539 0.132 1.249 0.0148 9.661 3.079 34.49

87

Sr/86Sr

0.70873 0.70727 0.71261 0.70686 0.71731 0.71158 0.75432

Sm Nd (ppm) (ppm) 4.76 6.74 11.27 5.16 9.73 9.01 12.46

28.93 36.34 68.97 30.51 46.50 42.52 59.20

147

Sm/144Nd 0.0992 0.1120 0.0987 0.1021 0.1263 0.1280 0.1271

143

Nd/144Nd

0.512470 0.512423 0.512311 0.512644 0.512761 0.512760 0.512565

1Nd now 23.28 24.19 26.38 0.12 2.40 2.38 21.42

TDM Geol. age (Ga) (Ga) 0.91 1.09 1.11 0.69 0.68 0.70 1.03

*Samples were not spiked for Rb– Sr; 87Rb/86Sr calculated from ICPMS trace element data. 87 Sr/86Sr internal precision (2s)  +0.00002, external precision (2sd) +0.00004. External precision for 87Rb/86Sr +0.5%. All 87Sr/86Sr adjusted to SRM987 ¼ 0.71023, after normalization to 88Sr/86Sr ¼ 8.37521. 143 Nd/144Nd internal precision (2s)  +0.000010, external precision (2sd) +0.000020. External precision for 147Sm/144Nd 0.2%. All 143Nd/144Nd adjusted to La Jolla ¼ 0.511860 after normalization to 146Nd/145Nd ¼ 2.07194253 (equivalent to the more familiar 146Nd/144Nd ¼ 0.7219). CHUR is 0.1967 and 0.512638. TDM relative to a present-day depleted mantle model with 0.2136 and 0.513141. l87Rb ¼ 1.42E 2 11/yr; l147Sr ¼ 6.54E 2 12/yr. Analytical procedures at the University of Melbourne are detailed in Maas et al. (2005).

0.113 0.112 0.102 0.102 0.102 0.102 0.097

Initial Sri

Initial 1Ndi

0.70626 0.70705 0.71080 0.70684 0.70331 0.70712 0.70678

21.87 22.98 25.10 1.35 3.32 3.27 20.56

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Fig. 5. Major element plots for felsic mid Cretaceous rocks from New Zealand (a) Shag, Canavans, Stitts, French Creek, Whataroa and Mount Somers are all high potassium. (b) Silica v. total alkalis plot. Most samples have moderate to high alkali contents, except for Eweburn, Stitts and Motu tuffs whose igneous compositions have modified by sediment dilution and/or other post-eruptive processes. The Houhora samples show a bimodal distribution between dominant trachyandesite and high-silica rhyolite.

Fig. 6. (a) Primitive mantle normalised multi-element plots of Cretaceous felsic (.68% SiO2) igneous rocks show a range from subduction related Median Batholith to intraplate French Creek Granite. Eweburn Tuff and Shag Valley Ignimbrite (112 Ma) are similar to subduction-related LoSY Median Batholith (average . 65% SiO2) except for their very high Sr and Ba contents. Compared to the Median Batholith, Houhora, Stitts and DSDP207 exhibit shallower negative Nb anomalies, higher large ion lithophile elements, deeper Ba, Ti and Sr (apart from one Motu sample) negative anomalies, and greater positive Y and Zr anomalies. French Creek granite exhibits a small positive normalised Nb anomally. (b) Chondrite-normalised rare-earth element (REE) patterns. Shag Valley Ignimbrite (112 Ma) and Motu Tuff (102 Ma) show similar, moderately negative, Eu anomally patterns to the Median Batholith LoSY average (.65% SiO2; grey band), and also have moderate light REE enrichment. The younger rocks, Houhora, DSDP 207A and French Creek Granite have large negative Eu anomalies, flatter heavy REE patterns and an overall flatter light–heavy REE pattern. One Eweburn Tuff has only a very slight negative Eu anomaly, suggesting that its very high Sr content is associated with little involvement of plagioclase during generation or fractionation. Symbols same as for Fig. 5.

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DSDP 207A and Mount Somers are high-K and have transitional calc-alkalic to alkali-calcic alkalilime indices. Multielement plots show all have negative Nb-anomalies, with DSDP 207A being the shallowest. This age group also includes mostly mafic rocks of the Tapuaenuku and Mandamus Igneous Complexes, both of which have ocean island basalt (OIB) like compositions. 85 Ma group. The Whataroa and French Creek Granite (and associated rhyolites) are both high-K granites with alkalic alkali-lime indices. On multielement plots French Creek Granite has a slight positive Nb-anomaly; strong negative Ba and Sr negative anomalies and positive Rb anomalies suggest strong fractionation, as previously described by Waight et al. (1998a). Whataroa has a shallow negative Nb-anomaly, and an overall smoother pattern. REE plots of selected samples reveal two groups. (1) An older Shag, Eweburn and Motu group with relatively steep patterns and modest heavy REE depletion. This group is similar to the patterns of the subduction-related low Sr/Y belt of the Median Batholith but are much younger. Shallow negative Eu-anomalies suggest minor plagioclase fractionation. (2) A younger Houhora, French Creek and DSDP 207A group has slightly higher light REE than the older group, but a distinctly higher and flatter heavy REE pattern, and large negative Eu-anomalies suggestive of significant plagioclase fractionation.

Discrimination plots: tectonic setting Various trace element discrimination plots (Fig. 7; Pearce et al. 1984; Whalen et al. 1987) indicate A-type characteristics for Houhora, DSDP 207A, Whataroa and French Creek siliceous rocks, whereas Eweburn, Shag Valley and Canavans rocks plot in the Volcanic Arc and I/S fields, along with Median Batholith averages. Motu and Stitts Tuffs overlap these boundaries, likely due to sedimentary contamination. Mount Somers rhyolites also overlap but in this case likely due to source contamination (Barley 1987). Further subdivision of the A-type felsic igneous rocks is presented in Figure 7c, d. On the Nb– Y –Ce plot of Eby (1992; Fig. 7d), only the French Creek granites and rhyolites fall clearly within the intraplate A1 A-type group. This group represents mantle differentiates from OIB, intraplate and rift zone sources. On the Ce/Nb v. Y/Nb plot of Eby (1992) the French Creek granites and rhyolites fall in the OIB field, with the remainder of the New Zealand A-types defining a linear trend from the French Creek samples towards the Island Arc Basalt field. Canavans, Shag, Eweburn and Mt

Somers, along with Median Batholith, appear to form a linear array which diverges from the main trend which defines the spectrum of compositions with varying degrees of A-type character.

Crust and mantle contributions to magmas An indication of juvenile mantle v. crustal contributions may p beffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi represented ffi by the parameter Nb/ Nb* ¼ NbN/ (ThN LaN ) which provides a potential index of crustal contamination including that from arc-related magmas and derived sediments (Eisele et al. 2002). Figure 8a shows the temporal variation of juvenile mantle input to magmatism from 160 to 80 Ma in Zealandia, as indicated by excursions above Nb/Nb* values of 0.2–0.3, the crustal values represented by Early Cretaceous Median Batholith arc magmas and Early Palaeozoic Western Province metasedimentary rocks. Other than the 145 Ma spike of Stewart Island dykes, significant juvenile components first appear at 101 Ma and 97 Ma, and Nb/Nb* continues to increase though 87 Ma (Whataroa) to the French Creek granites at 83 Ma seafloor spreading and Gondwana break-up. Figure 8b shows that a plot of 87 Sr/86SrT v. Nb/Nb* displays an apparent positive trend, (the opposite of that reported by Eisele et al. (2002) for generation of Pitcairn EM1 by mixing of DM and mixtures of 0.9–1.7 Ga continental crust and pelagic sediment). This trend appears to represent some sort of control on the maximum proportion of crustal contamination, but it is unclear how this works. Perhaps the degree of assimilation is constant and the lower bound is controlled by series of mixing lines from Western Province host rocks to mafic dykes of varying Nb/Nb*. An indication of the degree of crustal involvement in the various magmatic rocks is also indicated by 87Sr/86SrT v. 1Nd variation (Fig. 9). The 112 Ma group (Shag Valley and Eweburn) lies within the field defined by the composition of its host Eastern Province (Rakaia Terrane) sandstones (Adams et al. 2005), and thus was likely derived largely from partial melting of the Eastern Province accretionary prism. The 101 Ma group shows a wide range in crustal components (Stitts . Motu . Houhora). The Houhora ignimbrite has a relatively primitive isotopic composition (and high Nb/Nb*) indicating only minor crustal contribution. This is apparently at odds with the abundant and diverse zircon inheritance, something that is worthy of further study. The two felsic samples of the 97 Ma group (DSDP 207A and Takahe rhyolites) have 1Nd close to zero, suggesting only modest crustal contamination and absence of major fractionation (as is the case with Mount Somers; Tappenden 2003). They do, however, have much more enriched

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Fig. 7. (a–b) Discrimination of A-type granites from I- and S-type granites (Whalen et al. 1987; Pearce et al. 1984). NZ samples Houhora, DSDP, Whataroa and French Creek lie entirely in the Within Plate field on all plots, whereas Eweburn, Shag and Canavans plot in the Volcanic Arc fields. Motu and Stitts over lap these boundaries, due to sedimentary contamination at source or during deposition; Mt Somers also overlaps, due to source contamination. (c) On the Ce/Nb vs Y/Nb plot of Eby (1992) the French Creek granites and rhyolites fall in the OIB field, with the remainder of the NZ A-types defining a linear trend from the French Creek samples towards the IAB field. The I-type granites (Median Batholith, Canavans) and Shag Ignimbrite form a minor, but distinct, discordant horizontal trend, suggestive of Y-depletion due to residual garnet. (d) On the Nb–Y– Ce plot of Eby (1992), only the French Creek granites and rhyolites fall clearly within the intraplate A1 A-type group. Supracrustal contaminants: Greenland Group (þ, average composition from Roser et al. 1996), Rakaia Terrane sandstone (X, average composition from Roser & Korsch 1999).

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crustal assimilation and major fractionation as discussed by Waight et al. (1998a).

Discussion Age, chemistry and tectonics summary The 20 Ma interval within which our newly dated Zealandia rocks lie, 112–82 Ma, represents the time in which long-lived subduction ceased beneath Zealandia, and was replaced by a regime first of widespread continental rifting and stretching, and finally, at 83 Ma by sea floor spreading on two axes in the Tasman Sea and Southern Ocean. The widespread geographic distribution of our samples, combined with their high precision dating, enables for the first time, an interpretation of age progression trends of 112–82 Ma magmatism across the entire width and half the length of Zealandia.

Silicic magmatism within the schist accretionary wedge

Fig. 8. (a) Juvenile vs crustal contributions as ffiindicated pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi by the parameter Nb/Nb* ¼ NbN/ (ThN LaN ) (normalised to primitive mantle) which provides a potential index of crustal contamination (continental crust, including arc-related magmas, and derived sediments). Temporal variation of juvenile mantle input to magmatism during the arc-breakup period, as indicated by excursions above Nb/Nb* values of c. 0.2– 0.3, the values for Early Cretaceous arc magmas and Early Paleozoic metasedimentary rocks (Greenland Group). A juvenile component first appears at 101 Ma and 97 Ma, and continues to increase through 87 Ma to the French Creek granites at c. 83 Ma breakup. (b) Plot of Nb/Nb* vs 87Sr/86SrT displays an apparent positive trend, (the opposite of that reported by Eisele et al. (2002) for generation of Pitcairn EM1 by mixing of DM and mixtures of 0.9– 1.7 Ga continental crust and pelagic sediment). This trend appears to represent some sort of control on the maximum proportion of crustal contamination.

isotopic compositions than the mafic members (Mandamus, Tapuaenuku) of this age group and the more mafic Houhora rocks (Mortimer et al. 1998). The French Creek granites and rhyolites are less primitive than the associated mafic lamprophyres and are pulled to the right of the mantle array by various combinations of very minor

The 112 Ma group (Shag Valley and Eweburn) is coeval with the Median Batholith (230 –105 Ma) arc magmatism. Although there is complex local fault control, the two sites lie 50 km apart and tentatively define a trend subparallel to the then-active late Early Cretaceous arc (Median Batholith) and trench. Although the tuffs occur in association with a major Cretaceous normal fault system (Waihemo Fault), taken at face value they have arc rather than A-type characteristics. The 112 Ma group rocks represent the only known occurrence of this age of rhyolitic volcanism in Zealandia. By comparison with clast sizes in modern day Taupo Volcanic Zone ignimbrites (Walker & Wilson 1983), a vent within 30 –40 of km of the Waihemo Fault is suggested, thus ruling out the Median Batholith as a source. Why are they there? If there was still a slab beneath the Eastern Province at this time (a fairly likely proposition) then this rare accretionary wedge magmatism might represent subduction of a spreading ridge and/or hotspot trace, which caused shallow melting of the greywackes to produce the subalkaline rhyolites with no emergent mafic magmatism. Ridge subduction has been proposed as the primary reason for the change from convergent to extensional tectonics between 105 and 83 Ma in Zealandia (Bradshaw 1989). Ridge subduction models have also been proposed for similar tectonic changes in the NE Pacific (Sisson & Pavlis 1993; Sisson et al. 2003). Slab rollback, steepening and/or break-off might represent alternative models (e.g. Tappenden et al. 2002; Tappenden 2003, Gray & Foster 2004) that would also be consistent with the normal fault setting.

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Fig. 9. Crustal involvement as indicated by 87Sr/86SrT v. 1Nd variation. The c. 112 Ma group (Shag and Eweburn) lie within the field defined by the composition of their host Rakaia sandstones (Roser & Korsch 1999), and thus were likely derived largely from partial melting of the base of the Torlesse accretionary prism. The 101 Ma group show a wide range in crustal components (Stitts . Motu . Houhora). The Houhora ignimbrite has a relatively primitive isotopic composition indicating only minor crustal contribution. The apparent conflict with abundant and diverse zircon inheritance might be explained by accidental incorporation. The two felsic samples of the c. 97 Ma group (DSDP 207A and Takahe rhyolites) with 1Nd  0 reflect modest crustal contamination and absence of major fractionation. They have much more enriched isotopic compositions than the mafic members (Mandamus, Tapuaenuku) of this age group. The French Creek granites are pulled to the right of the mantle array by combinations of very minor crustal assimilation and major fractionation, as discussed by Waight et al. (1998b).

Because our 112 Ma ages for the tuffs in the hanging wall of the normal Waihemo Fault are 5– 12 Ma older than that inferred from K– Ar dating (Adams & Raine 1988), they have new implications for the Cretaceous exhumation of the Otago Schist (the core to the Eastern Province accretionary wedge). Bishop & Laird (1976) interpreted the deposition of Kyeburn Formation to be coeval with movement on two high angle normal faults (Waihemo and Dansey Pass Faults, Forsyth 2001). This high-angle brittle fault movement would thus seem to be as old as 112 Ma, and now overlaps the youngest ages of extensional ductile shear zone formation in the schist core (109 –112 Ma; Forster & Lister 2003). Butz (2007) has proposed a model structurally linking the high angle Waihemo Fault to a nearby low angle fault, Footwall Fault, in the schist, via an extensional rolling hinge model. Thus, our new 112 Ma ages allow, for the first time, a plausible demonstration of synchronous deformation at middle crustal ductile levels and

supracrustal brittle levels of the Eastern Province accretionary wedge. It is also possible that the source of the igneous-derived hydrothermal fluids at the Macraes gold mine (within 20 km of the Waihemo Fault) (de Ronde et al. 2000) is related to our 112 Ma rhyolitic magmatism.

Initiation of widespread crustal extension We are not sure if the 112 Ma Eastern Province extension has anything to do with later, more widespread, Gondwana rifting (Forster & Lister 2003) or just represents intra-wedge exhumation (Deckert et al. 2002). The overlap in age between extension and deposition of the youngest Torlesse Eastern Province sediments deposited in a trench environment (Cawood et al. 1999) would seem to favour the latter. Our work has revealed discrete episodes of rhyolitic volcanism between 112 and 82 Ma, but it is much more difficult to relate these to punctuated episodes of extension. The development of the

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Fig. 10. Map showing reconstructed geological setting of sampled rocks, and related rocks of similar age mentioned in the text, in the Cretaceous interval 120– 80 Ma. Rocks dated in this study are in bold type. Median Batholith (230–105 Ma) is comprised of a 230–130 Ma low Sr/Y gabbro– diorite–granite belt on its ocean side, and a 130–105 Ma high Sr/Y diorite–granite belt on its continent side. Thick dashed line shows position of continent – ocean boundary (palaeotrench). 120 Ma ridge is schematic, position of Takahe locality is projected onto diagram from further east. Tasman– Southern Ocean spreading ridge at c.52 Ma is from Sutherland et al. (2000). Pap, Paparoa core complex; Ssz, Sisters Shear zone; Fre, French Creek Granite and Hohonu Dike Swarm; Gal, Galleon Volcanics; Som, Mount Somers; Tap, Tapuaenuku; Man, Mandamus; BPt, Buttress Point. Reconstruction after Mortimer et al. (2005) based on Sutherland (1999), Gaina et al. (1998) and Eagles et al. (2004). Modified here to include Fiordland block, located as follows. Given the parallelism of Sisters Shear Zone and Paparoa ductile stretching lineations we rotate stretching lineations of the intervening Fiordland block (average of Gibson 1990; Klepeis et al. 2007 ¼ 060o) 728 counterclockwise to parallel Sisters Shear Zone and Paparoa core complex lineations. This agrees with the angular

CRETACEOUS FELSIC VOLCANISM

Paparoa metamorphic core complex in the Paparoa Range, dated by the 100 Ma Stitts Tuff, is generally regarded as a ‘Gondwana break-up’ precursor. Episodes of pre-101 Ma extension in the Median Batholith and Western Province have been proposed as indicating regional continental extension as early as 108 –111 Ma on the basis of granulite-facies metamorphism (Gibson & Ireland 1995), deformed plutons (Muir et al. 1994) and cross-cutting dykes (Scott & Cooper 2006; Klepeis et al. 2007). However, just as with the Eastern Province discussed above, extensional events at this time and in these places is not inconsistent with overall oceancontinent plate convergence, as a subducting slab in the mantle does not preclude coeval intra- or back arc continental extension in the crust. Unambiguous subduction-related magmatism (outboard low Sr/Y I-type belt of Median Batholith) largely ceased at about 138 Ma (Kimbrough et al. 1994, Tulloch & Kimbrough 2003); the subsequent 131– 105 Ma high Sr/Y Separation Point Suite is possibly explained by partial melting of a mafic underplate derived from 230 –138 Ma magmatism with the older, insulating, slab removed.

Tectonic controls on Cretaceous rhyolitic volcanism The shape of some individual intrusive bodies and the outcrop length of volcanics and/or controlling faults are shown in Figure 10. Included are structures from the Sisters Shear Zone (Kula et al. 2007) and the Late Cretaceous Galleon volcanics of offshore South Island. Ages of initiation of rapid exhumation of mid-crustal levels core complexes can be correlated with the igneous episodes. The 101 Ma igneous event may relate to an early Paparoa core complex event (the lower plate was not exhumed at that time; Tulloch & Palmer 1990) and the 97 Ma event may correlate with a more rapid cooling/exhumation episode in the Paparoa core complex (Spell et al. 2000). Likewise, the 88 Ma igneous event (Whataroa) coincides with initiation of rapid exhumation on the Sisters Shear Zone (Kula et al. 2007). Although undated, the Galleon volcanics lie within Albian–Santonian (112 –83 Ma) ENE-striking grabens, e.g. Clipper

111

Basin (Field et al. 1989). Figure 10 emphasizes that our 101 and 97 Ma groups occur right across the width of Zealandia from very close to the inferred palaeotrench (Houhora, Takahe) to interior Gondwana well back from the then-recently extinct Median Batholith arc. The structural trends in the 101 and 97 Ma groups strike at c. 308 to the inferred palaeotrench and Median Batholith orogenic trend, but are, in general, subparallel to the closest segment of then-yet-to-be formed spreading centres, particularly the Southern Ocean centre between Zealandia and West Antarctica. Our data do not unambiguously resolve models of how subduction ceased beneath the Zealandia part of the Gondwana margin. A subducted spreading ridge model (Bradshaw 1989) is perhaps the most likely explanation for the Eweburn and Shag Valley accretionary wedge magmatism at 112 Ma, although slab roll back or foreland slab tear (Shoonmaker et al. 2005) may also be possibilities. The lack of parallelism of the widespread and strong (Fig. 10) 101–97 Ma extension directions with the palaeotrench suggests that the Gondwana– Zealandia continental plate and the mosaic of Pacific – Phoenix oceanic plates had effectively become welded by that time, as predicted by the platecapture hypothesis of Luyendyk (1995). The arrival of the Hikurangi Plateau at the Gondwana margin may also have been involved in cessation of subduction (Davy et al. 2008). Figure 10 shows there is no space-time trend in the distribution of either the 101 Ma or the 97 Ma igneous activity (unless the c. 1 Ma difference between Stitts and Motu tuffs is considered significant); there was a flare-up of 101 and 97 Ma age magmatism across the entire length and width of Zealandia. This would seem to argue against models involving propagation of a triple junction, slab window, or laterally spreading plume head beneath Zealandia (e.g. Bradshaw 1989; Weaver et al. 1994). If any of these scenarios were the case, moderately fast horizontal plate motion rates of 100 mm a21 would have resulted in variations in the age of magmatic and tectonic features within Zealandia by 4 Ma across distances of 400 km. Instead, with our new dataset of precisely dated siliceous magmatism across and along .1000 km baselines in Zealandia (Fig. 10) twin pulses of

Fig. 10. (Continued) difference between the Median Batholith strike and the overprinting extension direction (408), in both Fiordland and Stewart Island. In this orientation, the Puysegur Group basin subparallels the Pororari Group (including Stitts Tuff) of Westland, and the mafic dyke swarm on Five Fingers Peninsular subparallels the lamprophyric dyke swarms of Westland. The block is then moved southwards until the southeastwards continuation of the Buller– Takaka terrane boundary in southern Fiordland (Allibone et al. 2007) passes between southern Stewart Island (Takaka Terrane; Allibone & Tulloch 2004) and the southern most wells in the Great South Basin (Buller Terrane; Cook et al. 1999). A corollary of this reconstruction is that some of the northeasterly trending faults in Fiordland (e.g. Dusky, Breaker Point) may be post-break-up transfer faults such as those in Westland (e.g. Beggs et al. 2008) and Stewart Island (Kula et al. 2007).

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widespread magmatism are seen, suggesting that at 101 and 97 Ma the entire Zealandia underwent effectively instantaneous tectonomagmatic changes. Ages of igneous and tectonic events from further along the margin in West Antarctica are similar (Siddoway et al. 2004). Collectively, these observations lend further support to the Luyendyk (1995) model of rifting control by plate capture. We also note that eventual break-up occurred well inboard from the axis of what would have been the hottest lithosphere: the Median Batholith. The latter also exhibits considerable extension in Fiordland and Sisters Shear Zone core complexes, yet did not lead to break-up. We speculate that break-up did not occur in the vicinity of the Median Batholith because the hotter lithosphere was able to accommodate stretching via ductile flow. That break-up eventually occurred well within Gondwana, and cut across the Median Batholith and its continuation into Marie Byrd Land (Bradshaw et al. 1995) suggests that a colder, more brittle lithosphere was more amenable to fracturing.

Controls on degree of A-type chemistry A-type and peralkaline magmatism is a minor component of the mainly I-type Median Batholith plutonic record (Mortimer et al. 1999). For example, a small pulse occurs at the change from low Sr/Y to high Sr/Y suites at about 134 Ma (Tulloch & Kimbrough 2003), but a number of minor A-type bodies remain undated. Thus minor A-type magmatism appears to be a normal component in long-lived subduction related magmatism that is dominated by I-type granitoids. In the 101 Ma group, the Houhora ignimbrites, including our dated sample, have relatively strong A-type characteristics, and Stitts Tuff (despite contamination) has a trace of A-type character. Stitts Tuff is clearly spatially associated with early extension in the Paparoa Metamorphic Core Complex (Tulloch & Kimbrough 1989; Tulloch & Palmer 1990). The 97 Ma group contains a wide range of chemical types, from A-type felsic rocks variably contaminated with crust (DSDP 207A, Takahe), mafic undersaturated OIB (Tapuaenuku, Mandamus), and an oversaturated basalt to rhyolite suite reflecting major contamination of a depleted mantle derived magma by Eastern Province greywacke (Mount Somers; Barley 1987). We speculate that the degree of anorogenic (A-type) v. orogenic (I-type) characteristics of the rhyolitic magmas, in the absence of contemporaneous subduction, is basically related to crustal thickness. The principle here is that A-type character would be less likely to be affected by crustal contamination if the crust was thinner. Overall we see an increase in felsic A-type character from the 101

and 97 Ma groups to the 85 Ma group that we relate to progressively decreasing crustal thickness during this time interval. This analysis excludes the strongly silica-undersaturated mafic magmas of Tapuaenuku and Mandamus.

Conclusions New U –Pb and Ar/Ar ages for seven Cretaceous felsic igneous rocks in Zealandia, have been combined with previously published data to reveal a strong episodicity in silicic magmatism in the interval between cessation of arc magmatism and continental break-up. 112 Ma tuffs now recognized in the Eastern Province are unlike coeval plutonic rocks of the Median Batholith, and we speculate that a subducted mid-ocean ridge may have been responsible for this rare and unusual magmatism in the accretionary wedge. Both 101 and 97 Ma groups of rhyolites and tuffs occur across the entire width of Zealandia from near the palaeotrench to the continental interior indicating widespread, near instantaneous, extension. An 82 Ma group are known only from western South Island. Overall, the rhyolites, tuffs and granites show an increase in A-type character with time, from 112–101– 97 –88–83 Ma. We attribute this to the progressive thinning of the Zealandia continental crust whereby, with time, there is less opportunity for crustal contamination. Available extension directions associated with 101, 97 and 82 Ma magmatism and associated core complex exhumation in accretionary wedge, Median Batholith and interior Gondwana are all oriented c. 308 oblique to the then recently inactive subduction margin. This orientation is consistent with rifting being caused either by basal traction on a subducted slab that had been captured and pulled oceanwards by the Pacific plate, and/or southwestwards propagation of a ,83 Ma oceanic ridge between Zealandia and West Antarctica. Our observations are less supportive of rifting models based on laterally migrating triple junctions, slab windows or plume heads.

Appendix 1. Radiometric dating analytical methods U – Pb TIMS Samples analysed at MIT were single grains, and most have been abraded to reduce Pb-loss (Krogh 1982) followed by washing in 4 N HNO3 at 80 8C. Selected zircons for analysis were loaded into FEP Teflonw microcapsules, spiked with a mixed 205Pb– 233U – 235U tracer solution and dissolved in concentrated HF at 220 8C for 48–60 hours. Dissolution was followed by conversion to chloride form using 6 N HCl at 180 8C for 12 hours. Pb

CRETACEOUS FELSIC VOLCANISM and U were separated using a miniaturized HC1-based ion-exchange chromatography procedure modified after Krogh (1973). Both Pb and U were loaded with a silica gel-H3PO4 emitter solution on single degassed Re filaments and their isotopic compositions were measured on the VG Sector 54 TIMS at MIT using methods similar to those described in Schmitz & Bowring (2001). The most recently analysed samples at MIT have been thermally annealed and leached to reduce the effects of Pb-loss, after the methods described by Mattinson (2005). Age calculations used Isoplot v3.0 (Ludwig 2003). In order to minimize the effects of Pb loss, the grains were pre-treated in one of two ways; conventional mechanical air-abrasion (Krogh 1982) followed by fluxing in 4 N HNO3 at 80 8C, or a version of the thermal annealing and acid leaching (also known as chemical abrasion or CA-TIMS) technique of Mattinson (2005), prior to addition of a mixed 205Pb– 233U– 235U tracer solution and complete dissolution. Details of zircon pre-treatment, dissolution and U and Pb chemical extraction procedures are described in Ramezani et al. (2007). Pb and U were loaded together onto a single degassed Re filament in a silica-gel/ phosphoric acid mixture (Gerstenberger & Haase 1997) and their isotopic compositions were measured on a VG Sector 54 thermal ionization mass spectrometer. Pb isotopes were measured by peak-hopping using a single Daly photomultiplier detector and U isotopic measurements were made in static mode using multiple Faraday collectors. Mass fractionation for Daly measurements was determined to be 0.25 + 0.04%/amu over a wide temperature range based on long-term measurements of the NBS-981 Pb standard. U mass fractionation was calculated in real-time using a double spike. All common Pb was attributed to procedural blank. Data reduction, age calculation, and generation of concordia plots were carried out using the algorithms of Ludwig (1980), and the statistical reduction and plotting program ISOPLOT (Ludwig 2003).

U– Pb TIMS data interpretation and error treatment Our most recent results at MIT involved analyses of single zircon grains with total radiogenic Pb contents of as little as 3 pg and internal uncertainties of measured dates in the order of 0.1% or better. Furthermore, the application of the CA-TIMS method (Mattinson 2005) has significantly enhanced the accuracy of U–Pb dates by selectively removing zircon crystal zones prone to radiation damage and Pb loss. With improved analytical precision, however, additional sources of error that were previously deemed negligible have become significant and must be taken into account. It has now become evident that at high precision even statistically coherent and highly reproducible sets of zircon U– Pb data exhibit slight age discordance

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manifested by 207Pb/206Pb dates that are c. 0.2% older than the associated Pb/U dates. This discordance is likely due to imprecision in one or both of the U decay constants (Jaffey et al. 1971; Mattinson 2000; Schoene et al. 2006) rather than open system behaviour in zircon, and thus limits the application of age concordance (e.g. ‘Concordia age’ of Ludwig 1998) in the assessment of high-precision U– Pb data. This source of systematic uncertainty is particularly important when isotopic dates measured from different chronometers (e.g. U –Pb and 40 Ar/39Ar) are compared. Another systematic uncertainty of significance is the spike calibration error that must be accounted for when comparison is made between highprecision U– Pb dates produced in different laboratories (i.e. using different spike solutions). The interpretation of magmatic crystallization ages from U–Pb data involves the choice between multiple isotopic dates, including the 206Pb/238U, 207Pb/206Pb and concordia intercept dates. In general, results from a single, inheritance-free, population of zircons for which Pb loss is non-existent or has been eliminated is represented by a coherent cluster of data (termed ‘equivalent’ by Ludwig 1998), in which all scatter can be accounted for by analytical uncertainty. In these cases the weighted mean 206 Pb/238U date of the cluster is considered the most precise and accurate representation of crystallization age. When coherent data sets are not obtainable due to persistent Pb loss that is inferred to have occurred in geologically recent time, we instead rely on the 207 Pb/206Pb date as the best estimate of the crystallization age. In this case, uncertainties in the decay constants can be generally ignored. If the condition of young-age Pb loss (i.e. lower concordia intercept close to zero) is not satisfied, or zircon inheritance is involved, estimation of the crystallization age will inevitably rely upon the concordia intercept dates. In order to demonstrate all relevant sources of uncertainty the U–Pb date errors are reported as follows: +X[Z], where X is the internal uncertainty in absence of all systematic error (tracer calibration and decay constants), and Z includes the tracer calibration (using a conservative estimate of the 2s standard deviation of the Pb/U ratio in the tracer to be 0.05%) as well as decay constant errors of Jaffey et al. (1971) (external uncertainty). For 207 Pb/206Pb dates, tracer errors are negligible and Y is not reported (so it reads +X[Z]). The MSWD (mean square of the weighted deviates: York 1967, 1969) of the 206 Pb/238U weighted mean date is calculated prior to the addition of systematic uncertainties.

U – Pb ICPMS U –Pb dating of zircons from the Shag River ignimbrite was done at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia, Vancouver, Canada. Zircons were separated using conventional mineral separation methods. Approximately 25 of the coarsest, clearest, most inclusion-free grains

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were selected, mounted in an epoxy puck along with Plesovice standard zircon (Sla´ma et al. 2007) and brought to a very high polish. The grains were examined using a stage-mounted cathodoluminescence unit and no inherited cores or alteration zones were found to be present. The surface of the mount was then washed for approx. 10 minutes with dilute nitric acid and rinsed in ultraclean water. The same instrumentation and very similar operating parameters as those described by Chang et al. (2006) are used in the PCIGR, including a New Wave UP-213 laser ablation system and a ThermoFinnigan Element2 single collector, double-focusing magnetic sector ICP-MS. Line scans rather than spot analyse were employed in order to minimize elemental fractionation during the analyses. A 40% laser power and a 25 micron spot were used. Backgrounds were measured with the laser off for seven seconds, followed by data collection with the laser on for approximately 23 seconds. The time-integrated signals were analysed using the GLITTER software package described by van Achterbergh et al. (2001) and Jackson et al. (2004). Corrections for mass and elemental fractionation were made by bracketing analyses of unknown grains with replicate analyses of the 1099.1 Ma FC-1 zircon standard (Schmitz et al. 2003). Errors on the final calculated age of the sample are given at the 2 sigma level.

Ar/Ar Kyeburn Tuff biotite samples and irradiation monitor TCR-2 sanidine (Taylor Creek Rhyolite; Age ¼27.87 Ma; Lanphere & Dalrymple 2000) were irradiated in position E6 of the FRG-1 nuclear reactor at the GKSS Research Center, Geesthacht, using Cd shielding. Step-heating 40 Ar/39Ar analyses were carried out with a 20 W argon-ion laser, and an MAP 216 mass spectrometer. System blanks were measured prior to each sample, and after each fifth sample heating step, typically making up 10%, 1%, and 2% of the measured 36Ar, 39Ar, and 40Ar isotopes, respectively. We thank Mike Isaac and the Ocean Drilling Program repository for supplying some of the rocks for this study. Sam Bowring supported the ID-TIMS work at MIT. Technical help was provided by John Simes, Belinda Smith Lyttle and Philip Carthew. Tod Waight generously provided unpublished whole rock analyses. Keith Klepeis, Joseph Kula and James Scott are thanked for discussion. The manuscript was improved by comments from John Bradshaw and Christine Siddoway. Funded by the New Zealand Foundation for Research Science and Technology.

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Structure and evolution of the western Corinth Rift, through new field data from the Northern Peloponnesus EMMANUEL SKOURTSOS* & HARALAMBOS KRANIS Department of Dynamic, Tectonic and Applied Geology, Faculty of Geology and Geoenvironment, University of Athens, Panepistimiopolis, Athens GR 15784, Greece *Corresponding author (e-mail: [email protected]) Abstract: Extensional structures with geometrical and kinematic features analogous to the known Gulf of Corinth faults, are found further to the south of what is considered to be the southern margin of the of Proto-Corinth Gulf, reaching south to the northern flanks of Mt Mainalon. This mountain front is marked by the North Mainalon Fault Zone, which comprises a series of normal fault segments with NNE dips. Assuming a listric or ramp-flat geometry for the North Mainalon Fault Zone, it could flatten at a depth of 6–8 km, underneath Mt Khelmos. Its southern, shallow part has been truncated by NNE- and NNW-trending faults, which may be linked to northward propagation of the east–west extension in the Southern Peloponnesus, causing further uplift in the central and northern Peloponnesus, while its deeper part is still active and may reach further north and sole onto the hypothesized detachment zone beneath the Gulf of Corinth.

One of the most important results of extended continental crust is the creation of extensional basins, which host structures that have been active during the development of the basins, as well as sedimentary sequences, the dating of which offers useful constraints on the activity of the tectonic structures (Schlische 1991; Gawthorpe & Leeder 2000). The formation of these basins is mainly controlled by normal faults, the geometry and development of which have been investigated by diverse means and methodologies (seismological investigations, geophysical imaging, tectonic studies, etc.). Although there are significant differences in the faulting pattern encountered in various extensional basins, there are several aspects that are generally agreed upon, namely: (i) basin marginal normal faults have dips between 308 and 608 close to the Earth’s surface; (ii) fault length is dependent on the thickness of the seismogenic upper crust – the latter also controls the dimensions of fault-bounded blocks; (iii) fault-bounded blocks are often rotated, albeit not always to the same degree (Jackson 1987; Jackson & White 1989; Scholz & Contreras 1998; D’Agostino et al. 2001). On the other hand, what is still a matter of debate is fault geometry at depth and particularly whether marginal faults are of listric or planar type and if they sole at depth on a low-angle normal fault (detachment) (e.g. Lister & Davis 1989; Koyi & Skelton 2001; Fossen et al. 2003). One such region that has been in the limelight of the geoscientific interest, at least for the last 20 years, is the Corinth Rift, in the inner part of the

Hellenic Arc (Fig. 1). The Corinth Rift offers the opportunity of testing the validity of these models, as it is one of the representative examples of young continental rifting (Sorel 2000), with high seismicity and deformation rates (e.g. Bernard et al. 1997; Clarke et al. 1998; Briole et al. 2000; Avallone et al. 2004). Despite the variety of suggested models, its structure, evolution and the relationship between the surficial brittle structures and the deep, seismogenic ones, are still matters of debate. Central in this debate has been the manner in which net strain across the Corinth Gulf is accommodated. Rigo et al. (1996) identify a subhorizontal, north-dipping zone of microseismicity, at depths of 8–10 km and postulate that this may correspond to either a low-angle active fault or a low-angle detachment zone lying at 9–11 km depth. They also suggest that microseismicity is clustered at the intersection between this subhorizontal zone and the steep normal faults of the southern part of the gulf, a hypothesis that had also been previously made by Doutsos & Poulimenos (1992). Hatzfeld et al. (2000) propose that the background microseismicity at depths of 8–12 km is not related to any subhorizontal active fault, but to the brittleductile transition in the intermediate crust. Pham et al. (2000) identify a relatively conductive, 4 –km thick layer at a depth of 10 km, while in a tomographic study, Latorre et al. (2004) postulate that the sharp velocity change at 5–7 km corresponds to the contact between the Tripolis Unit carbonates and the underlying Phyllites –Quartzites

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 119–138. DOI: 10.1144/SP321.6 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Simplified geological map of North Peloponnesus and the Gulf of Corinth. Offshore faults are from Moretti et al. (2003). 1, synrift deposits; 2, Sub-Pelagonian Unit; 3, Parnassos Unit; 4, Pindos Unit; 5, Tripolis Unit; 6, Zarouchla Complex. Inset map shows the location of the study area.

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Unit (as also proposed by Le Pourhiet et al. 2004), which is a weak crustal zone between 7 and 9 km and constitutes a preferred environment for a continuous earthquake triggering. Che´ry (2001) argues that high-angle faulting may be responsible for detachment formation and suggests that the Gulf of Corinth is in the initial stages of core complex formation. Bernard et al. (2006) also speak about a subhorizontal seismic layer on which the steep main normal faults may root and Gautier et al. (2006) conclude that low-angle faulting contributes to the measured extension in the western part of the Gulf of Corinth. On the other hand, McNeill et al. (2005) argue that the series of high-angle faults between Aigion and the northern coast of the gulf is enough to accommodate the observed net strain across the gulf and that slip on the postulated detachment is superfluous to match extensional strain. Recently, Bell et al. (2008) suggest that slip on a detachment, at least during the Holocene, should be minimal and Cianetti et al. (2008) also suggest that a subhorizontal discontinuity is not necessary to model the observed extension. On the onshore southern shoulder of the Rift, Sorel (2000), Flotte´ & Sorel (2001) and Flotte´ (2002), mapped the southernmost outcrops of synrift deposits, related to the Corinth Rift on the northern slopes of Mt Khelmos. They suggest that (i) the surficial expression of the postulated discontinuity beneath the Corinth Gulf corresponds partially to the low-angle normal fault which marks the northern flanks of Mt Khelmos, where synrift sediments crop out and partially to the tectonic contact between the metamorphic Zarouchla Complex and the overlying carbonates of Tripolis and Pindos Units; (ii) the high-angle normal faults north of the detachment front are secondary structures that sole onto the detachment surface. This model, however, fails to account for a number of observations (see also Westaway 2002; Moretti et al. 2003), such as: (i) seismological data from northern Peloponnessus show hypocentral distribution at depths of 6–15 km, which is deeper than the suggested detachment zone (Jansky´ et al. 2004; Lyon-Caen et al. 2004; Zahradnı´k et al. 2004); (ii) the assumption that there is a northward younging of synrift deposits does not seem to be always valid, as is has been found that the older sediments in the various sub-basins are contemporaneous (Ori 1989; Rohais et al. 2007); (iii) the assumption of unimodal (northward) migration of fault activity may not be always valid, as recent geochronological data show out-of-sequence activity (Causse et al. 2004); (iv) the suggested detachment is too shallow to allow for the formation of hangingwall faults with length and throw in excess of 10 km and 2 km, respectively (Ghisetti & Vezzani 2005);

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(v) the overall uplift of Northern Peloponnessus may have resulted from processes other than simple, fault-induced isostatic rebound (e.g. Collier et al. 1992; Westaway 1998; Che´ry 2001). Ghizetti & Vezzani (2004, 2005) suggest that (i) a mechanical and geometrical barrier exists between the eastern and western parts of the Rift, almost perpendicular to the east–west normal faults, (ii) the active normal faults reach deeper in the eastern part of the Rift and (iii) the detachment layers, on which the normal faults may root are discontinuous and non-planar from east to west. This barrier (‘Zarouchla culmination’, according to Ghizetti & Vezzani 2005) has resulted from NNE to NE Miocene extension that affected the Hellenic nappe pile, a view also adopted by Papanikolaou & Royden (2007). This paper presents the results of field work in North Peloponnessus, focused on the extensional tectonics that postdates the Neogene Hellenic nappepile. The geological and tectonic mapping that was carried out shows that extensional structures are observed not only north of what is considered to be the southern margin of the Proto-Corinth Rift, but also south of it. These structures bear geometrical and kinematic characteristics that fit with those published for the Corinth Rift. The aim of this paper is the investigation of probable links between these structures and the deformation processes associated with the evolution of the Corinth Rift.

Geological and tectonic setting of the Corinth Rift: lithostratigraphy The Corinth Rift is a WNW –ESE extensional basin, which separates central mainland Greece from the Peloponnessus, extending from the Rio straits and Mt Panachaikon in the west to the Alkyonides Gulf and Corinth in the east (Fig. 1). It is a young basin, with its northern portion lying offshore, below the Gulf of Corinth; the latter is approximately 115 km long, its width increases from about 10 km in the west to 30 km in the east and its deepest part (900 m) is located at the central section. Brooks & Ferentinos (1984) characterize the Gulf as an asymmetrical tectonic basin (half graben), a suggestion later disputed by Moretti et al. (2003), who characterize it as a symmetrical graben. Stefatos et al. (2002) describe it as a composite asymmetric graben, with varying geometry along strike, with the master faults that control the rift being north-dipping ones in the eastern part and south-dipping in the western (Sachpazi et al. 2003; McNeill et al. 2005). The Rift also includes a 20 km wide landstrip, which is actually its southern shoulder, in northern Peloponnessus, where synrift deposits are the main

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outcrops, consisting mostly of fluvio-lacustrine facies and Gilbert-type fan deltas (Ori 1989; Rohais et al. 2007). Normal faulting has affected not only the synrift sediments, but also the pre-rift, Mesozoic– Early Cenozoic formations, which form a nappe pile consisting of carbonate, clastic and metamorphic rocks (Doutsos et al. 1988). The vast majority of researchers have adopted the notion that fault activity has shifted from south to north (e.g. Dufaure 1975; Ori 1989; Doutsos & Poulimenos 1992; Goldsworthy & Jackson 2001), a concept that led to the conclusion that the rift was wider at the early stages of its evolution. Doutsos & Piper (1990) suggest that these faults are of listric type; however, quite a few researchers argue that there is not enough evidence to support the suggestion for listric geometry and favour the model of planar faults (e.g. Westaway 2002; Moretti et al. 2003; Rohais et al. 2007). Sorel (2000), Flotte´ & Sorel (2001) and Flotte´ (2002) described the southernmost of these faults as a low-angle detachment fault, with shallow, high-angle faults soling onto it. The study of faults at the southern shoulder of the Corinth Rift and within the present-day gulf (e.g. Doutsos & Piper 1990; Poulimenos 1991; Flotte´ et al. 2005; McNeil et al. 2005), focal mechanism solutions (e.g. Rigo et al. 1996) and the results of geodetic investigations (Billiris et al. 1991; Briole et al. 2000) have shown that the rift is subjected to north–south extension at 1.0–1.5 cm a21. Armijo et al. (1996) suggest that 0.7 cm a21 of extension are accommodated through high-angle normal faulting, which accounts for 50–75% of the total extension rate. However, these are considered to be overestimated and Westaway (2002) suggests an extension rate of 0.2 –0.4 cm a21 for the last 2 Ma. An important issue is that the Northern Peloponnesus is an extending domain, which is also undergoing uplift and notably at different rates in its eastern and western parts (Zelilidis 2000; Houghton et al. 2003; Pirazzoli et al. 2004). The overall uplift has resulted in the Pliocene– Pleistocene sediments having been uplifted to altitudes of 1800 m (Dufaure 1975). Various suggestions have been made as to the driving mechanism behind it, such as isostatic response to subduction-driven sediment accretion (Le Pichon & Angelier 1981; Collier et al. 1992), compression-driven crustal thickening (Mariolakos & Stiros 1986), magmatism (Collier 1990), westward crustal inflow (Armijo et al. 1996), or southward lower-crustal flow from the beneath the modern Gulf to beneath Northern Peloponnesus (Westaway 1996). The synrift outcrops at the southern, onshore shoulder of Corinth Rift comprise sediments, the type of which was related to the magnitude of the (vertical) fault activity and the fluctuations of the sea level (Dart et al. 1994). The combination

of these two factors led to the deposition of either marine or terrestrial sediments. The synrift deposits can be divided in two lithostratigraphic units (Ori 1989), with a total thickness of up to 2.8 km (Rohais et al. 2007). The lower unit consists of alluvial fan to shallow-water lacustrine sediments, approximately 1500 m thick, while the upper comprises alluvial fan and Gilbert-type fan sediments (Rohais et al. 2007). The base of the upper unit, represented by a thick fanglomeratic unit represents, according to Ori (1989) a marked change in the basin evolution. The Gilbert-type deltas and deep sea sediments of the upper unit are characterized by thick, moderately to steeply inclined strata belonging to a steep margin of a deep basin; two sets of deltaic deposits have been distinguished, with the top formations of the older one (Evrostini delta) found now at altitudes of 1200–1700 m, while the topsets of the younger are at 300– 400 m (Rohais et al. 2007). The successive development of these deltaic deposits is controlled by the high-angle normal faults of Northern Peloponnesus. These two thick synrift units are overlain, either conformably or not, by Middle to Upper Pleistocene deposits, which include perched marine terraces and fluvial terrace deposits. (Keraudren & Sorel 1987; Armijo et al. 1996; Pirazzoli et al. 2004; Rohais et al. 2007). Pliocene and Quaternary sediments crop out further to the south, in the closed hydrological basins of Kandila and Levidi (Lu¨ttig 1976; Alexopoulos 1998; Hambilomati 2005) (Fig. 2). These are mainly grey lacustrine marls and sands with sapropel intercalations, with a maximum thickness of 95 m, overlain by thin alluvial fan and fluvial/terrestrial deposits. As far as their age of the marls and sands is concerned, they may be equivalent to the Late Pliocene sediments of the Tripolis and the Kalavryta basins (Lu¨ttig 1976; Papanicolaou et al. 2000). The synrift sediments have covered the Alpine nappe pile, which appears heterogeneous throughout the study area. The nappe stacking in western mainland Greece and Peloponnesus took place through progressive westward migration of thrust movements (Aubouin 1959; Richter 1976, 1993; Fleury 1980); this phase lasted from the Late Eocene to Early Miocene times. The deepest Unit of the pile is the Zarouchla Complex (Dercourt 1964; De Wever 1975), which was later distinguished by Lekkas & Papanikolaou (1979) and Dornsiepen et al. (1986) in two units: a lower, metamorphic one, consisting of mica schists and quartzites, which corresponds to the Phyllites – Quartzites Unit (Dornsiepen et al. 1986) and crops out within the Krathis valley and north of the Feneos and Stymfalia basins (Fig. 2); and an upper one, consisting of schists, tuffs, volcanics and limestones, which corresponds to the Upper

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Fig. 2. Geological –structural map of the study area (see Fig. 1 for location). EF, Eliki Fault; PMF, Pyrgaki–Mamoussia Fault; DoF, Doumena Fault; TF, Tsivlos Fault; KhF, Khelmos Fault; KrF, Krathis Fault; GF, Gardiki Fault; MF, Monodendri Fault; DF, Drossopigi Fault; LF, Lykouria Fault; PF, Penteleion Fault; SF, Saitias Fault; AgF, Agridi F; NMFZ, North Mainalon Fault Zone; OF, Orchomenos Fault; Ka, Kalavryta Basin; F, Feneos Basin; S, Stymfalia Basin; D, Dara Basin; K, Kandila Basin; L, Levidi Basin; T, Tripolis Basin. Lines with diamonds on major faults denote the mean slip vector orientation. The locations of the measurements projected in the stereographic plots of Figure 4 are shown by the respective circled numbers.

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Palaeozoic–Upper Triassic Tyros Beds. These two units are separated by a tectonic contact, which, according to Lekkas & Papanikolaou (1979), Dornsiepen et al. (1986) and Xypolias & Koukouvelas (2001) is a major thrust fault, although it has also been found to correspond to a large-scale detachment in southern Peloponnesus and Crete (Jolivet et al. 1996; Fassoulas et al. 1994; Ring et al. 2001; Skourtsos 2002; Skourtsos et al. 2004). The Tyros Beds are the volcanosedimentary base of the Tripolis Unit, a Mesozoic –Lower Cenozoic shallow-water carbonate sequence, topped by a thick clastic formation (flysch). The thickness of the Tyros beds varies and it has been found to be more than 800 m in south-eastern Peloponnesus, while the Tripolis carbonates are c. 2 km (Dercourt 1964; De Wever 1975; Zambetakis-Lekkas & Karotsieris 1986). The thickness of the flysch of the Unit is c. 300 m in Vytina, while in eastern Peloponnesus it exceeds 500 m. The uppermost unit of the nappe pile is the Pindos Unit, a pelagic sequence of limestones, radiolarites and clastic deposits of Upper Triassic – Upper Cretaceous, followed by Tertiary flysch, with a total thickness of c. 1300 m (Degnan & Robertson 1998; Skourlis & Doutsos 2000).

Structural setting and extensional fault geometry The study area extends from the Gulf of Corinth to the northern flanks of Mt Mainalon; and from Kleitoria in the west to Stymfalia in the east (Figs 1 & 2). The northernmost tectonic structure examined for the purposes of this study is the Valimi fault. The visible part of Valimi Fault (VF, Fig. 2) displays a marked curvature, both along strike and dip. Its western mappable part strikes NE –SW, its central segment strikes east –west and turns to NW–SE at its eastern visible part. Measurements taken from exposed fault surfaces at the crest of the fault showed dips of 20 –308, where the fault juxtaposes Pindos and Tripolis rocks against synrift conglomerates (Fig. 3a), while close to the base of the escarpment, dip values become higher. Corrugations on the main fault surface are oriented NNE; slip vectors from footwall fault arrays, within ,5 m from the main fault surface are oriented NW, NNW and north. The mean slip vector for the Valimi Fault is 002/52 (Fig. 4) and its throw is in the order of 1500 m. The lower members of the hanging-wall conglomerates have dips of 30 –408S and are truncated against the fault, while further to the west, where the fault tips out, the medial and upper members overlap the fault surface. The Tsivlos Fault (TF, Fig. 2) cuts and offsets Valimi Fault, according to our mapping. It strikes

NE– SW and corrugations and striations on its exposed surface are oriented N708E. The footwall of Tsivlos Fault is composed of the Tripolis carbonates, which overlie the Zarouchla Complex formations; the latter crop out within the Krathis valley but not north of the Tsivlos Fault. The lower stratigraphic horizons of the hanging-wall conglomerates dip at 30 –408SSW; dip values decrease gradually towards the stratigraphically younger horizons and eventually the upper-most members are found to seal the fault and the palaeorelief associated with it, west of the Krathis valley. The Khelmos Fault (KhF, Figs 2 & 5) is clearly visible on the western flanks of the Krathis valley, where it juxtaposes the lower stratigraphic horizons of Tripolis carbonates against Pindos rocks. Its mean strike is WNW –ESE and has moderate dips (c. 408N); mean slip vector is 350/35 (Fig. 4). The fault truncates the contact between the Tripolis carbonates and the underlying Tyros Beds. The nature of this contact is tectonic; however, according to our observations, it is considered to be a de´collement that developed between the platform carbonates and its volcanosedimentary base during the main nappe-stacking in Late Oligocene–Early Miocene times. Synrift deposits crop out at the hanging-wall of the Khelmos Fault, unconformable over the Pindos formations, dipping c. 208S, at Xirokampos (Fig. 5). This is the easternmost outcrop of the synrift sediments of the Kalavryta Basin, which are overlain unconformably by the cemented slope breccia fan deposits that cover the northern slopes of Mt Khelmos. The upper members of these breccias seal the Khelmos Fault, but the lower ones have been found to be faulted against the Tripolis carbonates (Fig. 3b). According to our observations, the Khelmos Fault continues east of the Krathis valley, juxtaposing the Pindos formations at Xerovouni (‘Xerovouni Fault’, Rohais et al. 2007) against the synrift conglomerates, while south-dipping slope breccias seal the fault, as it is the case on the west flank of the Krathis valley. The Zarouchla Complex occupies the core of a tectonic window in the topographic depressions within the mountains of Northern Peloponnessus: Mts Khelmos and Ziria. All along the north and NE margins of the Zarouchla Complex outcrops, slivers of Tripolis and/or Pindos Units are found between the formations of the Complex and the synrift sediments. The Krathis Fault (KrF, Figs 2 & 6a) juxtaposes the underlying ZC rocks against these wedges and is associated with intense cataclasis and dolomitization. The geometry of the Krathis Fault does not remain constant, despite the fact that its trace is quite linear in map view (Fig. 2). Kinematic analysis of fault slip data shows that the mean slip vector is towards the north (N048E) (Fig. 4).

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Fig. 3. Field photographs from exposures of fault surfaces. See Figure 2 for locations. (a) Exposed fault surface of the upper and gently dipping part of the Valimi fault, where the fault juxtaposes Pindos carbonates against synrift conglomerates. View towards the south. Horizontal field of view is approximately 10 m. (b) The Khelmos Fault, juxtaposing slope breccia deposits (Pt br) against Tripolis carbonates, at Valvoussi (see Fig. 2 for location). View towards the SE. (c) Striated surface of the Lykouria Fault, at the pass between Mts Penteleion and Saitias View towards the east. (d) The Kamenitsa fault, exposed on north of Vytina. View towards the ENE/E. (e) Polished and striated surface of the Levidi Fault. Footwall is composed of Tripolis carbonates (Tr). Faulting is accompanied by thick (.5 m) cataclasites. Person at lower left for scale.

The upper boundary of the Tripolis and Pindos slivers is the NW–SE, NE-dipping Monodendri Fault with a mean slip vector towards the NNE (N0148E) (MF, Figs 2 & 6a). The synrift sediments on the hanging-wall of this fault system present a synclinal geometry, with the synclinal axis subparallel to the fault trace. The synrift conglomerates on the southern flanks of Mt Ziria (Fig. 2), dip 35 –458S and their

southern boundary with the pre-rift formations is the NNE-dipping low-angle (c. 158) normal fault (Drossopigi Fault) (DF, Figs 2 & 7). The northern contact of these deposits is a steeply dipping unconformity surface (c. 458S). North of this location, the Pindos thrust dips towards the South, by c. 458. By restoring the conglomerates to their original geometry, the Drossopigi Fault acquires a dip of c. 608 NNE and both the

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Fig. 4. Equal-area, lower hemisphere stereographic projections of fault-slip data. Block arrow denotes mean vector orientation from Fischer statistics. For location of measurement stations see Figure 2. (a) Valimi Fault; (b) Monodendri Fault; (c) Khelmos Fault; (d) Krathis Fault; (e) Agridi Fault; (f) Kamenitsa Fault; (g) Levidi Fault.

unconformity and the Pindos thrust become almost horizontal. The southern and western boundaries of the Zarouchla Complex, although faulted, are characteristically different from the northern and eastern ones. Along this margin, the Tripolis platform overlies the Tyros beds, through a SW-dipping contact on Mt Khelmos and Mt Penteleion (Figs 2 & 6b, c) and maintains the same characteristics as on the western slopes of the Krathis valley. At the northern slopes of Mt Oligyrtos, the Tripolis/Tyros Beds contact corresponds to an east –west, south-dipping normal fault, the eastern prolongation of which

juxtaposes the Tripolis carbonates against the Phyllites –Quartzites Unit of the Zarouchla Complex (Fig. 2). Mt Khelmos is composed of Tripolis carbonates, dipping generally towards the SW (Fig. 6b). The peaks and the southwestern flanks of the mountain correspond to klippen of the Pindos nappe, which crops out a few kilometres further to the NW. The Pindos outcrops on Mt Khelmos are found at progressively lower altitudes towards the SSW, owing to faults of NW– SE strike as the Gardiki fault (GF; Figs 2 & 6b) (see also Dercourt 1964). The NE– SW, NW-dipping Penteleion Fault (PF,

Fig. 5. The Khelmos Fault, on the western slopes of the Krathis valley (see Fig. 2 for location).The fault truncates the tectonic contact between the Tripolis carbonates and the underlying Tyros Beds of the Zarouchla Complex. TB, Tyros Beds; Tr, Tripolis carbonates; Pi, Pindos carbonates; Pl-Pst c, synrift conglomerates.

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Fig. 2) separates Mt Khelmos from Mt Penteleion, where south-southwestward dipping carbonates of the Tripolis platform also crop out. The youngest stratigraphic members of the platform comprise imbricate thrust sheets. The topmost thrust sheet includes also the Tripolis flysch, upon which the Pindos formations have been thrusted, through a SW-dipping thrust (Fig. 6c), the geometry of which has been described by Dercourt (1964). The southern boundary of the Pindos platy limestones of Mt Penteleion is the NE–SW, Lykouria Fault, which dips 308NW (LF, Figs 2 & 3c). Two sets of slip vectors were identified on the Lykouria Fault: the older set, which is of pure normal character (slip vector towards NW), is overprinted by oblique-slip striations (slip vector oriented towards the West). The footwall of the Lykouria Fault is Mt Saitias, built of Tripolis carbonates with steady southerly dips (Fig. 6d). The flysch/carbonates contact in this area is defined by an east –west, south-dipping fault which has also affected the Pindos thrust. South of Mt Saitias is Mt Falkos, built of imbricated Pindos thrust sheets with overall south-southwesterly dips (Fig. 6d). The eastward continuation of Mt Saitias is Mt Oligyrtos (Kandila), which possesses identical geometrical and structural characteristics. North of Mts Saitias and Oligyrtos are the Stymfalia and Feneos basins (Fig. 2), which are two closed geomorphological features, have resulted from drainage reversals caused by back-tilting that may have taken at the beginning of the Late Pleistocene (Dufaure 1975; Seger & Alexander 1993; Zelilidis 2000; Rohais et al. 2007). Two closed continental basins are located south of Mts Oligyrtos and Saitias, namely the Kandila and Levidi basins. The Kandila basin (Fig. 2) is a flat depression north of Mt Mainalon, with a mean elevation of 630 m. It is filled with c. 95 m of grey lacustrine clays and silts, which locally contain detritic components from the Pindos and Tripolis Units (Hambilomati 2005). Two alluvial fans overlie the lacustrine deposits in the SW and SE of the basin: according to Lu¨ttig (1976), these fans developed during a glacial period of the Pleistocene. The Levidi Basin is located between the Kandila and Tripolis basins (Fig. 2), separated by minor basement culminations. It is a shallow basin, filled with Quaternary alluvium and debris. Two alluvial fans mark its southwestern border, along the foot of Mt Mainalon. The Dara Basin develops west of the Kandila basin, at a lower mean elevation (508 m) and it is filled with the lacustrine deposits of Kamenitsa (Karotsieris 1981), of uncertain age. These are multimictic, locally cohesive conglomerates with sandy matrix and clasts (max clast size: 30 cm) from both the Pindos and the Tripolis rocks.

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The aforementioned basins lie in the hangingwall of an extensional structure, hereinafter referred to as ‘North Mainalon Fault Zone’ (NMFZ). It comprises fault segments that mark the northern and northeastern flanks of Mt Mainalon, which dominates central Peloponnesus. Segments of this zone appear on the published IGME maps and in Karotsieris & Lekkas (1988) as the primary thrust surface of the Pindos Unit over the Tripolis platform. However, our field data show that it is an extensional structure, which overprints and truncates the pre-existing compressional ones. The Agridi Fault (AgF, Figs 2 & 8) is the western segment of the North Mainalon Fault Zone. It comprises two overlapping branches, one with low to moderate dips (35 –408), truncated by a steeper branch (mean dip 508NE) (Figs 6e & 8). The mean slip vector of the fault is toward N108E (Fig. 4). The footwall of the Agridi Fault consists of Palaeocene –Eocene Tripolis limestones, with gentle SW dips. These carbonates bear internal imbricated thrusts which dip towards the SW (Fig. 6e). The uppermost thrust sheet includes the flysch of the Unit, tectonically overlain by the Pindos formations, which crop out quite further to the south (west and south of Vytina) (Karotsieris 1981). A conservative estimate of the cumulative vertical displacement of the Agridi Fault is 800 m and the minimum estimated heave must be 4 km (in a north–south direction) (Fig. 6e). The eastern segment of the North Mainalon Fault Zone is the Orchomenos Fault (OF, Fig. 2), the trace of which is quite sinuous, owing to its low NNE dips and mean slip vector towards the NNE (N158E). The two aforementioned segments are truncated by NNW –SSE and NNE–SSW high-angle normal faults, which control the present-day morphology of Mt Mainalon, composed of Tripolis carbonates, folded in a large-scale asymmetrical anticline trending NNW–SSE (Karotsieris & Lekkas 1986). The gently dipping western limb is truncated by the NNE –SSW to north –south Kamenitsa fault, a westdipping normal structure (KF, Fig. 3d), with mean slip vector oriented towards the west (Fig. 4). It comprises two splays, which merge towards the south and its vertical throw is at least 500 m (Figs 2 & 6e, f). The northern part of the Kamenitsa Fault controls the eastern margin of the Dara Basin. The east-dipping Levidi Fault is a steep normal fault (LeF, Fig. 3e), with a mean slip vector oriented towards the NE (N358E) (Fig. 6f). Its maximum cumulative throw is not more than 250–300 m, decreasing towards the north. The Kamenitsa and Levidi faults converge towards the North, forming a north –south horst, which can be followed for several kilometres to the north, up to Mt Falkos, separating the Dara basin in the west from the Kandila basin in the east (Figs 2 & 6f).

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Fig. 6. Cross-sections across selected locations in the study area (see Fig. 2 for locations). 1, synrift deposits; 2, Pindos flysch; 3, Pindos carbonates (Cretaceous); 4, Pindos radiolarites and pelites (Upper Jurassic– Lower Cretaceous); 5, Tripolis flysch; 6, Tripolis carbonates; 7, Zarouchla Complex; 8, internal thrust; 9, thrust.

Discussion The geological cross-section of Figure 9 summarizes the results of our geological mapping and structural observations, presented in the preceding section. The section trends NNE– SSW, parallel to the main extension direction across the Corinth Gulf (e.g. Briole et al. 2000; Avallone et al. 2004) and has been selected to cross as many tectonic structures as possible, so as to ensure better control of the overall geometry. The pre-extensional configuration of the nappepile in the area is known to comprise a series of imbricated thrust sheets; it is therefore understandable that the knowledge of pre-extensional nappe geometry is a critical component in our understanding how post-nappe extension has affected the nappe-pile. Thrust surfaces in the Peloponnesus are described as horizontal, gently dipping or curved (e.g. Dercourt 1964; Lekkas 1978; Karotsieris 1981; Thiebault 1982; Bassias 1984). A large number of these tectonic contacts have been studied in

central and southern Peloponnesus (Lekkas 1978; Thiebault 1982; Bassias 1984), where late-orogenic stretching has affected the primary thrust sheets; therefore, many of them can now be described as extensional nappes (Skourtsos et al. 2001, 2004; Skourtsos 2002; Lekkas & Skourtsos 2004). Geophysical studies in northwestern Peloponnesus and mainland Greece – that is, in areas where extension has not affected the initial nappe geometry – show that the thrust sheets are characterized by a flat, long basal thrust that develops within mechanically weak rocks, roughly parallel to the bedding of the unfolded strata and climbs rapidly to higher levels through short ramps (Kamberis et al. 1998; Sotiropoulos et al. 2003). It is therefore reasonable to assume that the suggested nappe geometry in the study area corresponds to an antiformal stack, as seen in Figure 10b. The analysis of kinematic data from compressional structures (Kamberis et al. 1998, 2000; Skourlis & Doutsos 2000; Xypolias & Doutsos 2000) confirm that the cross-section of Figure 9 is nearly perpendicular to the shortening direction of the thrust system

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Fig. 6. (Continued).

and also crosses the nappe pile at an area where the thickness of the nappe column is maximum, on the crest of the antiformal stack. Therefore, the preextensional geometry is represented by a stack of nappes bounded by flat thrust surfaces, as shown in Figure 10c. The section can be distinguished in two parts, a northern, from the Corinth Gulf to Mt Khelmos and a southern one, from Mt Khelmos to Mt Mainalon in the south. The northern part, which corresponds to the onshore shoulder of the Corinth Rift, is characterized by thick outcrops of synrift deposits, bounded WNW– ESE and NW–SE highangle normal faults with northerly dips. The northernmost onshore fault is the Eliki Fault (EF, Figs 2 & 9) while offshore faults have been imaged in seismic reflection and bathymetric surveys studies, such as the Akrata Fault, the throw of which

exceeds 400 m (Stefatos et al. 2002). The cumulative displacement of the Eliki Fault has been estimated at 700–800 m (Ghisetti et al. 2001; Ghisetti & Vezzani 2005). The hanging wall of the Eliki Fault is composed of the Pindos Unit formations, overlain by a lower marl/siltstone sequence and the Akrata fan deposits, which are a lateral continuation of the 800 m thick Vouraikos fan where palynological analyses of the fan deposits gave an age of Lower –Middle Pleistocene (Malartre et al. 2004). The deposition of the fan sediments was controlled by the high-angle Pyrgaki–Mamoussia Fault, which has a throw of 800–1000 m (Malartre et al. 2004). The geometry of the fan delta conglomerates does not indicate southward tilt of the Pyrgaki–Mamoussia Fault block (see Moretti et al. 2003; Lemeille et al. 2004; Malartre et al. 2004), while the block south of the Pyrgaki–Mamoussia

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Fig. 7. View towards the east of the southern flanks of Mt Ziria, at Drossopigi, where the synrift deposits have been tilted by as much as 30– 458. The geometry of the synrift conglomerates (indicated by ball-and-line symbol) suggests that these are growth strata, deposited on the hanging-wall of the south-dipping normal fault that juxtaposes them against the Pindos carbonates (on the left of the photograph). Three successive low-angle normal faults are seen at the right (south), where ‘slivers’ of the Pindos and Tripolis formations are wedged between the Tyros beds and the synrift conglomerates. By restoring the conglomerates to an approximate horizontal geometry, the three low-angle faults become high-angle, north-dipping ones. TB, Tyros Beds; Tr, Tripolis carbonates; Pi, Pindos carbonates; Pl-Pst c, synrift conglomerates.

Fault are now tilted as much as 20– 408 (Flotte´ & Sorel 2001; Rohais et al. 2007; this study). Indeed, further south, the Valimi Fault appears to be a southward rotated fault, a notion corroborated by the geometry of the synrift deposits. The Tsivlos Fault is the northernmost extensional structure with NE–SW strike, while the orientation of its mean slip vector (N708E) differs significantly from the majority of extensional structures in the study area. Furthermore, the fact that it seems to truncate the: Valimi Fault is an indication of ‘out-ofsequence’ fault activity, in the sense that a southern fault is younger than a northern one. West of this location, Causse et al. (2004), using U –Th dating of calcrete cements in fault breccias, also suggests out-of-sequence migration of fault age activity. The Tsivlos Fault also truncates Khelmos Fault, which is, according to Sorel (2000), the surficial expression of the detachment (‘Khelmos detachment’) that controls the opening of the Gulf of Corinth. The fact that the slope breccias that seal the fault dip c. 208S suggests a southward tilt that postdates the Khelmos Fault activity (Flotte´ & Sorel 2001). This is indicative either of footwall rebound due to activity of the faults in the north or to the presence of a north-dipping normal fault further to the south. Intense southward back-tilting is observed on Mt Ziria as indicated by the geometry of the synrift deposits on the southern flanks of the mountain. The restoration of the synrift sediments in their original geometry, results in the Pindos thrust becoming

horizontal, a suggestion also made by Flotte´ & Sorel (2001) for the same thrust on Mt Khelmos. A significant difference between the northern and southern parts is that in the latter the outcrops of synrift deposits are limited to thin outcrops of Pleistocene continental deposits that have filled the basins of Dara, Kandila and Levidi (Figs 2 & 7). As regards the pre-rift formations, the Pindos Unit is the main outcrop, with considerable extent between the mountainous masses of Saitias and Oligyrtos in the north and Mainalon in the south (Fig. 2); Dercourt (1964) describes this area as the ‘Chotoussa syncline’ (Chotoussa being an alternative name for Kandila). This extensive Pindos outcrop is bounded by the south-dipping Pindos thrust in the north and by the Agridi Fault and Orchomenos Fault in the south (Figs 2 & 9). The Pindos nappe is cut by NE–SW and NW –SE normal faults (Penteleion, Gardiki and Lykouria faults), which result in it being found at progressively lower altitudes, from c. 1900 m on Mt Khelmos in the north, to below sea level at the Dara basin. Following the assumption for pre-extensional nappe geometry, these fault blocks and consecutively the thrusts and thrust sheets within them have been subjected to southward tilting, by 25 –408. It can therefore be reasonably assumed that the geometry of the Pindos thrust, which displays an overall southward tilt for the part south of Mt Khelmos, can be indicative of the geometry of the underlying nappe sheets, as well. It is thus suggested that the exposure of the Zarouchla

Fig. 8. (a) Panoramic view (see Fig. 2 for location) of the Agridi Fault, western segment of the North Mainalon Fault Zone, which juxtaposes the Tripolis carbonates against the Pindos formations. Both branches of the Agridi Fault are visible: it is truncated by a younger, steeper normal fault. (b) Close-up view of the striated surface of the Agridi Fault. Tr, Tripolis carbonates; fT, Tripolis flysch; Pi, Pindos carbonates.

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Complex within the Krathis valley, at the northern part of Feneos basin and the southern slopes of Mt Ziria (Figs 2 & 12), can also be attributed to this southward tilt. The nature of the tectonic contact between the Zarouchla Complex and the overlying pre-rift Units has not been completely resolved: Xypolias & Doutsos (2000) and Xypolias & Koukouvelas (2001) describe it as the upper boundary of a topto-the-WNW shear zone, combined with chevron and kink folds with top-to-the-ENE shear sense, attributed to back-thrusting and extrusion of the Phyllites –Quartzites and resulting in the arcuate geometry of the shear zone and the formation of a normal fault on its eastern margin. Recently, Ghizetti & Vezzani (2005) suggest that the exhumation of Zarouchla Complex is attributable to NE – SW Miocene extension and interpret this tectonic window as a NNW– SSE crustal bulging (‘Zarouchla culmination’). Moreover, Papanikolaou & Royden (2007) describe a detachment fault in Feneos and place its activity in the Late Miocene – Early Pliocene, suggesting that this is the northward continuation of an extensional fault on the eastern slopes of Mt Parnon, in southeastern Peloponnesus. However, our geological mapping and structural observations show that: (i) the brittle faults and related structures along the northeastern margin of the Zarouchla Complex (Krathis Fault) exhibit a dominant top-to-north/NNE sense of slip, while the hanging-wall synrift sediments of some of these faults (i.e. Drossopigi and Monodendri Faults) are of Pliocene– Middle Pleistocene age; (ii) the dimensions and cumulative displacement of these faults suggest that they reach deeper than the Phyllites –Quartzites PQ Unit: indeed, geological mapping has shown that the Gardiki Fault does truncate the contact between the Phyllites – Quartzites and the Tyros Beds (Fig. 2); (iii) the Drossopigi Fault, at the southern slopes of Mt Ziria is not an original low-angle fault, but a southwardrotated high-angle structure; (iv) Rohais et al. (2007) show that the older synrift formations that contain clasts derived from the Phyllites –Quartzites Unit are the Lower to Upper Pleistocene alluvial fan and giant Gilbert fan deltas, which means that the exposure of the Phyllites –Quartzites Unit in the earth surface must have taken place around this time period. Based on all the above, it is suggested that the position of a low-angle extensional detachment that may control the formation and evolution of the Corinth Rift should be deeper than the Zarouchla Complex and the exposure of the Phyllites–Quartzites Unit on the earth surface is a late feature, younger than the Late Miocene. If this is true, the effects of the Late Miocene extension seem to be less pronounced in northern than in southern Peloponnesus or Crete. Indeed,

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Fig. 9. Geological cross-section, from the south coast of the Gulf of Corinth to the northern flanks of Mt Mainalon. Lithological symbols are the same as in Figure 2.

Xypolias & Koukouvelas (2001) suggest that no net extension has taken place in the overall system during the extrusion of the Phyllites –Quartzites Unit, while Skourtsos (2002) and Skourtsos et al. (2004) point out that the exhumation of the Phyllites –Quartzites Unit in Southern Peloponnesus took place in the Pliocene based on the fact the older synrift deposits that overlie unconformably the metamorphic rocks are of Late Pliocene age and that the differences in the age of the older synrift sediments between Crete and the Peloponnesus may show a slower exhumation rate towards the NNW, as also indicated recently by Jolivet et al. (2006), who suggest that extension is less efficient in the Peloponnesus than in Crete.

The overall suggested geometry can be seen in the section of Figure 11. The area south of the Khelmos Fault is a large, southward-rotated fault block, bounded in the south by the NW– SW to WNW – ESE North Mainalon Fault Zone, The faults, fault blocks and the amount of southward tilting observed south of Mt Khelmos are believed to be controlled by the North Mainalon Fault Zone, which exhibits significant cumulative displacement, as evidenced by the relative position of the Pindos thrust (Figs 7 & 12). The endorheic basins of Kandila and Levidi have been formed at the relatively southern, more intensely downthrown part of the tilted hanging-wall block, adjacent to the master fault (North Mainalon Fault Zone), while, going north, the deeper thrust

Fig. 10. (a) Schematic representation of the Peloponnesus nappe pile during the Middle Miocene (after Jacobshagen et al. 1978). (b) Schematic diagram of the late-stage thrust configuration of the Hellenide nappes in northwestern Peloponnesus. (c) Derivation of the section undergoing subsequent orogen-parallel extension (redrawn from Ferranti & Oldow 1999).

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Fig. 11. Suggested geometry for North Peloponnesus and the Gulf of Corinth. According to this configuration, the North Mainalon Fault Zone (NMFZ) is the southernmost of the extensional structures related to the Corinth Rift. Assuming a listric or ramp-flat geometry, the NMFZ should flatten at a depth of 6– 8 km, below Mt Khelmos, while its deepest part may link to the postulated detachment zone beneath the Gulf of Corinth with the major normal faults of North Peloponnesus soling onto it. Two major crustal blocks can be distinguished: a southern, now inactive and a northern one, which hosts the present-day seismic activity of the Rift. The distinction between upper and lower plate does not necessarily imply difference in lithology and/or metamorphic grade (as it is common in metamorphic core complexes). Block arrows show the suggested localized uplift driven by the northward propagation of east–west extension.

sheets (Tripolis and Zarouchla Complex) crop out at its relatively uplifted part. The northern boundary of this fault block is the Khelmos Fault and all the northand NE-dipping faults that mark the northern boundary of the Zarouchla Complex. If the North Mainalon Fault Zone has a listric or ramp-flat geometry, analogous to the generally accepted models in various continental extensional fields (Schlische 1991), then it should flatten at a depth of 6–8 km, underneath Mt Khelmos (Fig. 11). The deepest part of this fault may reach further north and merge in the postulated detachment zone below the Gulf of Corinth suggested by seismological and geophysical investigations (Rigo et al. 1996; Taylor et al. 2003; Bernard et al. 2006); the major normal faults in North Peloponnesus, such as the Eliki and Pyrgaki –Mamoussia faults would also sole onto this detachment (Fig. 11). This configuration is comparable with the existence of a lowangle normal fault at depths of 6– 7 km, suggested by Doutsos & Poulimenos (1992). It is also compatible with the hypocentral distributions and focal mechanisms of the earthquake sequences analyzed by Rigo et al. (1996) and Lyon-Caen et al. (2004), while it also allows for the existence of faults with lengths in excess of 10 km and with considerable cumulative displacement. The suggested detachment beneath the Northern Peloponnesus does not correspond to the Phyllite–Quartzites unit of the Zarouchla Complex, but it may lie within a mechanically weak zone of a deeper Unit of the Hellenides. The southern part of the detachment is now truncated by NNE and NNW extensional faults

(Kamenitsa and Levidi faults) (Fig. 2), which are found for at least 6– 7 km within the hanging wall of the North Mainalon Fault Zone, a fact that proves that it is no longer active, at least along its southern (and shallower) segment. The activity of the Kamensita and Levidi faults may be linked to northward propagation of the east–west extension in the southern Peloponnesus, causing further uplift in the central and northern Peloponnesus. This extension may be due to gradual and more localized uplift of the Plattenkalk Unit (metamorphic equivalent of the Ionian Unit), which is characterized by the formation of north– south to NNW –SSE oriented mountain chains, such as Mts Taygetos and Parnon (Skourtsos et al. 2004). Should this be the case, then the mechanical barrier suggested by Ghizetti & Vezzani (2005), composed of metamorphic rocks in northern Peloponnesus, may not be the result of Miocene extension, but the outcome of a much younger phase. In view of the aforementioned observations and suggestions, an alternative scenario for the evolution of the Corinth Rift could be given (Fig. 12), bearing in mind the ambiguity and the insufficiency of data regarding crucial aspects, such as the initial width of the Corinth Rift and the position, type and evolution of the northern margin; the suggested interpretation assumes that this northern boundary is fixed. The evolution of the Corinth Rift may have taken place in two stages (e.g. Ori 1989). In the proposed model, the first rifting phase involved the formation of two large fault blocks: a northern

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Fig. 12. Suggested evolutionary model for the Corinth Rift. 1, Synrift deposits; 2, Parnassos Unit; 3, Pindos U.; 4, Tripolis U. (a) Initial conditions, after nappe emplacement. Parnassos Unit, a Mesozoic carbonate sequence that crops out only north of the Corinth Rift, is the topmost nappe. (b) Early stage of rift evolution and onset of the uplift of North Peloponnesus. The northern fault block is a symmetrical one, bounded from the south by the Khelmos Fault. The sedimentation within the northern block is fed by drainage systems draining the southern, asymmetrical one, bounded from the south by the North Mainalon Fault Zone (NMFZ). (c) As extension increases, the southern block becomes more tilted, together with the Khelmos fault and new, steep hanging-wall faults develop north of it. This stage corresponds to the second phase of the Corinth Rift evolution, where sedimentation in the northern block is characterized by giant alluvial and Gilbert-type fans. Back-tilting and antithetic faulting lead to the formation of endorheic basins on the hanging-wall of the NMFZ. (d) Present-day configuration. The NMFZ is domed beneath Mt Khelmos, with its southern part locked (the NNW and NNE faults that truncate NMFZ cannot be seen, as they run parallel to this section). The active part of the detachment is confined to beneath and north of Mt Khelmos, with the north-dipping high-angle normal faults of North Peloponnesus and the Gulf of Corinth soling onto it. Dotted line represents the actual relief. NMFZ, North Mainalon Fault Zone; KF, Khelmos Fault; TF, Tsivlos F.; VF, Valimi F.; PMF, Pyrgaki– Mamoussia F.; EF, Eliki Fault.

STRUCTURE AND EVOLUTION OF CORINTH RIFT

block, bounded by south-dipping faults located along the present-day northern coast of the Gulf and the Khelmos Fault (Proto-Corinth Gulf, according to Ori 1989); this block must have been a symmetrical one, at least during this early stage. Sedimentation within the Proto-Corinth Gulf was characterized by continental deposits, from the southern block that corresponds to the modern Mts Khelmos, Saitias and Oligyrtos, between the Khelmos Fault (northern boundary) and the North Mainalon Fault Zone. The formation of these two blocks can be attributed to the extension that followed the overall Peloponnesian uplift, which was in turn initiated by the underplating processes along the Hellenic subduction zone (Le Pichon & Angelier 1981; Jolivet et al. 1996; Skourtsos 2002). As uplift of the Northern Peloponnesus continued to take place, displacement on the North Mainalon Fault Zone and the Khelmos Fault increased, forming the required accommodation space for the deposition of the upper sequence of synrift deposits, which now comprised giant alluvial fans and Gilbert-type fan deltas. Further displacement on the North Mainalon Fault led to hangingwall backtilting, the deformation of the Khelmos Fault and the northward migration of fault activity on the Pyrgaki –Mamoussia and Eliki faults. Progressively, the Khelmos Fault became sealed by younger members of the synrift sequence. Meanwhile, the Zarouchla Complex was gradually exposed, as a consequence of the aforementioned processes, feeding sediments of the upper sequence with metamorphic detritus (Rohais et al. 2007). The continued uplift and fault-block back-tilting in the Northern Peloponnesus brought about dramatic changes in the geomorphology, including the reversal of drainage in some of the rivers that originally flowed towards the north (i.e. Seger & Alexander 1993), feeding the Proto-Corinth Gulf. The northward propagation of east –west extension is manifested in the formation of Kamensita and Levidi faualts, which caused the localized uplift of the Northern Peloponnesus, the deactivation of the North Mainalon Fault Zone and the blocking of the southern, up-dip segment of the Corinth Rift detachment, the active segment of which lies north of Mt Khelmos and may correspond to the up-dip, southward prolongation of the postulated detachment zone below the Gulf of Corinth. The horizontal extension derived from the cross section of Figure 12 is at least 20 km, using the Pindos thrust as a marker. This figure could be considerably higher, though, as this estimation does not take into account second- and higher-order faults. The known geodetic measurements across the Gulf of Corinth show that the average extension rate ranges between 6.4 and 14 mm a21 (Billiris et al. 1991; Clarke et al. 1998; Briole et al. 2000).

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Assuming a constant extension rate, this means that the initiation of the Corinth Rift is placed between 3.12 and 1.42 Ma, which is congruent with the age of the older synrift deposits.

Conclusions In this paper an alternative consideration on the structure and evolution of the Corinth Rift is presented, in an attempt to explain some problematic aspects of its geometry, kinematics and seismicity. The existence of a detachment beneath Mt Khelmos can explain the occurrence of faults with length and throw in excess of 10 km and 2 km, respectively, the earthquake activity in North Peloponnesus at depths of 6–15 km and the northward migration of faulting. The uplift of North Peloponnesus is attributed to underplating processes related to the Hellenic subduction zone. The authors are grateful to Mary Ford and Brian Taylor, whose reviews greatly improved the clarity, focus and consistency of the manuscript, while Uwe Ring provided valuable recommendations and assistance. Sincere thanks are also extended to Spyros Lekkas, for his sharing his experience and reviewing the earliest version of the manuscript. We would also like to thank Denis Sorel for field guidance in North Peloponnesus and Konstantinos Soukis and Francis Lemeille for fruitful discussions.

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S KOURTSOS , E., A LEXOPOULOS , A. & L EKKAS , S. 2001. The tectonic structure of the upper plate of the Vlachokerasia metamorphic core complex: implications for the origin of the low-angle normal fault. Bulletin of the Geological Society of Greece, 34, 217–225. S KOURTSOS , E., L EKKAS , S. & A LEXOPOULOS , A. 2004. The structural evolution of Mt Parnon, Peloponnesus, Greece: implications for synconvergence extension in arc regions. In: C HATZIPETROS , A. & P AVLIDES , S. (eds) Proceedings of the 5th International Symposium on Eastern Mediterranean Geology, Thessaloniki, Greece, 189–192. S OREL , D. 2000. A Pleistocene and still-active detachment fault and the origin of the Corinth–Patras rift, Greece. Geology, 28, 83– 86. S OTIROPOULOS , S., K AMBERIS , E., M ARIA , V., T RIANTAPHYLLOU , M. V. & D OUTSOS , T. 2003. Thrust sequences in the central part of the External Hellenides. Geological Magazine, 140, 661 –668. S TEFATOS , A., P APATHEODOROU , G., F ERENTINOS , G., L EEDER , M. & C OLLIER , R. 2002. Seismic reflection imaging of active offshore faults in the Gulf of Corinth: their seismotectonic significance. Basin Research, 14, 487–502. T AYLOR , B., G OODLIFFE , A., W EISS , J., S ACHPAZI , M., H IRN , A., L AIGLE , M. & S TEFATOS , A. 2003. Detachment tectonics in the Gulf of Corinth Rift. Geophysical Research Abstracts, 5, 07222. T HIEBAULT , F. 1982. Evolution ge´odynamique des Helle´nides externes en Pe´loponne`se me´ridional (Gre`ce). The`se d’Etat, Universite´ de Lille. W ESTAWAY , R. 1996. Quaternary elevation change of the Gulf of Corinth in central Greece. Philosophical Transactions of the Royal Society of London, A354, 1125– 1164. W ESTAWAY , R. 1998. Dependence of active normal fault dips on lower-crustal flow regimes. Journal of the Geological Society, London, 155, 223–253. W ESTAWAY , R. 2002. The Quaternary evolution of the Gulf of Corinth, central Greece: coupling between surface processes and flow in the lower continental crust. Tectonophysics, 348, 269–318. X YPOLIAS , P. & D OUTSOS , T. 2000. Kinematics of rock flow in a crustal-scale shear zone: implication for the orogenic evolution of the southwestern Hellenides. Geological Magazine, 137, 81–96. X YPOLIAS , P. & K OUKOUVELAS , I. 2001. Kinematic vorticity and strain rate patterns associated with ductile extrusion in the Chelmos Shear Zone (External Hellenides, Greece). Tectonophysics, 388, 59–77. Z AHRADNI´ K , J., J ANSKY´ , J., S OKOS , E., S ERPETSIDAKI , A., L YON -C AEN , H. & P APADIMITRIOU , P. 2004. Modeling the ML 4.7 mainshock of the February– July 2001 earthquake sequence in Aegion, Greece. Journal of Seismology, 8, 247– 257. Z AMBETAKIS -L EKKAS , A. & K AROTSIERIS , Z. 1986. Le Jurassique supe´rieur de la zone de Tripolitza dans la re´gion de Vitina (Pe´loponne`se Central-Gre`ce). Revue de Paleobiologie, 5, 269– 279. Z ELILIDIS , A. 2000. Drainage evolution in a rifted basin, Corinth graben, Greece. Geomorphology, 35, 69–85.

Timing and nature of formation of the Ios metamorphic core complex, southern Cyclades, Greece STUART N. THOMSON1*, UWE RING2, STEPHANIE BRICHAU3, JOHANNES GLODNY4 & THOMAS M. WILL5 1

Department of Geology and Geophysics, Yale University, New Haven, CT 06511, USA

2

Department of Geological Sciences, Canterbury University, Christchurch 8140, New Zealand 3

School of Earth Sciences, University and Birkbeck College, London WC1E 7HX, UK 4

GeoForschungsZentrum Potsdam, 14473 Potsdam, Germany

5

Geodynamics and Geomaterials Research Group, Department of Geography, Universita¨t Wu¨rzburg, 97074 Wu¨rzburg, Germany *Corresponding author (e-mail: [email protected])

Abstract: We apply low-temperature thermochronology, Rb/Sr geochronology, petrological data, and structural mapping to constrain the timing and kinematics of the Ios metamorphic core complex. Top-to-north extension in the lower plate Headland Shear Zone was active at 18–19 Ma under metamorphic conditions of 475–610 8C and 0.65– 1.1 GPa. The South Cyclades Shear Zone/Ios Detachment Fault (SCSZ/IDF) system shows top-to-south extensional shear active at c. 19 Ma at 380–550 8C, with local top-to-north bands. Extensional shear above the SCSZ/IDF is dominantly top-to-south to top-to-SW. PT estimates from an eclogite boudin constrain Eocene high-pressure metamorphism to 430– 560 8C and 1.21 + 0.42 GPa to 0.66 + 0.37 GPa. Similar low-temperature thermochronometric ages across Ios demonstrate that ductile extensional movement ceased by c. 15 Ma. Exhumation to shallow crustal levels took place between c. 15 and 9 Ma at cooling rates of up to 120 8C Ma21 with a slow down to ,20 8C Ma21 between 12 and 9 Ma, most likely accommodated by extensional slip at rates of c. 3 km Ma21 along the top-to-SW Coastal Fault System. We propose a model of bivergent extension for exhumation of the Ios core complex between 19 and 9 Ma, with Ios forming a secondary antithetic top-to-south to top-to-SW extensional fault system to a more dominant top-to-north Naxos/Paros detachment system.

The Cycladic Islands of the central Aegean Sea, Greece, are part of an extensive region of continental crustal extension related to the rollback of the Hellenic subduction zone (McKenzie 1978; Lister et al. 1984; Buick 1991). Most of the islands are characterized by distinctive Eocene eclogite/ blueschist high-pressure–low temperature metamorphic rocks of the Cycladic blueschist unit (Du¨rr et al. 1978) that have been subsequently exhumed. Miocene extensional deformation under greenschist-facies metamorphism resulted in the formation of a number of distinct metamorphic core complexes (Lister et al. 1984; Avigad & Garfunkel 1991; Buick 1991; Gautier et al. 1993; Avigad et al. 1997). This study focuses on the island of Ios at the southern margin of the Cyclades (Fig. 1), the first recognized metamorphic core complex within the Cyclades extensional province (Lister et al. 1984; Lister & Forster 1996). Here dominantly top-to-the-south (top-S) directed extensional shear has been associated with the crustal

scale low-angle South Cyclades Shear Zone (SCSZ). Subsequent study of other core complexes throughout the Aegean region has since shown that this top-S movement is peculiar to just a few islands at the southern margin of the Cyclades archipelago, including Ios (Vandenberg & Lister 1996; Forster & Lister 1999a), Serifos (Brichau et al. 2009) and Kithnos (Grasemann et al. 2007). Most tectonic models of Aegean extension are instead based on more extensive study of the islands of the central and northern Cyclades dominated by Miocene core complexes with prominent top-N to top-NE extensional shear zones and low-angle detachment faults such as Naxos (Buick 1991; Gautier et al. 1993; Brichau et al. 2006; Seward et al. 2009), Tinos and Syros (Avigad & Garfunkel 1991; Ring et al. 2003a; Brichau et al. 2007), Mykonos (Faure et al. 1991; Lee & Lister 1992; Brichau et al. 2008), Ikaria and Samos (Kumerics et al. 2005), Amorgos (Rosenbaum et al. 2007; Ring et al. 2009), as well as Andros and Evia (Ring et al. 2007).

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 139–167. DOI: 10.1144/SP321.7 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Simplified geological map of the Aegean region (after Brichau et al. 2009) showing major Oligo-Miocene detachment faults annotated with the dominant direction of upper plate slip. The location of Ios is highlighted. The Lycian nappes are Cretaceous to Miocene thrust sheets that lie tectonically above the blueschist rocks of the Cyclades in the western part of the Aegean (Ring et al. 1999).

Several fundamental questions regarding the nature of Miocene crustal extension in the Cyclades currently remain unresolved. These include how and whether the top-N and top-S core complexes, extensional shear zones and low-angle detachment faults relate to each other in time and space (Lister & Forster 1996; Forster & Lister 1999a) and how and to what extent did this extension contribute to the ultimate exhumation of the Eocene HP–LT rocks of the Cyclades blueschist unit. To best answer these questions requires a better understanding of the nature, extent, evolution, and most particularly the timing of the less studied top-S extension of the southern and western Cyclades. Furthermore, such information is critical in testing how and whether large-scale displacement along the low-angle top-S detachment on Ios interacted with the Miocene top-N Cretan detachment system (Fassoulas et al. 1994; Thomson et al. 1998, 1999; Ring et al. 2001a, b) and opening of the Cretan

Sea further south (Lister et al. 1984; Meulenkamp et al. 1988). We address these questions here by presenting an extensive new low-temperature fission track, (U –Th)/He thermochronometric and Rb/Sr geochronological dataset from the island of Ios (the best-studied of the top-S metamorphic core complexes of the Aegean) in conjunction with complementary new detailed petrographic and structural analysis. These data provide the first constraints on the timing, rates, and kinematics of movement of the final unroofing and exhumation of the Ios core complex, as well as robust new constraints on the relative timing of ductile deformation along several of the main extensional shear zones exposed on the island. Interpretation of the Ios detachment system has been complicated by the presence of multiple generations of later detachment faults and, in contrast to most other Aegean core complexes, an almost complete lack of

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unmetamorphosed rocks in the hanging wall of the uppermost detachment and, moreover, major highlevel brittle deformation along any of the low-angle extensional faults on Ios (Lister & Forster 1996; Forster & Lister 1999a).

Regional setting The geology of the Cyclades comprises three main structural units, which are from bottom to top. (1) The Basal unit, as part of the External Hellenides, which is made up of a thick (.1000 m) sequence of marble and also other metasedimentary and volcanic rocks of Permian to Cenozoic age. The Basal unit crops out below the Cycladic blueschist unit in some windows on Evia, Tinos and Samos. (2) The Cycladic blueschist unit consists of a basement nappe of Carboniferous orthogneisses and schists, overlain by a late to post-Carboniferous nappe consisting of metabasic and metasedimentary rocks (Du¨rr et al. 1978). Prior to emplacement, both these nappes were intruded by middle Triassic granitoids. The uppermost nappe of the Cycladic blueschist unit is made up by an ophiolitic me´lange (Ring et al. 1999). The Cycladic blueschist unit experienced at least two episodes of metamorphism during the Cenozoic. The oldest event is an Eocene regional blueschist- to eclogite-facies metamorphism (Andriessen et al. 1979; Schma¨dicke & Will 2003; Tomaschek et al. 2003). During the Miocene greenschist- to locally amphibolite-facies metamorphic conditions were attained (Altherr et al. 1982; Wijbrans & McDougall 1988; Buick & Holland 1989) and I- and S-type granites intruded during the middle and late Miocene. (3) The Upper unit is rarely exposed and mainly consists of a weakly to non-metamorphic composite Cycladic ophiolite nappe. All these structural units are unconformably overlain by sedimentary basins filled with early Miocene and younger sedimentary rocks (Le Pichon & Angelier 1979; Weidmann et al. 1984). Both the sedimentary basins and the granite intrusions are the result of strong Miocene north–south extension that initiated at c. 23 Ma across the entire Aegean Sea region (Ring et al. 2010) and ultimately caused the opening of the Aegean Sea basin.

The geology of Ios The rock outcrop pattern on Ios (Fig. 2) is governed by the dome-like structure of the Ios metamorphic core complex. A structurally lower basement nappe comprises predominantly folded garnet – mica schists surrounding an internal core of orthogneiss derived from a Variscan granite protolith. The upper parts of the basement unit are

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characterized by intense Cenozoic deformation and mylonites associated with a domed extensional ductile shear zone named the South Cyclades Shear Zone (SCSZ) by Lister et al. (1984). The basement rocks and SCSZ are separated by a later arched lowangle normal fault, the Ios Detachment Fault (IDF) sensu Lister & Forster (1996), from a structurally higher nappe of metasedimentary rocks of the Cycladic Blueschist Unit. The upper nappe is dominated by intercalated marbles, pelitic schists with local eclogite lenses. Some of the latter have Triassic protolith ages (Bro¨cker & Pidgeon 2007). The IDF and the Coastal Fault System (CFS) subdivide the nappe pile on Ios into three tectonic units, which we refer to as the lower plate (basement), the middle plate (post-Carboniferous metasedimentary nappe unit) and the upper plate above the CFS. The Andre´ fault (Fig. 2) is interpreted by Forster & Lister (1999a) as a laterally continuous low-angle normal fault within the middle plate that cuts the older IDF, but is itself cut by the CFS.

Metamorphic and geochronological history Past petrological and geochronological studies have revealed two major post-Variscan phases of metamorphism (M1: Eocene HP –LT blueschist facies and M2: Oligo-Miocene greenschist facies) seen in the upper, middle and lower plates (Andriessen et al. 1979, 1987; Henjes-Kunst & Kreuzer 1982; Van der Maar & Jansen 1983; Gru¨tter 1993; Baldwin & Lister 1998; Keay & Lister 2002). M1 represents the classic eclogite- to blueschistfacies metamorphism of the Cyclades Blueschist Unit seen throughout the Cyclades, which occurred during Eocene subduction and accretion at the southern margin of Eurasia (e.g. Okrusch & Bro¨cker 1990; Avigad & Garfunkel 1991). An Eocene age of M1 metamorphism on Ios has been inferred by ages between about 40 and 55 Ma obtained using various isotopic age dating methods from the lower plate basement and rocks of the middle, and upper plates on Ios, as well as other neighbouring islands (e.g. Andriessen et al. 1979, 1987; Henjes-Kunst & Kreuzer 1982; Van der Maar & Jansen 1983; Baldwin & Lister 1998). Based on reanalysis of previous 40Ar/39Ar ages from the region, Forster & Lister (2005) proposed that M1 metamorphism may have been more protracted, with as many as four separate metamorphic events between c. 53 – 49 Ma and 35– 30 Ma. However, which of these hypothetical M1 sub-events affected the rocks of Ios remains unclear, and hence the exact timing of the Eocene HP metamorphism is unresolved. Only a few estimates of M1 pressure– temperature (PT) conditions have been published for Ios. Van der Maar & Jansen (1983) estimated minimum PT conditions of 0.9–1.1 GPa and 350–400 8C with an

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Fig. 2. (a) Simplified geological map of Ios (major faults after Forster & Lister 1999a) showing locations of samples used for petrological, geochronological and thermochronological analysis in this study. Profile A –A0 marked line of cross-section, while profiles B– B0 and C–C0 mark the transects used to plot and project the thermochronometric age– distance relationships shown in Figure 7 (SCSZ/IDF ¼ South Cyclades Shear Zone/Ios Detachment Fault). Note that crosscutting relationships between the Andre´ fault and the CFS are not resolved on the NW coast of Ios (see Forster & Lister 1999a, figs 6 & 7). (b) Schematic cross-section after Forster & Lister (1999a) showing location of samples projected at right-angles onto the section.

upper pressure limit of 1.5 GPa, albeit with no differentiation between middle and lower plate rocks. Gru¨tter (1993) determined values of 1.2–1.3 GPa and 450– 500 8C for the high-pressure metamorphism in the middle plate, which are

currently the best-constrained PT conditions for the high-pressure event on Ios. M2 is an Oligo-Miocene Barrovian-type greenschist-facies metamorphic event that overprints rocks in all three plates on Ios, its effects

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being much more pervasive in the basement rocks of the lower plate. Van der Maar & Jansen (1983) estimated maximum PT conditions of c. 0.5–0.7 GPa with temperatures of 380– 420 8C, a drop of some 0.6 GPa in comparison to the M1 metamorphic conditions, whereas Gru¨tter (1993) reported slightly higher PT values of c. 0.9 GPa at 400 –500 8C for the middle plate. The age of M2 metamorphism is only poorly constrained. K– Ar and Rb/Sr ages (e.g. Andriessen et al. 1979, 1987; Henjes-Kunst & Kreuzer 1982) imply an age of c. 30 –23 Ma. A phengite 40Ar/39Ar plateau age of 20.6 + 0.1 Ma is interpreted by Baldwin & Lister (1998) to date recrystallization during M2 metamorphism. Some younger mid-Miocene ages have also been reported from the lower plate basement rocks on Ios. These include a Rb/Sr phengite-whole rock age of 13.2 + 0.4 Ma from a relic granite within the orthogneiss basement of the lower plate (Henjes-Kunst & Kreuzer 1982). Ages as young as c. 14 Ma are indicated by several potassium feldspar 40Ar/39Ar age spectra from a mylonitic augen gneiss in the upper part of the lower plate on Ios (Baldwin & Lister 1998). In both studies the ages are interpreted as a Miocene thermal resetting related to magmatic activity on neighbouring Naxos Island. However, Baldwin & Lister (1998) point out that the 13 Ma Rb/Sr age of Henjes-Kunst & Kreuzer (1982) may also date fluid activity and recrystallization related to movement along the SCSZ and Ios detachment. More precise dating of the M1 and M2 metamorphic events and related deformation on Ios is complicated by a complex and varied pattern of mixed and/or partially reset post-Variscan, and preM1 and M2 radiometric ages (e.g. Henjes-Kunst & Kreuzer 1982; Andriessen et al. 1987; Baldwin & Lister 1998). Particularly notable is the preservation of pre-Alpine ages on Ios despite the rocks having experienced minimum temperatures of c. 350 8C during Eocene M1 HP–LT metamorphism, and temperatures of over 380 8C during Oligo-Miocene M2 greenschist-facies metamorphism. Similar observations have been made in rocks of the Cyclades blueschist unit elsewhere in the Aegean (e.g. Wijbrans & McDougal 1988; Ring & Layer 2003). However, such high temperatures would be expected to completely reset many of these mineral-isotopic systems. For example, the widely cited experimental diffusion parameters for Ar in Kfeldspar, muscovite and biotite (e.g. Hodges 2003; Harrison & Zeitler 2005) predict almost total thermally activated volume diffusion loss of radiogenic argon, and hence reset ages, at temperatures of c. 200 8C for K –feldspar, 350 8C for muscovite and 325 8C for biotite (90% loss for a hold time of 10 Ma: Reiners & Brandon 2006). Using thermal time –temperature modelling with a similar

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assumption of thermally activated volume diffusion, Baldwin & Lister (1998) proposed that the only way to preserve old ages through such metamorphic temperatures was by having each metamorphic event being induced by a short-lived ‘thermal spike’ where elevated temperatures lasted for less than 1 Ma (or ,0.01 Ma to preserve premetamorphic K –feldspar ages). Such a scenario is, however, very difficult to reconcile with the pervasive M1 and M2 metamorphism, and the time needed to accrue the amounts of strain caused by associated high-temperature deformation. Instead, the preservation of pre-Alpine ages more likely reflects a higher than commonly assumed temperature for diffusion of radiogenic argon in several of these minerals. For example, recent studies have demonstrated that in the absence of fluids or recrystallization, white mica can retain radiogenic argon to temperatures as high as 500– 550 8C (Scaillet et al. 1992; Villa 1998; Di Vincenzo et al. 2001, 2004; Mulch & Costa 2004; Federico et al. 2005; Ring et al. 2007). K –feldspar 40 Ar/39Ar ages too have been shown to be very susceptible to the effects of partial recrystallization caused by geothermal and metamorphic fluids (e.g. Parsons et al. 1999) meaning that complexities in any obtained age spectra (including older ages) relate to the fluid –rock interaction history rather than to any thermal effects. Resetting or partial resetting caused by local recrystallization and deformation rather than by thermal spikes also offers a far more reasonable explanation for the very heterogeneous K –Ar and 40Ar/39Ar ages related to M1 and M2 metamorphism obtained from adjacent samples and even the same individual rock sample in previous studies on Ios (e.g. Henjes-Kunst & Kreuzer 1983; Baldwin & Lister 1998).

Deformation history Current understanding of the deformation history of Ios is primarily based on the recognition by Lister et al. (1984) that the basement rocks were exhumed in the footwall of the major crustal scale SCSZ and later IDF in the Miocene to form a domed metamorphic core complex. Almost all of the subsequent detailed structural analysis and interpretation on Ios has been carried out by the research group of Gordon Lister (e.g. Vandenberg & Lister 1996; Lister & Forster 1996; Baldwin & Lister 1998; Forster & Lister 1999a). These authors distinguish three main phases of rock fabric generation which they subdivide into five major ductile deformation events: a relic pre-Alpine phase (D1), an ‘Alpine’ collisional phase associated with M1 HP –LT metamorphic conditions (D2), and a phase related to Miocene extension under M2 greenschistfacies metamorphic conditions (D3 –D5). Other

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notable structural and metamorphic work on Ios includes detailed mapping in northern Ios by Gru¨tter (1993). The most dominant fabric recognized on the island is a D2 penetrative augen gneiss (mylonitic) foliation and schistosity within the mica schists seen throughout both the lower and middle plates. Association of D2 with glaucophane lineations and lack of kinematic indicators led Vandenberg & Lister (1996) to infer that D2 deformation was largely co-axial and took place during M1 HP–LT metamorphism related to Eocene collision. D3 is characterized by tight to isoclinal folds with locally penetrative axial planar crenulation cleavage during Miocene M2 greenschist-facies metamorphism. D4 deformation re-orients and overprints all previous structures and is dominated by the development of syn- and post-M2 discrete non-coaxial 10–500 m scale shear zones associated with north–south extension and the development of the SCSZ (Vandenberg & Lister 1996; Forster & Lister 1996). Within the SCSZ the upper few hundred metres of the lower plate and the bottom part of the middle plate preserve D4 kinematic indicators with a consistent top-S shear sense, both in the north and south of Ios (Vandenberg & Lister 1996). However, in the structurally lower parts of the SCSZ and at the contact of the lower plate augen gneiss with the garnet –mica schist, overprinting D4 top-N shear zones, including the Headland Shear Zone system (Fig. 2a), have also been recognized (Lister & Forster 1996; Forster & Lister 1999a). D5 is marked by local upright and open north–south-trending folds associated with later structural doming of the island in the late stages of extensional unroofing and metamorphic core complex formation. The late stages of D4 deformation are also marked by at least two generations of discrete lowangle normal fault systems. The most prominent of these is the IDF (Lister & Forster 1996; Forster & Lister 1999a), which accomplished the final juxtaposition of the basement rocks of the lower plate with metasedimentary rocks of the middle plate (Fig. 2). The IDF truncates earlier ductile fabrics of the SCSZ, but has a similar domal form, with two separate surface expressions: one in the north and a less steeply dipping equivalent in southern part of the island. The IDF formed within rather than at the top of the SCSZ. Forster & Lister (1999a) assign a dominant top-S movement sense on both the northern and southern parts of the IDF. In contrast, Gru¨tter (1993) mapped alternating top-S and top-N shear-sense indicators along the northern strand of the IDF, with the top-N kinematic indicators being dominant. Gru¨tter (1993, p. 50) explicitly assigned a top-N sense of shear for the major extensional event in north Ios.

Several other discrete low-angle faults are associated with the IDF system. The most prominent of these being the Andre´ Fault (Forster & Lister 1999a; see Fig. 2), which runs parallel to and also occasionally crosscuts the IDF at a low angle. The Coastal Fault System (CFS) forms another lowangle extension detachment system recognized by Gru¨tter (1993) and Forster & Lister (1999a) and is restricted to limited exposures on the NW coast of the island (Fig. 2). The CFS facilitated the emplacement of the upper plate on Ios. Gru¨tter (1993) ascribed a top-S sense of shear to the CFS. In contrast, Lister & Keay (1996) report east –westtrending stretching lineations associated with the CFS. Forster & Lister (1999a) postulated that the east –west movement operated at a higher crustal level (at lower greenschist-facies conditions) and from that they infer that the CFS is the youngest extensional fault system on Ios. The precise timing, magnitude, rates (and even dominant direction) of movement on each of the detachment fault systems of Ios remain controversial or unknown, as does the timing and nature of final exhumation of the rocks of the Ios core complex from greenschist-facies conditions to the surface. With the exception of rare breccia in the hanging wall to the CFS and a few late east –west striking high-angle normal faults (Fig. 2) no late stage brittle activity has been documented for any of the low-angle normal faults. Furthermore, highlevel upper plate non-metamorphic rocks are either not exposed or not present on the island (Lister & Forster 1996; Forster & Lister 1999a).

New structural work and sampling We mapped kinematic indicators in mylonitic rocks in the Headland Shear Zone, the SCSZ and associated IDF and in the middle plate between the IDF and CFS (Fig. 2). Our major goal was to link the mylonitic structures with P–T estimations and Rb/Sr dating of the mylonites to obtain a better idea of the tectonic significance of the various extensional systems. In this section we describe the shearzone structures starting from the bottom to the top of the sequence.

Headland Shear Zone We mapped and sampled the Headland Shear Zone in a small area southwest of Mylopotas beach. Vandenberg & Lister (1996, fig. 7) have previously described the structures from this shear zone in detail and reported that top-N shearing overprinted earlier top-S extensional shear. The extensional Headland Shear Zone is associated with north – south stretching lineations (Figs 2 & 3a) expressed by elongated quartz–feldspar aggregates (Fig. 4a).

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Fig. 3. Orientation data of mylonitic foliations and stretching lineations related to the various extensional shear zones on Ios. (a) Headland Shear Zone. (b) Northern part of South Cyclades Shear Zone and associated Ios detachment fault. (c) Southern part of SCSZ/IDF. (d) Andre´ Fault. (e) Varvara eclogite boudin and surrounding mylonites. (f ) Coast Fault System.

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Fig. 4. (a) Microphotograph of moderately deformed augen gneiss from Headland Shear Zone (north is to the right). (b) Lower-plate mylonitically sheared garnet–mica schist from South Cyclades Shear Zone at southern tip of Ios Island.

Shear bands in orthogneiss and intercalated garnet-mica schist indicate a top-N sense of shear. We sampled the garnet –mica schist from a lowstrain part of the shear zone for P –T work (Ios 03-15), a neighbouring ultramylonitically sheared dark orthogneiss (Ios 04-1) for constraining the age of mylonitization, and took orthogneiss samples for low-temperature thermochronometry (Ios 04-1, -2, -4).

variance and pronounced disequilibrium of the metamorphic assemblages. We sampled a marble mylonite (Ios 04-6) and a mylonitic quartzite (Ios 04-7) from the middle plate section of the SCSZ for constraining the age of mylonitization and a mylonitic schist (Ios 04-9) for low-temperature thermochronometry of the middle plate.

South Cyclades Shear Zone and Ios Detachment Fault

The middle plate includes a number of boudins several metres in size relatively close to the CFS and in the direct footwall of the CFS that have eclogitic mineral assemblages including omphacite and garnet, overprinted by later blueschist-facies, and in places greenschist-facies assemblages. One of these boudins, the Varvara boudin, shows exceptional preservation of early eclogite-facies minerals and has been previously described in some detail by Forster & Lister (1999b). We sampled the Varvara boudin (samples Ios 03-1, -2, -3, -4) for constraining aspects of its P–T –t evolution, as well as to place some constraints on the role the CFS played in the exhumation of the eclogites. Sample Ios 03-5 also provides the potential to constrain the cooling history in the direct footwall of the Andre´ fault, while the Varvara boudin samples the cooling history in the footwall of the CFS. The structures in the footwall of the Andre´ fault are more phyllonitic in character than the mylonites in the structurally deeper SCSZ/IDF. The phyllonitic zones are characterized by pronounced chloritization, quartz ribbons that are locally cataclasized, and also calcite and epidote-rich zones. The stretching lineation in the phyllonites trends north–south (Fig. 3d) and associated shear sense indicators are again variable; top-S shear bands and asymmetric clasts dominate but top-N shear bands were also found (Fig. 2).

Northern part. Stretching lineations within the SCSZ trend north–south (Figs 2 & 3b) and are associated with top-S kinematic indicators. However, some leucogranite outcrops (e.g. N368450 0300 , E258180 1200 ) show top-N shear bands. We did not find any clear overprinting relationships between the top-S and top-N shear bands. Our finding of alternating top-S and top-N kinematic indicators is in line with Gru¨tter (1993), but somewhat contradicts the uniform top-S shear sense reported by Vandenberg & Lister (1996). One sample from this part of the SCSZ/ IDF system was collected for low-temperature thermochronometry (Ios 03-5, -6). Southern part. The structures in the SCSZ/IDF zone at the southern tip of Ios were described briefly by Vandenberg & Lister (1996). We found a penetrative moderately SW-dipping mylonitic foliation in garnet –mica schist of the lower plate (Fig. 4b) and mica schist, garnet-bearing quartzite and marble of the middle plate (Figs 2 & 3c). An associated stretching lineation plunges down-dip. Kinematic indicators show a consistent top-S sense of shear. Unfortunately, the mylonitic garnet –mica schist in the lower plate was not suitable for P –T and geochronological work because of the high

Middle plate

TIMING AND EXHUMATION OF IOS CORE COMPLEX

The eclogite and blueschist facies rocks of the Varvara boudin have a moderately NW- to westdipping foliation and a north- to NW-trending glaucophane lineation (Fig. 3e). The boudin is surrounded by mylonitic marble, calcareous schist and quartzite, which also show north- to NW-trending stretching lineations expressed by strongly aligned quartz and calcite aggregates (Fig. 3e). According to Vandenberg & Lister (1996) these structures predate mylonitic structures associated with the Headland Shear Zone and SCSZ. The footwall immediately adjacent to the CFS has a similar structural expression to that of the Andre´ fault, with development of a thin (5–10 m) zone of shallowly to moderately west- to NWdipping phyllonite. The trend of the stretching lineations is east– west according to Forster & Lister (1999a) or north–south according to Gru¨tter (1993). Our mapping shows that a strong north– south stretching lineation in the middle plate is progressively deformed and overprinted in the phyllonite zone of the CFS, an observation already made by Forster & Lister (1999a). Stretched quartz –chlorite aggregates and brittle striae in the phyllonite zone and fault zones associated with the phyllonite zone depict variable orientations, which statistically suggest a NE– SW-oriented extension direction (Fig. 3f ). Kinematic indicators show a predominantly top-SW shear sense.

New petrographic analysis We undertook detailed petrographic analysis on four samples from structurally well-defined sites of the Varvara boudin, the Headland Shear Zone and the SCSZ. The goal is to link the metamorphic evolution with the structures and the geochronological data given below and to augment the few published P –T estimates of the various phases of metamorphism on Ios (e.g. Van der Maar & Jansen 1983; Gru¨tter 1993). Sample Ios 03-15 is from a weakly deformed pod in the Headland Shear Zone (Fig. 2). The weakly foliated rock contains garnet, barroisitic amphibole, paragonite, albite, epidote and zoisite. Texturally late chlorite and muscovite formed during retrogression and, locally, replace the outermost garnet domains. In places, the garnet porphyroblasts are strongly embayed and contain inclusions of quartz, epidote and titanite. The garnet composition is dominated by the almandine and grossular components with minor amounts of spessartine in the garnet cores and pyrope at the rims. Amphibole has uniform XMg values [¼ Mg/(Mg þ Fe2þ)] of 0.65– 0.67, the pistacite component [Xpis ¼ Fe3þ/ (Fe3þ þ Al[6])] in epidote is 0.13–0.21, and the paragonite component [XPa ¼ Na/(Na þ K)] in mica is 0.85 –0.93.

147

Sample Ios 04-7 is from the southern part of the SCSZ within 10 m from where Ios 04-6 was collected. It is a mylonitic quartzitic schist characterized by quartz ribbons, which define a strong stretching lineation. Asymmetric albite porphyroclasts occur in the quartz-ribbon matrix and provide a top-S sense of shear. In addition it contains potassium feldspar, chlorite, white mica and small garnet grains. Sample Ios 03-1 is a strongly foliated glaucophane-bearing, garnet– paragonite schist collected near the Varvara boudin in the middle plate above the northern IDF (Fig. 2). The rock contains up to 2.5 mm large (in diameter) poikiloblastic garnet porphyroblasts that occur in a strongly foliated, fine-grained matrix of paragonite, glaucophane and quartz. The foliation wraps around the pre- to syn-kinematic garnet grains. Inclusions in garnet are paragonite, quartz, rutile and minor amounts of epidote and titanite. The syn-kinematic glaucophane grains contain rare inclusions of paragonite and quartz. Texturally late biotite and chlorite are intergrown with each other and replace garnet, glaucophane and paragonite at their respective grain margins. Garnet is an almandine– grossulardominated solid solution with minor amounts of spessartine and pyrope. The glaucophane is almost a pure Na-amphibole with XNa [¼ Na/(Na þ Ca)] close to 1 and a narrow XMg range of 0.30 –0.38. The composition of the Na-mica is close to the paragonite end-member with XPa values of 0.94 –0.97, and epidote has uniform Xpis values of 0.21– 0.22. Sample Ios 03-4, collected at the same locality as sample Ios 03-1, is a massive, fine- to mediumgrained garnet-glaucophanite with glaucophane, epidote and albite as major constituents and minor garnet, blue-green barroisitic amphibole, quartz and rare muscovite and biotite. Elongated glaucophane and epidote grains define a weak planar foliation. From textural evidence the metamorphic peak assemblage is inferred to be represented by glaucophane, epidote, garnet and muscovite; albite, barroisitic amphibole and biotite formed during retrogression. Garnet is an almandine– grossular– spessartine solid solution, with the garnet cores being rich in the almandine and spessartine components, whereas the rim areas are strongly enriched in the grossular component. The glaucophane grains are very close to their end-member composition and have XMg values ranging between 0.45 and 0.56. The white mica is almost pure muscovite and epidote has pistacite values of 0.20– 0.23.

Pressure – temperature analysis Pressure –temperature (PT) conditions of formation of peak metamorphic parageneses were calculated, assuming local equilibria, using the average PT

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S. N. THOMSON ET AL.

method of Powell & Holland (1994) and an updated version (5/2001 data) of the thermodynamic data set of Holland & Powell (1998). Once the endmembers of the minerals in an equilibrium assemblage are distinguished, all the reactions among those end-members can be balanced. With the thermodynamic data available, each reaction can be used to calculate the PT conditions of formation of the assemblage. Average PT estimates can be calculated from a linearly independent set of end-member reactions. For example, five independent reactions can be balanced for the assemblage barroisitic hornblende–epidote–chlorite–plagioclase–chlorite– quartz –H2O. The calculated pressures and temperatures are correlated and the best average PT value is obtained by least-squares techniques, which allows determination of the standard deviations on the results. Additionally, a x 2 test is applied to the average PT result to check the reliability of the estimate. For example, for five independent reactions, the x 2 value should be ,1.6. In all our calculations this test was passed (Table 1), which indicates the reliability of the results and our initial assumption of local equilibrium. The activities of the mineral end-members in the solid solutions were calculated with the program ax (http://www.esc.cam.ac.uk/ astaff/holland/ax.html).

Pressure – temperature estimates The results of the PT estimations are summarized in Table 1. Several mineral assemblages of three domains in the garnet –barroisitic amphibolebearing schist of sample Ios 03-15 from the Headland Shear Zone constrain epidote –amphibolitefacies metamorphic conditions of 506 + 31 8C/ 0.78 + 0.13 GPa, 523 + 43 8C/0.80 + 0.17 GPa and 571 + 36 8C/0.93 + 0.17 GPa. We interpret these results to reflect the prevailing PT conditions during ductile top-N extensional shearing. The PT

estimates are largely similar to those estimated by Gru¨tter (1993) for the M2 greenschist-facies overprint in the middle plate and thus argue for a rather continuous M2 metamorphism across the SCSZ/IDF. Unfortunately the paragenesis of sample Ios 04-7 from the top-S SCSZ did not allow the determination of exact PT data, although the mineral paragenesis indicates that metamorphic temperatures during shearing were above 380–400 8C and probably of the order of 450–550 8C. This crude estimate is in the same order as the temperature results obtained for the sample from the Headland Shear Zone. The garnet–glaucophane schists Ios 03-1 and Ios 03-4 from the Varvara boudin in the middle plate yielded PT conditions of 500 + 61 8C and 1.21 + 0.42 GPa and 490 + 58 8C and 0.66 + 0.37 GPa, respectively. Whereas the temperatures are basically identical, the pressures differ considerably, even though their 2s uncertainties overlap. We consider these results to indicate the PT conditions reached during the Eocene high-pressure event in the middle plate. The metamorphic temperatures for the samples from the Varvara boudin are again similar to those from the extensional shear zones. However, the large pressure uncertainties make it hard to compare the pressure estimates from the Varvara boudin with those obtained from the Headland Shear Zone.

New Rb/Sr geochronology Rationale Our main goal is to provide robust age constraints for mylonitic deformation in the top-N Headland Shear Zone and the top-S SCSZ to see whether there is a true time sequence in the differently displacing extensional shear zone. We also dated a mylonite marble that surrounds the Varvara

Table 1. PT conditions of formation (including 2s standard deviations) Sample, location & unit

Rock type

Ios 03-1 Foliated grt–nam Varvara boudin (MP) schist Ios 03-4 Massive grt Varvara boudin (MP) glaucophanite Ios 03-15 grt –cam schist Mylopotas beach (LP)

P

2s (P)

T

2s (T )

x2

grt–nam–ep–pa – ab – q

12.1

4.2

500

61

1.12 (1.61)

grt–nam–ep–mu – ab – q

6.6

3.7

490

58

1.33(1.61)

grt–cam–ep–pa – q (A1) grt–cam–ep–pa – q (B1) grt–cam–ep–mu – q (B2)

9.3 8.0 7.8

1.7 1.7 1.3

571 523 506

36 43 31

1.15(1.45) 0.85 (1.45) 1.07 (1.45)

Assemblage

The results of the x 2 test are given in the last column. The test is passed if the calculated number is smaller than the cut-off value given in parentheses. In sample Ios 3-15 mineral assemblages in three equilibrium domains (A1, B1 and B2) were analysed. Mineral abbreviations: grt, garnet; nam, Na-amphibole; cam, Ca-amphibole; ep, epidote; pa, paragonite; mu, muscovite; ab, albite; q, quartz. Unit abbreviations: MP, middle plate; LP, lower plate.

TIMING AND EXHUMATION OF IOS CORE COMPLEX

boudin to place some constraints between deformation during the high-pressure stage and the extensional shear zones. Rb/Sr isotopic data from penetratively deformed rocks can be used to date the waning stages of mylonitic deformation, provided that deformation occurred below the temperature range at which thermal diffusional resetting is activated. The Rb/ Sr system of white mica is thermally stable to temperatures of c. 550 8C, but may be fully reset by dynamic recrystallization at lower temperatures (Inger & Cliff 1994; Freeman et al. 1997; Villa 1998). For example, Sr isotopic re-equilibration between white mica and coexisting phases during mylonitization has been shown to occur at temperatures as low as 350 8C (Mu¨ller et al. 1999). Careful study of the correlation between microtextures and isotopic signatures, both by conventional mineral separation techniques (Mu¨ller et al. 1999) and by Rb –Sr microsampling (Mu¨ller et al. 2000; Cliff & Meffan-Main 2003) has shown that complete synkinematic recrystallization in mylonites is usually accompanied by isotopic re-equilibration. In our samples, late increments of deformation and related white mica recrystallization (during Miocene extension-related M2 metamorphism) occurred largely at temperatures below 550 8C, which ensures that Rb/Sr isotopic signatures record deformation and/or greenschist-facies recrystallization. To detect possible Sr-isotopic inhomogeneities resulting from long-term incomplete dynamic recrystallization, from diffusional Sr redistribution and/or from alteration processes, white mica was analysed in several, physically different (in terms of magnetic properties and/or grain size) fractions whenever possible. This approach ensures control on possible presence of unequilibrated, pre-deformational white mica relics (see Mu¨ller et al. 1999).

Methods For Rb/Sr analysis the internal mineral isochron approach was used (Glodny et al. 2002, 2005). Small samples (c. 20– 100 g) were chosen, the assemblages of which could be clearly tied to certain tectonic or metamorphic events. White-mica fractions were ground in ethanol in an agate mortar and then sieved in ethanol to obtain pure, inclusionfree separates. All mineral concentrates were checked and finally purified by hand-picking under a binocular microscope. Rb and Sr concentrations were determined by isotope dilution using mixed 87 Rb – 84Sr spikes. Determinations of Rb and Sr isotope ratios were carried out by thermal ionization mass spectrometry (TIMS) on a VG Sector 54 multicollector instrument (GeoForschungsZentrum Potsdam). Sr was analysed in dynamic multicollection mode. The value obtained for 87Sr/86Sr of

149

NBS standard SRM 987 was 0.710268 + 0.000015 (n ¼ 19). The observed ratios of Rb analyses were corrected for 0.25% per a.m.u. mass fractionation. Total procedural blanks were consistently below 0.15 ng for both Rb and Sr. Because of generally low and highly variable blank values, no blank correction was applied. Isochron parameters were calculated using the Isoplot/Ex program of Ludwig (2003). Standard errors, as derived from replicate analyses of spiked white-mica samples, of 0.005% for 87Sr – 86Sr ratios and of 1.5% for Rb/Sr ratios were applied in isochron age calculations. Individual analytical errors were generally smaller than these values.

Results Headland Shear Zone. Sample Ios 04-1 from an ultramylonitically deformed orthogneiss yields a deformation age of 18 + 1 Ma (Table 2 and Fig. 5a). It is not a perfect isochron, because of some Srisotopic disequilibrium between apatite and feldspar. This disequilibrium is probably related to weathering effects. Nevertheless, due to high 87 Rb/86Sr ratios in the white mica the age is well constrained. It is interpreted on textural grounds as the age for the end of ductile deformation in this sample, within the stability field of garnet at 475–610 8C and 0.65 –1.1 GPa. South Cyclades Shear Zone. Carbonate mylonite sample Ios 04-6 (SCSZ) gives a perfect isochron age of 18.8 + 0.3 Ma, with no correlation between mica grain size and apparent age (Fig. 5b). This indicates that the age dates cessation of ductile deformation in this specific sample. Mylonitic quartzite sample Ios 04-7 is Sr-isotopically not equilibrated (Fig. 5c). The data pattern shows a clear correlation between white mica grain size and apparent ages, with highest apparent age for the biggest grains. This pattern can be best interpreted as pointing to incomplete dynamic recrystallization during deformation meaning that this age represents a maximum age for deformation in the shear zone at the southern tip of Ios. This precludes derivation of a clear-cut deformation age. However, the smallest white mica analysed (160 –125 mm) show the youngest apparent ages (c. 25 Ma) suggesting these grains are more fully recrystallized, and hence provide a maximum age for deformation in the shear zone at the southern tip of Ios. We report the data from sample Ios 04-7 for reasons of completeness, but stress that no clearcut age for any geological event can be derived from these data. They are nevertheless consistent with the information from sample Ios 04-6 collected nearby, that syn-kinematic recrystallization and thus ductile shearing ceased at 18.8 + 0.3 Ma. It appears that

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Table 2. Rb/Sr analytical data Sample no. Analysis no.

Material

Rb (ppm)

Sr (ppm)

IOS04-1 (18 + 1 Ma, MSWD ¼ 116, Sri ¼ 0.71754 + 0.00054) PS1485 wm 355 –250 mm 279 16.4 PS1486 apatite 1.48 445 PS1487 feldspar 19.2 47.9 PS1488 wm 250 –160 mm 280 14.8 PS1489 wm 160 –125 mm 277 18.4

87

Rb/86Sr

87

Sr/86Sr

87

Sr/86Sr 2sm (%)

49.6 0.00962 1.16 54.9 43.7

0.730641 0.717317 0.718087 0.731729 0.731729

0.0022 0.0012 0.0016 0.0016 0.0016

IOS04-6 (18.8 + 0.3 Ma, MSWD ¼ 1.7, Sri ¼ 0.708369 + 0.000035) PS1525 wm . 160 mm 202 34.3 17.0 PS1527 wm , 160 mm 203 40.2 14.62 PS1533 carbonate 0.05 136 0.00100

0.712875 0.712300 0.708368

0.0014 0.0014 0.0014

IOS04-7 (27 + 8 Ma, MSWD ¼ 54) PS1523 wm 355 –250 mm PS1528 wm 250 –160 mm PS1534 wm 160 –125 mm PS1535 feldspar þ quartz

23.7 24.5 22.9 1.81

0.768808 0.768708 0.767500 0.760136

0.0014 0.0014 0.0014 0.0012

0.0258 0.0210 36.0 26.1 40.0

0.709531 0.709504 0.727244 0.721861 0.728850

0.0012 0.0016 0.0036 0.0012 0.0018

366 378 368 14.0

45.0 44.9 46.8 22.4

IOS03-2 (34.5 + 2.5 Ma, MSWD ¼ 14, Sri ¼ 0.70947 + 0.00092) PS1193 carbonate LBRF 0.68 76.3 PS1194 carbonate HBRF 0.54 75.2 PS1195 wm 250 –200 mm 331 27.4 PS1196 wm 160 –80 mm 307 34.0 PS1197 wm 200 –160 mm 360 26.1

An uncertainty of +1.5% (2s) is assigned to Rb/Sr ratios. wm, white mica; m/nm, magnetic/non-magnetic on Frantz magnetic separator, 138 inclination, at electric current as indicated. LBRF, light bromoform fraction, density ,2.85 g cm23; HBRF, heavy bromoform fraction, density .2.85 g cm23. Typical sample weights are 3 –10 mg for apatite, carbonate, feldspar; and 10 –15 mg for white mica phases.

synchronously at 19 –18 Ma ductile deformation along the top-S SCSZ and along the top-N Headland Shear Zone came to an end. Varvara boudin of the middle plate. Carbonate mylonite sample Ios 03-2 yields an age of 34.5 + 2.5 Ma (Fig. 5d). Closer analysis of the data reveals that there is a slight but significant correlation between mica grain sizes and apparent ages. This suggests either a prolonged process of deformation at decreasing temperatures, or incomplete recrystallization during the latest increments of deformation. Despite the uncertainty on the PT data from the Varvara boudin, there are no obvious indications for decreasing temperatures. Therefore incomplete recrystallization during a late stage of deformation during waning highpressure metamorphism is more likely.

Thermochronology Apatite (U –Th)/He dating methodology Dated crystals were handpicked and inspected under a high-powered binocular microscope with crosspolarization to eliminate grains with inclusions. Suitable grains were then measured in two

orientations for later alpha-ejection correction, and loaded either as single or multiple grains into 1 mm Pt tubes. Degassing of He was achieved by heating with a Nd-YAG laser in a high-vacuum laser cell connected to the He extraction and measurement line. Concentration of 4He was determined by spiking with a known volume of 3He and analysing the isotope ratio in a quadrupole mass spectrometer according to the procedure outlined in Reiners et al. (2003). For U, Th and Sm analysis, degassed apatite grains were dissolved in situ from Pt tubes in HNO3 and spiked with a calibrated 229 Th, 233U, and 147Sm solution. U, Th, and Sm concentrations were determined by inductively coupled plasma mass spectrometry (Reiners & Nicolescu 2006). Alpha ejection was corrected using the formula of Farley (2002). Based on the long-term reproducibility of Durango apatite standard analyses at Yale University, an analytical uncertainty of 6% (2s) was applied to the apatite (U– Th)/He age determinations. Radiogenic He accumulated in apatite is lost by diffusion at even lower temperatures than the annealing of fission-tracks in the same mineral. Extrapolation of laboratory He diffusion experiments in apatite to geological time (Farley 2000), supported by evidence from borehole data (House et al. 1999), show that He begins to be

TIMING AND EXHUMATION OF IOS CORE COMPLEX

151

Fig. 5. Rb/Sr isochron plots of samples analysed in this study (wm, white mica; m/nm, magnetic and non-magnetic fraction – see Table 1; HBRF/LBRF, heavy and light bromoform heavy liquid fraction, respectively; MSWD, mean square of the weighted deviates).

measurably lost above about 45 8C, and entirely lost above about 85 8C for a typical grain with radius of 60 + 20 mm (Wolf et al. 1998). This range of temperatures, called the helium partial retention zone, is analogous to the apatite fission track partial annealing zone (e.g. Gallagher et al. 1998; Reiners & Brandon 2006; Wagner & Van den Haute 1992). The closure temperature for He in apatite of typical grain radius (60 + 20 mm) is c. 70 8C at a cooling rate of c. 10 8C Ma21 (Farley 2000).

Fission track analysis methods The methodology follows that outlined in Thomson & Ring (2006). CN5 and IRMM540R glass was used to monitor neutron fluence during irradiation at the Oregon State University Triga Reactor, Corvallis, USA. Zeta calibration factors (Hurford & Green 1983) of 342.5 + 3.8 (CN5 apatite), 352.4 + 12.1 (IRMM540R apatite) and 371.5 + 14.0 (CN5 zircon) were used in age calculation. Central ages

(Galbraith & Laslett 1993), quoted with 1s errors, were calculated using the IUGS recommended Zetacalibration approach of Hurford & Green (1983), which allows for non-Poissonian variation within a population of single-grain ages belonging to an individual sample. Owing to the low spontaneous track counts in several of the samples, ages are also presented with 95% confidence intervals for a binomial parameter (Table 3) (Brandon et al. 1998; Galbraith 2005). For apatite of typical Durango composition (0.4 wt% Cl) experimental and borehole data (Green et al. 1989; Ketcham et al. 1999) show that over geological time tracks begin to anneal at a sufficient rate to be measurable above c. 60 8C, with complete annealing and total resetting of the apatite fission-track age occurring at between 100 8C and 120 8C. This range of temperatures is labeled the apatite fission-track partial annealing zone. For samples that have undergone moderate to fast cooling, a closure temperature of 110 + 10 8C can be reasonably assumed for the most common

152

Table 3. Apatite and zircon fission track data Sample no.

IOS 03-3 IOS 03-4

50 50 90

Location

36.750658N; 25.264668E 36.750658N; 25.264668E 36.750088N; 25.264828E

Mineral

Apatite Apatite Apatite Zircon

IOS 03-5

201

36.751178N; 25.303628E

Apatite Zircon

IOS 03-6

187

36.750328N; 25.297828E

Apatite Zircon

IOS 03-7

240

36.676958N; 25.364238E

Apatite Zircon Zircon Zircon

IOS 04-1

1

36.704138N; 25.287128E

Zircon

No. of crystals

20 17 20 15 17 20 19 6 4 20 20 6 20

Track density (106 tr cm22)

rs (Ns)

ri (Ni)

rd (Nd)

0.0282 (16) 0.0611 (39) 0.0045 (5) 2.362 (283) 0.1437 (127) 1.325 (229) 0.2804 (133) 1.521 (56) 0.5677 (45) 2.090 (360) 2.006 (512) 2.073 (163) 1.774 (414)

0.7906 (449) 0.9783 (624) 0.1174 (130) 7.669 (919) 2.574 (2275) 4.529 (783) 3.550 (1684) 5.404 (199) 9.349 (741) 7.065 (1217) 7.929 (2024) 7.681 (604) 6.096 (1423)

1.375 (4289) 1.365 (4261) 1.048 (3270) 0.2642 (1649) 1.393 (4346) 0.2651 (1655) 1.033 (3223) 0.2661 (1660) 1.384 (4318) 0.2670 (1666) 0.2679 (1672) 0.2688 (1678) 0.2698 (1684)

Age Dispers (Px 2ion) 2.07% (76%) 1.11% (69%) 0.11% (99%) ,0.01% (99%) ,0.01% (99%) ,0.01% (99%) ,0.01% (99%) ,0.01% (99%) ,0.01% (54%) 2.92% (61%) ,0.01% (99%) 8.97% (29%) 0.01% (98%)

Central age (Ma) (+1s)

50% (Median) Age for binomial parameter (95% C.I.)* 8.7

8.4 + 2.1 14.9 14.6 + 2.4 7.8 6.9 + 3.2 15.1 15.1 + 1.2 13.3 13.3 + 1.2 14.4 14.4 + 1.3 14.0 14.0 + 1.3 14.1 13.9 + 2.2 14.6 14.4 + 2.2 14.7 14.7 + 1.1 12.6 12.6 + 0.8 13.5 13.5 + 1.4 14.6 14.6 + 1.0

þ5.0 24.0 þ5.3 24.6 þ8.7 25.6 þ2.1 22.0 þ2.5 22.3 þ2.3 22.1 þ2.6 22.4 þ4.7 23.9 þ4.8 24.2 þ1.8 21.7 þ1.3 21.2 þ2.5 22.3 þ1.7 21.6

Apatite mean track length (mm+1 s.e.) (no. of tracks)

Standard deviation (mm)

–

–

–

–

–

–

–

–

14.27 + 0.24 (17) –

0.95 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

S. N. THOMSON ET AL.

IOS 03-2

Elevation (m)

–

–

0.72

–

–

–

–

14.37 + 0.11 (42) –

–

– 14 Zircon

2 Apatite

14

1 IOS 04-9

36.652028N; 25.368178E

20 Zircon

20 Apatite

IOS 04-4

36.704008N; 25.288308E

4 Zircon 36.704228N; 25.287188E 3 IOS 04-2

Summary of data

Analyses by external detector method using 0.5 for the 4p/2p geometry correction factor. Ages calculated using dosimeter glass: CN5 with zCN5 ¼ 342.5 + 3.8 (apatite). zCN5 ¼ 371.5 + 14.0 (zircon). Px 2 is the probability of obtaining a x 2 value for v degrees of freedom where v ¼ no. of crystals 21. *Binomial Age with 95% confidence limits (see Brandon 1992; Galbraith 2005) applicable where grain counts (Ns) are dominantly low (Ns  5). † IRMM 540R with z540R ¼ 352.4 + 12.1 (apatite).

13.9 13.8 + 1.2

,48.5†

13.6 + 0.9

40.9

13.7

13.0 13.6 + 1.9

13.0 + 1.0†

13.7

14.1 13.9 + 2.2†

,0.01% (98%) ,0.01% (92%) ,0.01% (.99%) ,0.01% (99%) ,0.01% (68%) ,0.01% (99%) 1.917 (5982) 0.2707 (1689) 1.907 (5951) 0.2716 (1695) 1.897 (5920) 0.2725 (1701) 1.141 (1023) 6.917 (259) 9.111 (6454) 6.704 (1749) 0.1122 (14) 6.322 (793) 0.0469 (42) 1.869 (70) 0.3515 (249) 1.813 (473) 0.0008 (1) 1.730 (217) 20 Apatite

153

apatite compositions (e.g. Reiners & Brandon 2006). In zircon, tracks are stable to higher temperatures. For pristine zircon, annealing over geological time begins at about 250 + 20 8C, with total resetting occurring above about 310 + 20 8C (Tagami et al. 1998), although these temperatures are lower in zircons with high accumulated radiation damage (Brandon et al. 1998; Rahn et al. 2004). This translates into a closure temperature for fission tracks in zircon of average radiation damage at moderate to fast cooling rates of c. 250 + 30 8C.

þ4.8 24.2 þ4.0 23.5 þ1.7 21.6 þ1.4 21.4 þ114.1 240.4 þ2.2 22.0

–

TIMING AND EXHUMATION OF IOS CORE COMPLEX

A total of ten zircon fission track (ZFT), nine apatite fission track (AFT) and four (U –Th)/He (AHe) new thermochronometric ages are presented (Tables 3 and 4) from samples taken across the SCSZ/IDF and from the footwall of the CFS (Fig. 5). Samples were collected systematically in a transect from NW to SE across the island (B –B0 in Fig. 2) in an attempt to assess slip rates associated with lowtemperature brittle low-angle faulting (e.g. John & Foster 1993; Brichau et al. 2006). Also included in our analysis are three ZFT ages, two AFT ages and two AHe from Ios samples I2, I8, and I11 previously presented in Brichau (2004). Triplet ZFT, AFT and AHe ages were obtained from six individual samples, while paired ZFT and AFT ages were obtained from a further three samples. Three replicate ZFT ages were also determined from sample Ios 03-7 from the augen gneiss of the footwall basement in the southern part of the island. The ZFT ages range from 13.2 + 1.4 to 15.1 + 1.2 (1s) and all have low age dispersion of less than 10%, with the majority less than 0.1%, and all pass the x 2-test at the 5% level indicating that the ZFT sample ages represent a single population of single grain ages. The AFT ages vary between 6.9 + 3.2 and 14.4 + 2.2 Ma (1s). However, most show poor precision owing to the often very low uranium content (,1–10 ppm), and hence spontaneous fission-track densities of the analysed apatite. Indeed, in sample Ios 04-4 only one track was observed meaning the age can only be constrained as being ,48.5 Ma. The track counts in samples Ios 03-2 and 03-4 were also very low at 16 and 5 tracks, respectively. These lower precision ages are generally not included in the subsequent data analysis. Low track densities also meant that confined AFT length measurement was only possible with two samples (Table 3), both showing long mean track lengths (14.27 –14.37 mm) and low standard deviations (0.72 –0.95 mm) typical for samples that have cooled very quickly through the AFT partial annealing zone from temperatures of c. 110 8C to below c. 60 8C at the time represented by the AFT age (e.g. Green et al. 1989). These

5.48 0.15 9.03 2.86 16.1 2.83 3.25 19.9

0.341 0.053 2.777 0.144

Th (ppm) U (ppm)

Th/ U

0.036 0.052 0.011 0.240

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He (ncc)

154

values are consistent with long mean track length measurements of .14.3 mm presented by Brichau (2004). The four AHe ages (three from the footwall and one from the hanging wall of the IDF; Table 4 and Fig. 5a) vary from 9.3 + 0.6 to 11.4 + 0.8 Ma (Fig. 6a) and match well with lower plate AHe ages presented in Brichau (2004) of 9.5 + 0.4 Ma and 10.5 + 0.9 Ma from samples I2 and I8, respectively (Fig. 2a). AHe analysis was attempted on a further three samples (Ios 03-1, Ios 03-4 and Ios 04-9), but owing to their low apatite uranium (and thorium) content they released insufficient radiogenic helium to be measurable.

45.25 88.5 43.75 76.5 2.00 16.62 2.50 12.23 1 1 1 1 0.715 0.846 0.705 0.825 0.78 0.68 0.59 0.23 11.40 10.58 9.31 9.52 201 187 1 3 25.30362 25.29782 25.28712 25.28718 36.75117 36.75032 36.70413 36.70422 los03-5 los03-6 los04-1 los04-2

Longitude Latitude Sample

Table 4. Apatite (U – Th)/He data

Elevation (m)

Corrected age (Ma)

+2s (Ma)

Ft

No of grains

mass (mg)

mwar (mm)

Low-temperature cooling history The first important observation is that none of the new low temperature thermochronological data show any significant difference in age between the footwall and hanging wall of the SCSZ/IDF system in both northern and southern Ios (Fig. 6a, b), as well between the footwall and hanging wall of the Andre´ fault. The higher 250 + 30 8C closure temperature of the ZFT system implies that little or no later brittle movement has occurred along the SCSZ/IDF and the Andre´ fault since c. 15 Ma. This is in line with the findings of Brix et al. (2002) who demonstrated that where ductile deformation (in the form of quartz recrystallization) was observed (at temperatures of . c. 300 8C) ZFT ages become partially or completely reset. Movement along IDF and SCSZ must thus have been solely accomplished at higher temperatures between the ZFT closure temperature (.c. 250 + 30 8C) and Miocene greenschist-facies metamorphic temperatures. The ZFT age from the Varvara boudin in the footwall of the CFS is also around 15 Ma, whereas the three AFT ages show some spread but are younger with ages between 6.9 + 3.2 and 14.6 + 2.4 Ma (Table 3). Similar ZFT, AFT, AHe ages and long mean AFT lengths indicate that rapid cooling of the rocks on Ios continued between about 15 Ma and 10 Ma following cessation of movement along the IDF and SCSZ within the upper brittle part of the crust (Fig. 7a, b). These new data provide the first precise constraints on the timing, and rates of cooling of the final late-stage (brittle field) unroofing and exhumation of the Ios core complex. The large errors on the ZFT and AFT ages mean that precise cooling rates are difficult to constrain. However, some general trends can be observed. For example, the six samples from which triplet (ZFT, AFT and AHe) ages were obtained (Fig. 6a) are consistent in showing rapid cooling at rates of c. 120 8C Ma21 or more at around 15 –12 Ma between the closure temperatures for ZFT (250 + 30 8C) and AFT (110 + 10 8C), with a

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Fig. 6. Shaded topographic location map using SRTM 90 m resolution DEM data showing (a) apatite fission track and (U–Th)/He ages and (b) zircon fission track ages (ages in grey are from Brichau 2004; all ages given with 1s errors). Outcrop of Ios detachment fault after Forster & Lister (1999a) is shown for reference.

slow down in cooling rates (to about 10 – 20 8C Ma21) to below the closure temperature for AHe (70 + 10 8C) between about 12 and 10 Ma. Obviously rates must have slowed since about 10 Ma as the samples have remained at near surface temperatures below about 70 8C (in the upper few kilometres of the crust) since this time. Rapid cooling rates of .40 –100 8C have been observed in during extensional unroofing of the footwall of metamorphic core complexes, especially at lower temperatures between c. 300 8C and the surface (e.g. Holm & Dokka 1993; Lee 1995; Foster & John 1999; Wells et al. 2000). The time of rapid cooling recorded by the ZFT and AFT ages overlaps in time within error with the Rb/Sr phengite-whole rock age of 13.2 + 0.4 Ma obtained from an igneous relic within the lower plate basement orthogneiss by Henjes-Kunst & Kreuzer (1982), as well as the age minima of c. 14 Ma obtained in several potassium feldspar 40 Ar/39Ar age spectra from mylonitic orthogneiss in the upper part of the lower plate on Ios (Baldwin & Lister 1998).

Age – distance relationships A characteristic signature of the footwall of many extensional detachment faults is that thermochronometric ages show a decrease in age in the direction of slip and hence increasing palaeodepth and preand syn-extension temperature (John & Foster 1993; Foster & John 1999). If the thermochronometric age was zero (i.e. above its closure

temperature) before extension started or sometime thereafter, then the rate of slip and the slip direction of an extensional detachment can be determined from the inverse slope of age in the direction of slip of the hanging wall (e.g. Foster et al. 1993; Ketcham 1996; Foster & John 1999; Thomson & Ring 2006; Brichau et al. 2006). The technique is particularly valuable where kinematic indicators of slip direction in the brittle field are ambiguous or lacking and where hanging-wall rocks are absent, such as on Ios. Alternatively, the lack of an inverse slope, or a more complex pattern of ages with slip direction, is usually indicative of either synkinematic magmatism disturbing the crustal thermal profile (e.g. Lister & Baldwin 1993), coeval rapid erosion (e.g. Fayon et al. 2000; Thomson & Ring 2006), resetting of ages by the influx of synkinematic hot mineralizing fluids (e.g. Morrison & Anderson 1998; Carter et al. 2004), or more than one active detachment, potentially with bivergent directions of slip (e.g. Gessner et al. 2001; Brichau et al. 2009). Here, we have plotted the new thermochronometric data along two age–distance profiles (Fig. 8; location of profiles is given in Fig. 2a): (1) B – B0 (NW– SW) parallel the main top-S slip direction recorded in the footwall to the IDF and SCSZ and (2) C–C0 (NE–SW) parallel to the dominant extension direction we have observed in the phyllonites in the footwall to the CFS (Fig. 3f ) where kinematic indicators show a predominantly topSW shear sense, and hanging-wall breccias and brittle fault striations imply this is the youngest and

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Fig. 7. (a) Time– temperature plots (cooling histories) using low-temperature thermochronological mineral pair and mineral triplet sample ages obtained in this study and by Brichau (2004); (b) Comparison of our data with fission track data from Naxos published by Brichau et al. (2006) and Seward et al. (2009).

highest structural level fault. The sample locations are projected onto the plane of the section at right angles to the direction of the main profile line. For the NE –SW age-distance plots, data from samples in the south part of the island were excluded owing to their distance from the currently exposed segments of the CFS (i.e. samples Ios 03-7 and I11 and those further south; Fig. 2a). The slip rates for each thermochronometer and profile given in Figure 8 were determined using the software Regressþ 2.5.3 (McLaughlin 2006) to fit a line to the data using a y on x weighted least-squares linear regression with distance (x) as the independent variable, and age (y) as the dependent variable with the 1s age uncertainty used for weighting. Estimates of 95% confidence intervals for the slope were computed using a non-parametric bootstrap analysis (BCa method) with 3000 bootstrap samples, again using 1s age uncertainties. The computed slope and uncertainty estimates are calculated in units of Ma km21, which were inverted to give the estimate for slip rate in km Ma21.

In the NW–SE direction (B –B0 transect) both the ZFT and AFT ages (Fig. 8a, b) show no significant trend owing to the relatively low precision of the ages, their similarity, and the relatively short distances between the samples. The best-fit slopes for both thermochronometers are ambiguous, with the ZFT ages showing an inverse age-distance relationship, while the AFT ages show a normal relationship. Furthermore the weighted R-squared correlation coefficient for both is very poor (0.17 and 0.07 for ZFT and AFT slopes, respectively). Perhaps the most meaningful constraint is the minimum slip rates. Within the 2s age uncertainties (the error bars used in Fig. 8), then the data do not confirm either top-S or top-N slip, but instead could support either possibility, whereby the minimum top-N slip rate would have to exceed 3.9 km Ma21 for ZFT and 1.6 km Ma21 for AFT, and minimum top-S slip rates supported by the data are 2.4 km Ma21 for ZFT, and 1.7 km Ma21 for AFT. The NW–SE age–distance plot for the AHe data is also shown. However, the slope on the plot is effectively constrained by only four data points on the distance (y) axis (hence the reasonable, but misleading 0.59 R-squared correlation coefficient). One problem with interpreting the thermochronometric cooling ages as being a result of top-S extensional unroofing are the similar ages of each of the different thermochronometers across the major apparently top-S faults on Ios (SCSZ/IDF and Andre´ fault) implying that little or no later brittle movement occurred along these faults since c. 15 Ma (see previous section). If the data do record cooling caused by extensional unroofing, then cooling must have been caused by a higher level extensional fault operating largely in the brittle field. One obvious candidate for such a fault is the CFS exposed on the NW shores of Ios (Lister & Forster 1996; Forster & Lister 1999a). As described earlier, this fault is assumed to be younger than the IDF and operated at a higher crustal level (Forster & Lister 1999a) and hence movement along this fault is more likely to be recorded by low-temperature thermochronometers. It is dominated by NE–SW stretching lineations (Fig. 3f ) (or east–west according to Lister & Keay 1996; Forster & Lister 1999a) that overprint earlier formed north–south lineations and show kinematic indicators with a predominantly top-SW shear sense. To evaluate whether the thermochronometric data fit a model that assumes approximately top-SW slip along the CDF, we have projected the age data onto an NE–SW oriented transect (C–C0 in Fig. 8d, e & f ). Again no significant or clear trend is seen with the higher temperature ZFT and AFT thermochronometers. The big variation in the best-fit slopes for the ZFT and AFT age –distance

TIMING AND EXHUMATION OF IOS CORE COMPLEX 157

Fig. 8. Low-temperature thermochronometric age– distance plots showing zircon fission track (ZFT), apatite fission track (AFT) and apatite (U– Th)/He ages respectively (with 2s error bars) for samples projected at right angles onto profiles B–B0 and C– C0 of Figure 2. Value in box is slip rate calculated from a weighted least-squares linear regression (see text for details of calculations; C.I., confidence intervals). The negative value is indicative of an inverse relationship (equivalent to top-SE for profile B– B0 , and top-NE for profile C– C0 ). The grey lines are those that pass through all the data (within 2s error limits) and provide the limits for minimum slip rate limits given in grey text.

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relationships (22.3 and 235.4 km Ma21) is a consequence of the low-slopes and very low correlation coefficient, suggesting that little relationship between age and distance above the closure temperature of the AFT system (110 + 10 8C). In contrast, the lowest temperature AHe thermochronometer shows an excellent age– distance relationship (R2 ¼ 0.61) indicative of top-SW slip at a rate of 3.1 (þ2.3/21.0) km Ma21 between about 11 and 9 Ma. If the 2s age errors of the AFT and ZFT ages are considered, then these data too support top-SW movement, with minimum rates .2.4 km Ma21 between 14.6 Ma and 11.8 Ma still compatible with the ZFT data, and top-SW rates of .1.1 km Ma21 between 12.4 Ma and 11 Ma reconcilable with the AFT age data from the footwall of the CFS. Note that dominant top-SW movement of the CFS would require a break-away fault somewhere to the NE of Ios, and a further surface expression of the CFS (with hanging-wall rocks) now below sea level in the Aegean and Cretan Sea to the southwest of Ios and west of Santorini. An alternative explanation for the better age– distance correlation of the AHe data is that these ages record the time when the footwall was being exhumed along a relatively steep part of the main detachment at temperatures below about 100 8C, perhaps as part of a ‘rolling hinge’ (e.g. Buck 1988; Wernicke & Axen 1988; Lee 1995), hence yielding a better defined spatial gradient of AHe ages. The more uniform ZFT and AFT ages, on the other hand, would indicate that at the time represented by these ages and at temperatures of c. 250 + 30 8C to 110 + 10 8C the footwall to the active CFS was relatively flat. This would lead to relatively uniform cooling across the whole footwall in the direction of transport, hence the lack of age–distance relationship seen in these higher closure temperature thermochronometers. Such an interpretation would require a listric geometry for the CFS, with the fault shallowing out at around 100 8C (i.e. at about 3 –5 km for typical upper crustal geothermal gradients of between 20 and 30 8C km21.

Temperature –time relations In Figure 9a and b temperature (T ) –time (t) relations for the top-N Headland Shear Zone and the top-S SCSZ have been plotted to see whether the coeval activity at both shear zones as deduced from the Rb/Sr mylonite dating can be more thoroughly constrained. Our PT work indicates synshearing temperatures of between 400 and 600 8C. We note that previous PT work (Van der Maar & Jansen 1983; Gru¨tter 1993) never reported peak temperature much exceeding 500 8C, which in some respect narrows the large error bars shown

for the high-temperature part of the T –t curves in Figure 9 to about 500 + 50 8C. The data for the Headland Shear Zone are reasonably well constrained (solid curve in Fig. 9a) has been fitted through the data points for sample Ios 04-1) and show slightly slower cooling for the hightemperature shear zone history and an increase in the cooling rate between the ZFT and AFT closure temperatures. However, the relatively large errors bars on the temperatures and also on the age data also allow fitting a straight line through the data points (dashed line in Fig. 8a). Whichever solution is more appropriate cannot be decided from our data alone, although we note that the entire lowtemperature data set argues for only very limited slip along the Headland Shear Zone and also the SCSZ/IDF system after c. 15 Ma. No evidence for brittle movement along the Headland Shear Zone is observed, thus the rapid cooling indicated by largely similar ZFT and AFT ages was most probably caused by unroofing along another detachment above the Headland Shear Zone. The AHe data imply that the CFS is a candidate for such a fault. For the SCSZ of southern Ios we have only two data points (Fig. 9b). As noted above, the ZFT and AFT ages imply that slip on the SCSZ and IDF had all but ceased by c. 15 Ma, requiring that rapid cooling likely continued here to temperatures below that for closure of the AFT system (shown as dashed line of T– t curve in Fig. 9b). According to this treatment of the data, the T–t curves from the top-N Headland and the top-S SCSZ are remarkably similar. In Figure 9c, the metamorphic temperatures obtained for the Varvara boudin and the Rb/Sr age for the directly overlying marble mylonite have been combined with the thermochronometric data from the footwall of the CFS. One problem is the poor AFT ages from samples Ios 03-2 and 03-4. We restrict our discussion here to the well-defined ZFT age of 15.1 + 1.2 Ma from Ios 03-4 and the AFT age of 14.6 + 2.4 from Ios 03-3. This results in a T– t curve in Figure 8c remarkably similar to the ones from the Headland Shear Zone and the south SCSZ. A second problem is the nature of the T–t curve between the Rb/Sr age of 34 Ma and the ZFT age of 15 Ma. We can only speculate about this portion of the curve. If the similarity of the ZFT and AFT ages in the data sets is accepted, then it might be inferred that the cooling history is also largely similar and then the dashed line in Figure 8c might be realistic. This would imply that post c. 15 Ma cooling of the middle plate was largely accomplished by slip on the CFS. However, we realise that a slow and monotonic cooling curve (solid line in Fig. 9c) could also fit the data, in which case the tectonic importance of the CFS would be less.

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159

Fig. 9. Temperature (T )–time (t) plots derived from PT estimates, Rb/Sr dates from time of deformation, and low temperature thermochronological data for (a) the top-N Headland Shear Zone of the lower plate, (b) the top-S South Cyclades Shear Zone and Ios Detachment faults of south Ios and (c) the Varvara boudin in the footwall of the Coastal Fault System in NW Ios (dashed, dotted and solid lines signify alternative T – t paths, see text for details).

Discussion Timing and geometry of the extensional structures on Ios The most robust aspect of our data is that they demonstrate that movement on the higher temperature ductile Headland Shear Zone and SCSZ/IDF extensional structures exposed on Ios had ceased by c. 15 Ma. This is indicated by the commonality of ZFT and AFT ages in the footwall of the Headland Shear Zone, the SCSZ/IDF, and the Andre´ fault. Furthermore, the Rb/Sr data imply that most ductile movement had ceased even earlier: by 19– 18 Ma. Our inference is in line with some of the previous data reported by Henjes-Kunst & Kreuzer (1988) and Baldwin & Lister (1998). This conclusion indicates that for the last increment of slip

on the various ductile extensional faults there is no time sequence as inferred by Forster & Lister (1999a), at least within the geological resolution of our data (c. 1 Ma). AHe ages from the footwall of the higher structural level CFS indicate that top-SW motion along this fault was still active until about 9 Ma. The majority of displacement along the SCSZ/ IDF was accomplished while under ductile conditions (.250 + 30 8C) during and after peak Miocene M2 greenschist-facies conditions of up to 500 8C. Because the PT estimates for the M2 metamorphism are fairly similar across the SCSZ/IDF, the post-M2 vertical separation between the lower and middle plate has to be within the error limits of the pressure estimates (+c. 0.1 GPa) and thus less than c. 7 km. This implies that the SCSZ/IDF did not result in significant exhumation of the

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lower plate. The amount of offset (and hence horizontal extension) represented by the SCSZ/IDF is more difficult to determine as this is dependent on the angle of the fault. For example c. 7 km of vertical offset along a 308 fault would translate to a post-M2 displacement of less than c. 15 km, whereas for a lower angle fault of, say, ,108 could signify maximum displacement of more than 40 km. Furthermore, the PT estimates of van der Maar & Jansen (1983) for the high-pressure M1 metamorphism also suggest no distinct break in PT conditions across the SCSZ/IDF, suggesting there was hardly any exhumation and probably also not much displacement across this contact between M1 and M2. Despite the fact that ductile slip on the Ios faults had ceased by c. 15 Ma, the ZFT, AFT and AHe data require continued rapid cooling in the brittle field. Removal of overburden and exhumation to shallow crust of rocks on Ios was achieved rapidly between c. 15 Ma and 9 Ma as indicated by thermochronometrically determined average cooling rates of up to c. 120 8C Ma21. Such high rates are largely only observed during cooling related to tectonic exhumation in core complexes (e.g. Holm & Dokka 1993; Lee 1995; Foster & John 1999; Wells et al. 2000). But which structure accomplished fast cooling in the brittle crust? This question is closely related to the missing non-metamorphic upper crustal section on Ios. Clearly the data indicate that at least part of this cooling was accommodated by extensional unroofing related to top-SW movement on the CFS between about 15 and 9 Ma. Another candidate would be the Naxos –Paros detachment system for which a top-N displacement of .50 km has been proposed (Brichau et al. 2006). The rocks on Ios, including the exposed extensional faults, would thus represent the footwall rocks to these later detachments. This would require that movement on the top-S ductile detachments had all but ceased before the time of major extension on the CFS and/or Naxos–Paros detachment. The earlier slip history on the Ios extensional faults is constrained by some of our Rb/Sr mylonite ages. The data demonstrate coeval slip at both the top-N Headland Shear Zone and the top-S southern strand of the SCSZ at 19–18 Ma. Furthermore, the lack of subsequent resetting of the Rb/Sr system implies that little further ductile movement occurred after this time. Vandenberg & Lister (1996) argued that top-N shear in the Headland Shear Zone overprinted top-S shear related to the northern strand of the SCSZ. Unfortunately there are no data that indicate pre-19 Ma activity on any of these ductile extensional shear zones. We strongly emphasize that the disequilibria evident in our Rb/Sr data from mylonite sample Ios 04-7 do not allow any

movement on the SCSZ to be constrained before 19 –18 Ma. Overall, our data indicate that the extensional faults and exhumation of the Ios metamorphic core complex occurred between c. 19 Ma and c. 9 Ma. Structural evidence of Vandenberg & Lister (1996) suggests that top-S extensional slip occurred first. After this initial phase our geochronological data are probably best interpreted as being compatible with bivergent extension during overall coaxial north–south stretching as discussed below. Figure 10 shows our interpretation of the geometry of the Ios extensional system evolved through time. Based on the structural data of Vandenberg & Lister (1996) we propose that ductile extensional movement along the SCSZ ceased sometime around 18.8 + 0.3 Ma (Fig. 10a). At about the same time top-N movement in the Headland Shear Zone ceased (Fig. 10b). Given the lack of any recognized major faults between the Headland Shear Zone and SCSZ, their current position is likely close to what is was at about 19 Ma. The next time slice constrained by our data set is at c. 15 to 9 Ma when rocks from all three plates on Ios moved quickly through the closure temperatures for the zircon and apatite fission track system. The age-distance relationships can be best explained as being related to top-SW movement along the higher level CFS, possibly as a listric fault flattening out at about 4 km. However, at least some contribution to cooling at this time by coeval unroofing along the major top-N Naxos –Paros detachment system now exposed to the north of Ios cannot be ruled out. As described in detail above, neither data set unanimously supports monovergent top-S movement as advocated by Vandenberg & Lister (1996) and Forster & Lister (1999a), nor do the data clearly indicate solely top-N extensional faulting. From c. 15 to 9 Ma ongoing extension must have been resolved mostly higher up in the structural pile than that presently exposed on Ios. There is no indication for lower greenschist facies or brittle movement in the Headland Shear Zone. At the SCSZ/IDF there is evidence for greenschistfacies deformation but distinctly brittle movement zones are hardly exposed in this fault system. The Andre´ fault shows discrete lowermost greenschistfacies slip zones, while the CFS while largely greenschist facies in nature, also shows brecciation in the hanging wall and some brittle fault striations (Fig. 3f ) making it the most likely candidate for having accommodated post c. 15 Ma brittle field extensional slip on Ios. Another potential candidate for post c. 15 Ma brittle field cooling record on Ios is the Naxos – Paros extensional system, for which there is ample evidence for slip after 15 Ma (Brichau et al. 2006; Seward et al. 2009). However, the Naxos –Paros

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161

Fig. 10. Tectonic model showing our interpretation of evolution of the geometry of the Ios extensional system through the late Cenozoic. (a) At c. 19 Ma bivergent top-N and top-S extensional shearing during upper greenschist to lower amphibolite facies conditions is active. (b) Ios detachment fault overprints ductile extensional structures between 19 and 15 Ma. At about the same time the Andre´ Fault is active and extensional shearing at the Naxos/Paros extensional system is also active. (c) Movement on Headland Shear Zone and SCSZ/IDF has ceased and the Coastal Fault System is accommodating extension on Ios, in concert with still active Naxos/Paros extensional system. Black arrows are active faults, grey arrows represent inactive faults.

extensional system is top-N displacing; a geometry not fully supported by the age–distance relations of our low-temperature thermochronometric data. Note that we can rule out significant erosion as being the cause of this later cooling, although it would require 8– 10 km of erosion in a very short time span, and furthermore there is no evidence of any significant mid-Miocene sedimentation at this time in the Aegean region (e.g. Mascle & Martin 1990) where the products of such substantial erosion would have to have been deposited.

Regional tectonic implications The time of most rapid cooling recorded by lowtemperature thermochronometers on Ios occurred slightly, but significantly earlier than the main phase of c. 12 –10 Ma cooling record by the same thermochronometers on the neighbouring islands of Naxos and Paros to the north (Fig. 6b, Naxos data from Seward et al. 2009; Brichau et al. 2006). Metamorphic and geochronological data of Buick & Holland (1989), John & Howard (1995)

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and Keay et al. (2001) indicate that movement on the Naxos –Paros extensional system commenced before the peak of amphibolite-facies metamorphism at c. 20 –17 Ma and progressed thereafter. Therefore, our Rb/Sr ages from the bivergent Ios extensional system are roughly contemporaneous with the initiation of the Naxos –Paros extensional system. Nonetheless, slip on the exposed ductile extensional faults on Ios ceased before the main phase of cooling and extensional movement recorded on Naxos. However, Seward et al. (2009) using a larger and more densely sampled thermochronometric data set from Naxos instead interpret the younger cooling ages on Naxos as not being directly related to extensional roofing along the detachment, but rather to post-kinematic cooling related to variability of the geotherm during cooling and more specifically the juxtaposition of the exhuming core with the cold upper plate, or to in situ cooling following the c. 12 Ma syn-kinematic intrusion of the Naxos granite (Keay et al. 2001). This would favour a model whereby the differently vergent detachments on Ios and Naxos actually operated near synchronously, similar to bivergent core complexes seen elsewhere such as the nearby Menderes core complex of western Turkey (Gessner et al. 2001; Ring et al. 2003b) or as recently proposed for the slightly younger (c. 13– 8 Ma) Serifos detachment system to the west of Ios (Brichau et al. 2009). If as discussed above, there is a connection between the Ios extensional system and the Naxos–Paros extensional system, then our model shown in Figure 10c would support the hot core complex case of Gessner et al. (2007). In the numerical simulations of Gessner et al. (2007) hot core complexes form if the rheological contrast between the brittle upper crust and the ductile lower crust is very pronounced, i.e. when the lower crust is extremely weak and hot. The hot core complexes are characterized by large-scale normal faults with great slip and cooling rates, and the development of an overall symmetric geometry expressed by a master fault system and a secondary antithetic extensional fault system caused by symmetric, plumeshaped flow of the hot lower crust. In nature, hot core complexes are expressed by extension-related high-grade, in part migmatitic, domes. Adapting the modelling results to the Aegean extensional province, we propose that Naxos/ Paros and Ios belong to one single hot core complex. The top-N Naxos –Paros extensional system would represent the master fault with the extension-related migmatite dome on Naxos/Paros. This migmatite dome is an exceptional metamorphic feature in the entire Aegean Sea, where extension is otherwise characterized by greenschist-facies mineral

parageneses. We further propose that the Ios extensional system represents part of the antithetic extensional system that delimits the hot core complex to the south. Ios Island would basically mark the largely coaxial hinge zone between the top-N master fault and the antithetic top-S fault. Cooling on Naxos and Ios occurred rapidly with rates exceeding 100 8C Ma21. On a regional scale Naxos/Paros and Ios are situated in the centre of the Aegean extensional province. We speculate that asthenospheric flow associated with subduction-zone retreat caused increased heat input in the centre of the region affected by retreat causing the anomalous thermal conditions on Naxos and Ios. Other major normal faults were active at about 22 –18 Ma in the southern Aegean are the top-N Cretan detachment (Thomson et al. 1998, 1999; Rahl et al. 2009) and the top-N Amorgos detachment (Ring et al. 2009). For Amorgos Ring et al. (2009) discuss a model in which this early movement on both faults was not extensional and that normal faulting was a geometric effect of an extruding wedge of high-pressure rocks with a basal thrust and a normal fault at the top of the extruding wedge. True extensional movement on the Cretan detachment did not commence much before c. 12 Ma as evidenced by the development of extensional graben and associated sedimentation on Crete (Ring et al. 2001b; Seidel et al. 2007; Van Hinsbergen & Meulenkamp 2006). There is growing evidence for top-S normal faulting in the south-western Cyclades (Serifos, Kithnos, Ios). New structural field work on Sifnos shows that the latest phase of extensional deformation is also characterized by top-S movement (Ring et al. in prep.). At least the reasonably wellstudied detachment on Serifos (Grasemann & Petrakakis 2007; Brichau et al. 2009) shows a phase of pre-13 Ma slip under greenschist-facies metamorphism. The Serifos detachment was then domed-up associated with top-S and top-N components of slip. On Serifos, as on Ios, age-distance relationships of low-temperature thermochronometric data show a rather flat trend, which is not diagnostic for monovergent extension. In this regard the Serifos detachment has some similarities with the Ios extensional system. In any event, the data from Ios and Serifos show that extensional faulting at the south-western margin of the Cyclades was complicated and in contrast to simpler monovergent top-N extension in the central and northern Cyclades.

Conclusions In this study we have applied an integrated multidisciplinary approach, including low-temperature

TIMING AND EXHUMATION OF IOS CORE COMPLEX

thermochronology (zircon and apatite fission track analysis and apatite (U –Th)/He dating), Rb/Sr geochronology, inference of PT data, and new structural mapping, to provide robust new constraints on the timing, rates, and kinematics of movement of the final late-stage (brittle field) unroofing and exhumation of the Ios metamorphic core complex, as well as robust new constraints on the relative timing and kinematics of ductile deformation along several of the main extensional shear zones exposed on the island. Petrological and geochronological data reveal that the Headland Shear Zone in the lower plate of Ios was a top-N extensional ductile shear zone active at c. 19–18 Ma under epidote-amphibolitefacies metamorphic conditions of 506 + 31 8C to 571 + 36 8C and 0.78 + 0.13 to 0.93 + 0.17 GPa. The southern part of the South Cyclades Shear Zone and Ios Detachment Fault are associated with a consistent top-S sense of extensional shear observed in the lower and middle plates on both sides of these shear zones active at 19 Ma. Mineral assemblages are consistent with greenschist- to lowermost amphibolite-facies conditions of at least 380 8C up to possibly 450 –550 8C. The northern part of the SCSZ/IDF is associated with top-S extensional shear with local top-N shear bands in its lower parts, although the timing of this movement could not be constrained. The middle plate extensional shear is generally more phyllonitic in nature with dominantly top-S, but some top-N shear bands, especially in the direct hanging wall of the Andre´ Fault. The eclogitic Varvara boudin within the middle plate, in the footwall of the Coastal Fault System, demonstrates a more NE –SW-oriented dominant extension direction and top-SW shear sense. We interpret PT estimates of 500 + 61 8C and 1.21 + 0.42 GPa and 490 + 58 8C and 0.66 + 0.37 GPa from this boudin to indicate conditions reached during the Eocene high-pressure metamorphism. An Rb/Sr age of 34.5 + 2.5 Ma likely represents incomplete recrystallization during the latter stages of high-pressure metamorphism. Similar ZFT, AFT and AHe ages in both footwall and hanging wall of SCSZ/IDF, and Andre´ Fault demonstrate that most movement along these faults had ceased by c. 15–12 Ma, and that the majority of displacement along these faults was accomplished under ductile conditions (.250 + 30 8C) following Miocene greenschist-facies metamorphism. The thermochronometric data also indicate that the exhumation to shallow crustal levels was achieved rapidly between c. 15 Ma and 9 Ma at cooling rates of up to c. 120 8C Ma21: rates characteristic of tectonic exhumation in core complexes. A slow down in cooling rates to c. ,20 8C Ma21 between c. 12 and 10 Ma is

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implied by slightly younger AHe ages. Thermochronometric age–distance relationships are poorly constrained. The best constrained results are compatible with top-SW slip at rates of c. 3 km Ma21 between about 15 and 9 Ma along the CFS, with indications that this fault was consistent with a rolling hinge listric fault flattening out at about 4 km. We propose a model of bivergent top-N and top-S extension for the evolution and exhumation of the Ios core complex between c. 19 and 9 Ma, with Ios forming a secondary (and perhaps earlier) antithetic extensional fault system to a more dominant top-N extension represented by the Naxos/Paros detachment system. S. N. T. wishes to thank Peter Reiners and Stefan Nicolescu for valuable support and technical assistance with (U– Th)/He sample analysis. Funded by the Deutsche Forschungsgemeinschaft (grants Ri 538/16, /18, /23 and Graduiertenkolleg 392), the Brian Mason technical trust of New Zealand, and the College of Science of Canterbury University. Paulina and Yanni Ring helped during the 2004 sampling campaign. This contribution also benefitted considerably from stimulating and helpful input provided by volume co-editor Brian Wernicke, as well as from the comments by reviewer Jamshid Hassanzadeh.

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Timing of the Amorgos detachment system and implications for detachment faulting in the southern Aegean Sea, Greece UWE RING1*, STUART N. THOMSON2 & GIDEON ROSENBAUM3 1

Department of Geological Sciences, University of Canterbury, Christchurch 8140, New Zealand 2

Department of Geosciences, University of Arizona, Tucson, AZ 85721-0077, USA

3

School of Earth Sciences, The University of Queensland, Brisbane, Queensland 4072, Australia *Corresponding author (e-mail: [email protected]) Abstract: We present apatite and zircon fission-track (AFT and ZFT) ages from the Amorgos detachment system in the Aegean Sea, Greece. The Amorgos detachment system consists of a basal and an upper detachment. The lower Amorgos detachment occupies the same tectonic position as the regionally important large-magnitude Cretan detachment and therefore can provide improved constraints on the evolution of the latter. AFT ages from the footwalls of both detachments show that detachment-related cooling occurred in the early Miocene, coeval with an important phase of cooling in the footwall of the Cretan detachment on Crete. We interpret the footwall AFT ages to indicate an early Miocene age of movement on the Amorgos detachments, essentially simultaneously with slip on the Cretan detachment. ZFT ages from rocks above the lower Amorgos detachment are not reset indicating that metamorphic temperatures during the Tertiary Hellenic orogeny did not exceed c. 300 8C significantly. We discuss a model in which top-to-the-north movement on the Cretan/ Amorgos detachment commenced in the early Miocene. Soon after the inception of the Cretan/ Amorgos detachment, top-to-the-south movement on the South Cyclades shear zone deformed the latter and brought the Amorgos detachment into a higher crustal position.

One of the most prominent structures in the Aegean extensional province is the top-to-the-north Cretan detachment, which has been extensively studied on the island of Crete (Fassoulas et al. 1994; Jolivet et al. 1996; Thomson et al. 1998, 1999) (Fig. 1) (we use the term detachment here to refer to a low-angle normal fault). The Cretan detachment exhumed the high-pressure Phyllite –Quartzite Unit in its footwall from 30 –35 km depth and juxtaposed it against the weakly metamorphosed Tripolitza Unit in its hanging wall (Fassoulas et al. 1994). New low-temperature thermochronologic data imply that movement and exhumation on the Cretan detachment involved two phases of accelerated cooling and hence movement and exhumation on the Cretan detachment: (1) an earlier cooling/ exhumation phase at c. 19– 17 Ma as implied by zircon fission-track (ZFT) data (Thomson et al. 1999; Brix et al. 2002) and (2) a later phase at c. 13– 10 Ma as inferred from zircon (U– Th)/He and apatite fission-track (AFT) data (Jeffrey Rahl 2008 pers.comm.). This second phase is coeval with the initial development of extensional graben on Crete (Creutzburg et al. 1977; Meulenkamp et al. 1988; Ring et al. 2001a; Seidel et al. 2007). The first phase is harder to interpret tectonically because it is not associated with the development of extensional graben. At this stage the Cretan detachment has been proposed to have represented

the normal fault at the top of an extrusion wedge (Thomson et al. 1999). An extrusion wedge is defined by a thrust fault at its base and a normal fault at its top. The latter accommodates the upward extrusion of the wedge and does not necessarily result from any horizontal extension. Ring et al. (2007, in press) discussed the subtle differences between a normal and an extensional fault. The term normal fault relates to the relative sense of shear along a shear zone or fault and carries a priori no information as to whether faulting resulted from horizontal lithospheric shortening or extension. An extensional fault is always a normal fault (but not vice versa). It is inherent in the definition for an extensional fault that it resulted from horizontal extension. Ring et al. (2001b) proposed that the displacement on the Cretan detachment may exceed 100 km and that the Cretan detachment underlies much of the southern Aegean region. A testable aspect of this proposition is that the Cretan detachment should be exposed on Amorgos Island, which is made up of the same tectonic units that constitute footwall and hanging wall of the Cretan detachment on Crete (Fytrolakis et al. 1981; Du¨rr 1985; Jolivet et al. 2004). Rosenbaum et al. (2007) mapped two detachments on Amorgos. The lower Amorgos detachment indeed occurs between what is interpreted as an equivalent of the

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 169–178. DOI: 10.1144/SP321.8 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. (a) Tectonic map of the Aegean Sea region showing major tectonic units and the location of the present-day retreating Hellenic subduction zone (modified after Jacobshagen 1986 and Du¨rr et al. 1978). The island of Amorgos is located at the boundary between the Cyclades that are dominated by the Cycladic Blueschist Unit and the External Hellenides, which dominate the geology of the island of Crete. Inset: location map showing Miocene to Recent subduction zones and thrust fronts in the Mediterranean. (b) Schematic north–south cross-section showing the projected trace of the Cretan detachment below Amorgos and offset of the latter by the South Cyclades shear zone.

TIMING OF AMORGOS DETACHMENT

Phyllite– Quartzite Unit in the footwall and the Tripolitza Unit in the hanging wall. In this paper we constrain the timing of the Amorgos detachment system, which, if correlatable with the Cretan detachment, should show evidence for slip at c. 19– 17 Ma. Another important prediction of the Ring et al. (2001b) proposition is that the rocks that correlate with the Phyllite–Quartzite Unit should be characterized by higher pressure –temperature (P –T) conditions on Amorgos than on Crete.

Geology of Amorgos The island of Amorgos (Fig. 2) in the southern Aegean Sea occupies a critical tectonic position between the most deeply exhumed parts of the Hellenide orogen in the central Aegean, the Cycladic Blueschist Belt, and the External Hellenides on Crete (Fig. 1). Structural and metamorphic data provide evidence for two distinct high-pressure units on Amorgos: (1) the Metabasite Unit and (2) the Basal Conglomerate and Marble Unit (Rosenbaum et al. 2007). Above these units rests the Flysch Unit (Fytrolakis et al. 1981; Du¨rr 1985). The Metabasite Unit is characterized by a mineral assemblage of glaucophane, garnet and clinopyroxene, indicating P –T conditions of .13 MPa and 500 –600 8C (Rosenbaum et al. 2007). The Basal Conglomerate Unit yielded P –T conditions of 10 –14 MPa and 300 –450 8C (Rosenbaum et al.

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2007) and the Marble Unit .8 MPa and 300– 400 8C (Theye et al. 1997). The contact between the Metabasite and Basal Conglomerate Unit is interpreted as a top-to-the-northwest detachment fault (Rosenbaum et al. 2007), the lower Amorgos detachment (Fig. 3). Above the lower Amorgos detachment occurs a second detachment fault, the upper Amorgos detachment. The latter probably roots in the basal detachment and cuts up section into the Flysch Unit (Fig. 4). Fytrolakis et al. (1981), Du¨rr (1985) and Rosenbaum et al. (2007) described the various rock units that make up the succession on Amorgos in detail. Relevant here is that the glaucophane – garnet– clinopyroxene-bearing rocks of the Metabasite Unit are the only rocks that are exposed in the footwall of the basal Amorgos detachment and that the actual outcrop of this unit is only a few tens of square metres (Fig. 3). Above the lower Amorgos detachment follows, with a distinct metamorphic break, the Basal Conglomerate Unit consisting of Fe–Mg carpholite bearing metaconglomerates, micaschists and quartzites. (Fe– Mg carpholite is a chain silicate and a low-temperature blueschist-facies index mineral in metapelites.) The rocks immediately above the Amorgos detachment are strongly sheared and fibrous carpholite forms a distinct stretching lineation. The carpholite occurs as needles in a quartz matrix and is commonly synkinematically retrogressed to

Fig. 2. Simplified geological map of Amorgos and its satellite island of Nikouria (modified after Du¨rr 1985, Rosenbaum & Ring 2007 and Rosenbaum et al. 2007). The island of Amorgos is a horst structure in the Aegean extensional province and its morphology is mainly controlled by footwall uplift below the seismically active SE-dipping normal fault offshore the SE coast. The locality of samples AMO 05-27 and AMO 05-28 from the Flysch Unit in the upper plate is shown. Note that the detachments are only shown schematically because of constraints with scale. Inset shows location map in the regional framework and the South Cyclades shear zone.

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Fig. 3. (a) Detailed geological map of the area NW of Katapola (see location in Fig. 2) showing the two detachments and the localities of the analysed samples from the lower and intermediate plate. (b) East– west cross-section through sequence northwest of Katapola (map and cross-section modified after Rosenbaum & Ring 2007 and Rosenbaum et al. 2007).

Fig. 4. Sketch showing tectonic relationships of the basal and upper Amorgos detachments. The upper detachment cuts across various units leaving a largely coherent section from the Basal Conglomerate Unit through the Marble Unit into the Flysch Unit in between the two detachments and possibly above it. Limited displacement on the upper detachment is suggested by similar lithology and metamorphism of the rocks of the Flysch Unit. The positions of the dated samples have been projected into this schematic section.

TIMING OF AMORGOS DETACHMENT

chlorite (Rosenbaum et al. 2007) suggesting that the carpholite-defined stretching lineation reflects a large part of the exhumation history from highpressure P –T conditions. The Marble Unit above the Basal Conglomerate Unit has a thickness of c. 1500 m and is the most dominant rock unit on Amorgos covering large parts of the island (Du¨rr 1985). The massive marble sequence is Triassic to Eocene in age (Du¨rr 1985) and contains several small lenses of Cretaceous metabauxites in the northern part of the island. The metabauxites have the assemblage diaspore, hematite and Fe-carpholite (Minoux et al. 1980) for which Theye et al. (1997) estimated the above given metamorphic conditions of 300 –400 8C and .8 MPa. The top of the stratigraphic sequence on Amorgos is made up of the Flysch Unit, which is both in tectonic contact with its underlying units and also shows conformable contacts with the underlying Marble Unit (Rosenbaum et al. 2007) (Fig. 4). Marble olistoliths up to several hundred metres long occur within the flysch (Fytrolakis et al. 1981; Du¨rr 1985). Large nummulitic foraminifers of middle Eocene age have been reported from some limestones olistoliths (Du¨rr 1985). The presence of the Nummulites suggests that the clasts were derived from the uppermost carbonate sequence (now marbles), indicating that limestone deposition continued into the Eocene and that the flysch is middle Eocene and younger in age. This in turn shows that the metamorphism of the Marble Unit occurred after the middle Eocene. Because the high-pressure P –T conditions of the Basal Conglomerate Unit and the Marble Unit are largely similar and there is no evidence for polymetamorphism, we suggest a similar age of this high-pressure metamorphism. If so, the high-pressure metamorphism in the Basal Conglomerate Unit would also be post-middle Eocene in age. The succession above the lower Amorgos detachment collectively resembles the stratigraphy of the Tripolitza Unit of the external Hellenides (Jolivet et al. 2004). This implies that the Metabasite Unit represents the lowermost structural unit in the central Aegean Sea, which we correlate with the Phyllite– Quartzite unit on Crete (Rosenbaum et al. 2007).

Sampling and methodology To constrain aspects of the cooling and exhumation history and thus the timing of the Amorgos detachment system, samples for fission-track dating were collected from the three tectonic packages (lower, intermediate and upper plate), separated by the two detachments (Fig. 4). Given that footwalls of active detachments cool rapidly, the fission-track

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ages from the lower and intermediate plate should date the Amorgos detachment system directly. However, major constraints are the lithology and the very limited exposure of lower plate rocks on Amorgos. We managed to extract apatite grains from three metabasite samples, two of which only yielded one single datable grain (Table 1). Because the sample localities are within a few metres from each other and the samples were irradiated together, we pooled the fission-track ages from these two samples. From the intermediate plate we collected two samples from the carpholite-bearing shear zone directly above the lower Amorgos detachment and two samples from the Flysch Unit of the upper plate (Figs 2 –4, Table 1). For the same reason given above for the metabasite samples from the lower plate, we also pooled the fission-track ages from the two samples from the intermediate plate.

Fission-track analysis Apatite and zircon crystals were separated, mounted, polished and etched according to the techniques outlined by Thomson & Ring (2006). The samples were analysed applying the external detector method and irradiated at the Oregon State University Triga Reactor, Corvallis, USA. The neutron fluence was monitored using Corning uranium-dosed CN5 glass for both apatite and zircon. Spontaneous and induced FT densities were counted using an Olympus BX51 microscope at 1250-times magnification. Central ages (Galbraith & Laslett 1993), quoted with 1s errors, were calculated using the IUGS recommended Zeta-calibration approach of Hurford & Green (1983), which allows for nonPoissonian variation within a population of singlegrain ages belonging to an individual sample. The x 2 test indicates the probability that all grains counted belong to a single population of ages, a probability of less than 5% is taken as evidence for a significant spread of single grain ages. A spread in individual grain ages can result either from inheritance of detrital grains from mixed source areas, or from differential annealing in grains of different composition by heating within a narrow range of temperatures (Green et al. 1989). Due to the low spontaneous track counts in several of the apatite samples, ages are also presented with 95% confidence intervals for a binomial parameter (Table 1) (Brandon et al. 1998; Galbraith 2005). CN5 zeta calibration factors of 342.5 + 3.8 (apatite) and 371.5 + 14.0 (zircon) were obtained by repeated calibration against a number of internationally agreed age standards according to the recommendations of Hurford (1990). Fission tracks in both apatite and zircon shorten or anneal with increased temperature and duration

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Table 1. Fission-track data Sample no.

Lower plate Metabasite AMO 04-16 AMO 05-6 AMO 05-25

AMO 04-1 AMO 05-22 AMO 05-22 Upper plate Flysch unit AMO 05-27 AMO 05-28 AMO 05-28

Mineral

No. of crystals

36.838038N; 25.855928E 36.838038N; 25.855928E 36.838038N; 25.855928E

Apatite

12

Apatite

1

Apatite

1

36.838008N; 25.855838E 36.838008N; 25.855838E 36.838868N; 25.855888E 36.838868N; 25.855888E

Apatite

5

Zircon

8

Apatite

6

Zircon

8

Zircon

20

Apatite

3

Zircon

10

36.831978N; 25.861398E 36.822198N; 25.900588E 36.822198N; 25.900588E

Track density (106 tr cm22)

Age dispersion (Px 2)

Central age (Ma) (+1s)

50% (median) age for binomial parameter (95% C.I.)*

rs (Ns)

ri (Ni)

rd (Nd)

0.1001 (38) 0.1202 (3) 0.2289 (6)

1.212 (460) 2.003 (50) 3.739 (14)

1.249 (3896) 1.857 (5796) 1.838 (5734)

0.01% (94%) n/a

18.3 + 3.2

18.7

19.6 + 11.7

24.2

n/a

19.8 + 8.4

22.1

0.0796 (7) 11.12 (666) 0.1051 (16) 15.73 (638)

1.159 (102) 3.505 (210) 2.535 (386) 4.807 (195)

1.867 (5827) 2.762 (1724) 1.847 (5765) 2.772 (1730)

,0.01% (95%) ,0.01% (92%) 33% (27%) ,0.01% (96%)

22.5 + 8.8

24.8

123.3 215.9

14.1

18.1 26.4

9.122 (2169) 0.0572 (3) 21.57 (2383)

3.415 (812) 0.7249 (38) 4.979 (550)

2.781 (1736) 1.828 (5703) 2.790 (1741)

11.4% (13.7%) ,0.01% (80%) 10.3% (9.5%)

31.3

148.3 226.3

*Binomial age with 95% confidence limits (see Brandon 1992; Galbraith 2005) applicable where grain counts (Ns) are dominantly low (Ns  5). Analyses by external detector method using 0.5 for the 4p/2p geometry correction factor. Ages calculated using dosimeter glass: IRMM 540R with z540R ¼ 352.4 + 12.1 (apatite); IRMM 541 with z541 ¼ 121.7 + 3.2 (zircon). Px 2 is the probability of obtaining a x 2 value for v degrees of freedom where v ¼ no. of crystals 21.

16.9 25.9 136.3 220.2 122.5 215.0

160.7 + 14.6 15.2 + 4.5 166.3 + 15.5

134.0 + 9.0 25.4 + 15.3 219.0 + 16.2

U. RING ET AL.

Intermediate plate Carpholite conglomerate AMO 04-1

Location

TIMING OF AMORGOS DETACHMENT

of heating. For apatite of typical Durango composition (0.4 wt% Cl) experimental and borehole data (Green et al. 1989; Ketcham et al. 1999) show that over geological time fission tracks begin to anneal at a sufficient rate to be measurable above c. 60 8C, with complete annealing and total resetting of the apatite fission track age occurring between 100 and 120 8C. This range of temperatures is usually labelled the apatite fission-track partial annealing zone. For samples that have undergone moderate to fast cooling, a value of 100 + 20 8C is regarded as the closure temperature of fission tracks in apatite. In zircon, tracks are stable to higher temperatures. For pristine zircon grains, annealing over geologic time begins at 250 + 20 8C, with total resetting occurring above 310 + 20 8C (Tagami et al. 1998). This translates into a closure temperature for fission tracks in zircon at moderate to fast cooling rates of 280 + 30 8C, which correlates well with the brittle– ductile transition in silicic rocks (Brix et al. 2002).

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Interpretation of fission-track results AFT ages Our AFT ages involve large uncertainties and are thus not well constrained. However, they appear to be important for constraining aspects of the timing of the Amorgos detachment system. The AFT ages for rocks of the lower and intermediate plate are largely similar and show that detachment-related cooling below 100 + 20 8C occurred in the early Miocene. The AFT ages also strongly suggest that at least the final slip on the basal and upper detachment occurred simultaneously indicating that both detachments are indeed related to each other. Alternatively, the data could also mean that the cooling was related only to unroofing of the upper detachment, because the data do not rule out the lower detachment being older. However, the tectonic as well as stratigraphic relationships between the Marble and the Flysch Units (Fig. 4) imply only limited displacement on the upper detachment, which renders the alternative interpretation unlikely.

Fission-track data Lower plate The only relatively well-defined apatite fissiontrack (AFT) age from the footwall of the Amorgos detachment is from sample AMO 04-16. The sample yielded 12 apatite grains, which provide an age of 18.3 + 3.2 Ma (1s errors) (Table 1). The other two samples yielded similar but poorly defined early Miocene ages (Table 1). The pooled age of samples AMO 05-6 and AMO 05-25 is 19.8 + 6.8 Ma. The weighted mean age (i.e. just the ages, weighted using the error) for all three samples from the lower plate is 18.6 + 2.9 Ma.

Intermediate plate The two samples (AMO 04-1 and AMO 05-22) from the carpholite-bearing shear zone yielded largely similar early Miocene AFT ages than the samples from the lower plate (Table 1). The pooled AFT age from both samples is 16.9 + 4.0 Ma and the weighted mean age is 14.7 + 3.2 Ma. The two zircon fission-track (ZFT) analyses yielded largely identical middle Jurassic ages.

Upper plate The only AFT age from the Flysch Unit (AMO 05-28, Table 1) is again poorly defined but appears to be older than the two samples from the carpholite-bearing shear zone. The two ZFT ages from the upper plate yielded ages ranging from 219 to 134 Ma (Table 1), which are distinctly older than the deposition age of the flysch.

ZFT ages The ZFT ages from the intermediate and upper plate are distinctly older than the Tertiary metamorphic overprint that affected the rocks of these units. The fact that the ZFT ages have not been reset indicates that metamorphic temperatures did not exceed 310 + 20 8C (approximate upper limit of closure temperature of fission tracks in zircon, e.g. Brix et al. 2002) for any significant period of time that would be needed to equilibrate a metamorphic mineral assemblage. We are therefore in a position to better constrain the metamorphic temperatures given by Theye et al. (1997) and Rosenbaum et al. (2007) for the rocks of the Basal Conglomerate Unit and the Marble Unit to c. 300–330 8C. This also constrains the maximum pressure of the carpholitebearing rocks more precisely to 11–12 MPa (see pseudosection in fig. 11 of Rosenbaum et al. 2007).

Implications for detachment faulting in the southern Aegean There are a number of similarities between the Cretan detachment and the Amorgos detachment system. Both structures moved in the early Miocene at c. 19 Ma and normal movement was not associated with the development of extensional graben above the detachments. This suggests that normal movement was not necessarily a result of regional extension but instead due to the development of an extrusion wedge above the subduction thrust as shown for Crete by Thomson et al. (1999) (Fig. 5a). Furthermore, the rocks in the

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Fig. 5. Interpretative sketches illustrating the proposed tectonic evolution of the Cretan and Amorgos detachments in the early Miocene. (a) Development of an extrusion wedge bounded by the Hellenic subduction thrust at the base and the Cretan/Amorgos detachment at the top. This interpretation implies that the high-pressure metamorphism of the tectonic units making up the lower and intermediate plate on Amorgos is largely coeval with high-pressure metamorphism of the Phyllite– Quartzite Unit at c. 25 Ma on Crete (Seidel et al. 1982; Jolivet et al. 1996). Note that boxes refer to P–T conditions in the Phyllite–Quartzite Unit only. (b) Retreat of the Hellenic slab and development of the South Cyclades shear zone. The latter brings the deep Amorgos section of the Cretan detachment into high crustal levels. This interpretation would explain why the Tripolitza Unit on Amorgos has experienced a higher metamorphic pressure than its Cretan equivalent.

lower plate of the Amorgos detachment system (Metabasite Unit which we correlate with the Phyllite–Quartzite Unit) have distinctly higher P–T conditions than the Phyllite– Quartzite Unit on eastern Crete (.13 MPa, 500 –600 8C v. 8– 10 MPa, 300 –400 8C; Crete P–T data from Theye et al. 1992). This would be in line with deeper exhumation of the lower plate in Amorgos because it had a more internal position below the top-tothe-north normal detachment (Fig. 5a). Nonetheless, there is one aspect that makes a straightforward correlation of the Cretan and Amorgos detachments problematic. The lower-plate rocks of the Cretan detachment cooled below c. 280 8C in the early Miocene, whereas the lower and intermediate plates of the Amorgos detachment system already reached much higher crustal levels at this time as indicated by their cooling through the c. 100 8C isotherm. One way out of this problem would be top-to-the-south movement on the South Cyclades shear zone, which is exposed on southern Ios Island (Fig. 1b, inset Fig. 2, Fig. 5b), immediately after the inception of the Cretan/Amorgos detachment. Combined geochronological, structural and metamorphic work by Thomson et al. (2009) have revealed that the South Cyclades shear zone indeed moved at c. 18 Ma and could have thus brought the Amorogos detachment system rapidly into a higher crustal position soon after it formed. Top-S movement on the South Cyclades shear zone explains the different cooling levels of the Amorgos detachment system compared to the Cretan detachment, which is in the hanging wall of the South Cyclades shear zone (Fig. 5b).

The geometry in Figure 5b implies that the South Cyclades shear zone is an extensional structure that controlled the opening of the Cretan Sea basin in its hanging wall (Meulenkamp et al. 1988; Ring et al. 2001a). We propose that this extension was triggered by a new retreat phase of the Hellenic slab, which ultimately resulted in the switchover of the Cretan detachment from a normal fault bounding the extrusion wedge to an extensional detachment and the development of the extensional graben on Crete in the middle Miocene. Funded by the Deutsche Forschungsgemeinschaft, the Minerva postdoctoral fellowship programme of the Max Planck Gesellschaft and the Brian Mason technical trust of New Zealand. We thank the referees for helpful comments.

References B RANDON , M. T. 1992. Decomposition of fission–track grain–age distributions. American Journal of Science, 292, 535– 564. B RANDON , M. T., R ODEN -T ICE , M. K. & G ARVER , J. I. 1998. Late Cenozoic exhumation of the Cascadia accretionary wedge in the Olympic Mountains, northwest Washington State. Geological Society of America Bulletin, 110, 985–1009. B RIX , M. R., S TO¨ CKHERT , B., S EIDEL , E., T HEYE , T., T HOMSON , S. N. & K U¨ STER , M. 2002. Thermobarometric data from a fossil zircon partial annealing zone in high pressure–low temperature rocks of eastern and central Crete, Greece. Tectonophysics, 349, 309–326. C REUTZBURG , N., D ROOGER , C. W., M EULENKAMP , J. E., P APASTAMATIOU , J., S ANNEMANN , W., S EIDEL ,

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bleus. Comptes Rendus de l’Academie des Sciences, Paris, 291, 745– 748. R ING , U., B RACHERT , T. & F ASSOULAS , C. 2001a. Middle Miocene graben development in Crete and its possible relation to large-scale detachment faults in the southern Aegean. Terra Nova, 13, 297–304. R ING , U., L AYER , P. W. & R EISCHMANN , T. 2001b. Miocene high-pressure metamorphism in the Cyclades and Crete, Aegean Sea, Greece: Evidence for largemagnitude displacement on the Cretan detachment. Geology, 29, 395–398. R ING , U., W ILL , T., G LODNY , J. ET AL . 2007. Early exhumation of high-pressure rocks in extrusion wedges: The Cycladic blueschist unit in the eastern Aegean, Greece and Turkey. Tectonics, 26, TC2001, doi: 10.1029/2005TC001872. R ING , U., G LODNY , J., W ILL , T. M. & T HOMSON , S. N. 2009. The retreating Hellenic subduction system: High-pressure metamorphism, exhumation, normal faulting and large-scale extension. Annual Review of Earth and Planetary Sciences, 38, (doi:10.1146/ annurev.earth.050708.170910). R OSENBAUM , G. & R ING , U. 2007. Structure and metamorphism of Amorgos: a field excursion. In: L ISTER , G., F ORSTER , M. & R ING , U. (eds) Inside the Aegean Metamorphic Core Complexes. Journal of the Virtual Explorer, Electronic Edition, ISSN 1441-8142, Volume 27, Paper 7. R OSENBAUM , G., R ING , U. & K U¨ HN , A. 2007. Tectonometamorphic evolution of high-pressure rocks from the island of Amorgos (Central Aegean, Greece). Journal of Geological Society, London, 164, 425–438. S EIDEL , E., K REUZER , H. & H ARRE , W. 1982. A Late Oligocene/Early Miocene high temperature belt in the External Hellenides. Geologisches Jahrbuch, E23, 165– 206. S EIDEL , M., S EIDEL , E. & S TO¨ CKHERT , B. 2007. Tectono-sedimentary evolution of lower to middle Miocene half-graben basins related to an extensional detachment fault (western Crete, Greece). Terra Nova, 19, 39– 47. T HEYE , T., S EIDEL , E. & V IDAL , O. 1992. Carpholite, sudoite, and chloritoid in low-grade high-pressure metapelites from Crete and the Peleponnese, Greece. European Journal of Mineralogy, 4, 487–507. T HEYE , T., C HOPIN , C., G REVEL , K. D. & O CKENGA , E. 1997. The assemblage diaspore þ quartz in metamorphic rocks; a petrological, experimental and thermodynamic study. Journal of Metamorphic Geology, 15, 17–28. T AGAMI , T., G ALBRAITH , R. F., Y AMADA , R. & L ASLETT , G. M. 1998. Revised annealing kinetics of fission tracks in zircon and geological implications. In: V AN D EN H AUTE , P. & D E C ORTE , F. (eds) Advances in Fission-Track Geochronology. Kluwer Academic Publishers, 99–112. T HOMSON , S. N. & R ING , U. 2006. Thermochronologic evaluation of post-collision extension in the Anatolide Orogen, western Turkey. Tectonics, 25, TC3005, doi: 10.1029/2005 TC001833. T HOMSON , S. N., S TO¨ CKHERT , B. & B RIX , M. A. 1998. Thermochronology of the high-pressure metamorphic rocks of Crete, Greece: Implications for the speed of tectonic processes, Geology, 26, 259–262.

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T HOMSON , S. N., B RICHAU , S., R ING , U., G LODNY , J. & W ILL , T. M. 2009. Timing and nature of formation of the Ios metamorphic core complex, southern Cyclades, Greece. In: R ING , U. & W ERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 141– 170.

Cenozoic tectonic evolution of Naxos Island through a multi-faceted approach of fission-track analysis DIANE SEWARD1*, OLIVIER VANDERHAEGHE2, LUC SIEBENALLER2, STUART THOMSON3,6, CHRISTIAN HIBSCH2, ANATOL ZINGG1, PATRICK HOLZNER1, UWE RING4 & STEPHANIE DUCHEˆNE5 1

Geology Institute, ETH Zurich, 8092 Zurich, Switzerland

2

G2R, Nancy-Universite´, CNRS, Boulevard des Aiguillettes B.P. 239 F-54506 Vandœuvre le`s Nancy, France

3

Department of Geology & Geophysics, Yale University, New Haven, CT 06511, USA

4

Department of Geological Sciences, University of Canterbury, Christchurch 8140, New Zealand 5

CRPG, Nancy-Universite´, CNRS, 15 rue Notre-Dame-des-Pauvres B.P. 20 F-54501 Vandœuvre le`s Nancy, France

6

Present address: Department of Geosciences, University of Arizona, Tucson, AZ 85721, USA *Corresponding author (e-mail: [email protected]) Abstract: New zircon and apatite fission-track ages obtained on samples from all lithotectonic units exposed on Naxos Island are presented. Zircon ages of the exhumed metamorphic rocks range from 25.2 to 9.3 Ma and from 13.0 to 6.4 Ma for apatite. Zircon track-length analysis distinguishes partial overprinting of an earlier event (M1) in the south. Northwards no overprint is seen and the ages there represent rapid exhumation since c. 12 Ma. Both zircon and apatite ages are slightly older toward the north of the island probably due to variation of the geotherm in the proximity of the fault. Zircon fission-track ages of the granodiorite range from 13.7 to 12.2 Ma are statistically identical to previously determined U– Pb ages. Apatite fission-track ages however, yield a younging trend from south to north from 12.9 to 9.0 Ma. This could be due to differential depth of emplacement and/or to differential exhumation during tectonic unroofing by a top-to-the north detachment. Fission-track ages on detrital grains in Lower Miocene sediments indicate a source not identified within the present outcropping rocks of Naxos. Ages on boulders and grains in the Middle to Upper Miocene sediments point to rapid erosion until about 8.5–7 Ma.

The application of thermochronology to geological processes relies on the concept of a closure temperature (Dodson 1973) at which the ‘thermochronometric clock’ is activated. Low-temperature thermochronology in particular has been used to reconstruct the cooling history of metamorphic rocks during their exhumation from temperatures of about 300 8C to those at the Earth’s surface (e.g. Fitzgerald et al. 1991; Holm & Dokka 1993; Lorencak et al. 2001). In most such studies, thermochronometric data have been used to decipher the cooling history often interpreted in terms of exhumation rates and tectonics. In order to bring the rocks to the surface exhumation is generally achieved by some form of denudation. Two endmember cases may be considered: one that assumes exhumation is controlled by surface erosion, the other that attributes exhumation to tectonic denudation. Simple erosion-controlled exhumation may sometimes be confirmed by a correlation between

ages and altitude, assuming that the isotherms were not influenced by topographic variations between valleys and crests (e.g. Grasemann & Mancktelow 1993; Braun et al. 2006). Alternatively, interpretation in terms of tectonic-controlled exhumation may be identified by an age trend relating to structures accommodating the exhumation such as lowangle detachments (e.g. Fitzgerald et al. 1991; Holm & Dokka 1993) or high-angle faults (Lorencak et al. 2001). For low-angle detachments, some authors have considered a steady-state thermal regime, which yields a fixed geometric relationship between isotherms and the detachment fault, and thus steady state migration of material points of the footwall of the detachment across the isotherms (Brichau et al. 2006). Others have considered the time evolution of the isotherms perturbed by the activation of the detachment (Grasemann & Mancktelow 1993). In general, isotherms are displaced on each side of the fault and an estimation

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 179–196. DOI: 10.1144/SP321.9 0305-8719/09/$15.00 # The Geological Society of London 2009.

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of the geothermal gradient, and hence exhumation rates, is difficult without modelling procedures (e.g. Ehlers et al. 2001). Besides the geothermal gradient and its evolution, another challenge is to distinguish the record of simple monotonic cooling from high to low temperature, from the one that has undergone some cooling, followed by reheating before a final phase of cooling. In the case of fissiontrack analyses, track-length measurements may yield this sort of information about the different thermal pathways followed during exhumation (e.g. Ketcham 2005). This is generally applied to apatite analysis but a qualitative estimation of cooling history can be made also from zircon length distributions. Further information on the possible earlier phases of the cooling history of a rock is archived in the eroded sediments of the related depositional basins. Thermochronology of the basinal material has proved to be a very satisfactory method linking sedimentary rocks to their potential source ultimately leading to information concerning the thermotectonic events of the source (e.g. Ruiz et al. 2004). In Naxos, several previous thermochronometric studies have been conducted on the various exposed lithologies. Pioneer K –Ar and Rb –Sr studies have provided some constraints on the timing of metamorphism and the emplacement age of the plutonic rocks (Andriessen et al. 1979). More detailed argon thermochronology on white micas documents an age distribution (Wijbrans & McDougall 1988) that was correlated to the metamorphic isograds (Jansen 1973; Jansen & Schuiling 1976) and was interpreted to reflect partial resetting of the argon isotopic system (Wijbrans & McDougall 1986). Metamorphic core complex models for the tectonic evolution of Naxos attribute the exhumation of metamorphic rocks to a single event (Lister et al. 1984; Buick 1991a; Gautier & Brun 1994). Later, John & Howard (1995) and Brichau et al. (2006), using data based on 40Ar/39Ar, K– Ar, fission-track and (U –Th)/He methods, interpreted the distribution of ages with respect to a single low-angle detachment, in terms of differential cooling owing to progressive tectonic denudation. This reasoning yielded estimates of slip rates on the detachment. In this study, we present new zircon and apatite fission-track data for various lithologies representing the different lithotectonic units exposed on Naxos Island, namely metamorphic rocks, magmatic rocks (the granodiorite pluton) and sediments deposited into Cenozoic extensional basins. These new data are discussed with reference to the various models presented above and allow refinement of the thermal and thus the tectonic history of Naxos. More specifically, zircon and apatite fission-track data on the metamorphic rocks provide constraints to

evaluate the relative impact of reheating of metamorphic rocks and allow the reconstruction of a more complex cooling/exhumation history than previously proposed. Zircon and apatite fissiontrack data from the granodiorite pluton confirm its emplacement age but also document the temperature of the host rock and the probable role of tectonically controlled exhumation on its differential cooling. Zircon and apatite fission-track analyses on detrital grains sampled at different stratigraphic levels of the sedimentary sequence yield constraints on the provenance of the sediments, the cooling history of the source, and the thermal history of the sediments.

Geological context and previous work The geology of Naxos Island in the context of the Aegean geodynamics The island of Naxos is situated in the southern Aegean Sea and is part of the Attic –Cycladic Massif (Fig. 1). The geodynamic evolution of this part of the Mediterranean is associated with Cenozoic convergence between Africa and Eurasia and by subduction of the Aegean slab since that time. Subsequent rollback of the subducting slab, of the order of 1000 km (Fyktikas et al. 1984), from the Rhodope region of Bulgaria –northern Greece to the Hellenic Trench south of Crete, has controlled the geodynamic evolution of the Aegean region over approximately the last 40 Ma. The Attic Cycladic Massif forms a belt of metamorphic rocks that first record an Eocene high-pressure– low-temperature (HP –LT) phase that was later overprinted by a medium-pressure –mediumtemperature (MP –MT) (Jansen & Schuiling 1976; Buick & Holland 1989; Avigad 1998; Ducheˆne et al. 2006). This metamorphic evolution was followed by decompression and cooling owing to exhumation during Miocene extension (Lister et al. 1984; Gautier et al. 1990, 1993; Urai et al. 1990). The erosional products of the metamorphic rocks were deposited into Cenozoic sedimentary basins (Jansen 1973; Sanchez-Gomez et al. 2002). Naxos, the largest island of the Cyclades, lies between the Greek and Turkish mainlands, to the north of the Hellenic subduction zone. It contains one of the most complete structural cross sections of the Cyclades whereby pre-Mesozoic rocks and Cenozoic granitoids have been juxtaposed against Tertiary sediments or weakly metamorphosed rocks along a low-angle detachment (Jansen 1973; Buick 1991b; Gautier et al. 1993). Geological studies of Naxos began at the end of the nineteenth century but were dominantly focused on the marble deposits. At the beginning of the twentieth century Papavasiliou (1909) produced the first geological map of

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Fig. 1. Regional tectonic setting of the island of Naxos in the Aegean Sea.

the island, which was followed by the map of Jansen (1973). With the detachment model of Lister et al. (1984) new research was intensified and the tectonic evolution of the island caused much interest. Three tectono-metamorphic units are recognized (Fig. 2a). The upper unit comprises low-grade metamorphic rocks, an ophiolitic melange and Cenozoic

sedimentary rocks (Jansen 1973). This upper unit is dissected by high-angle faults and cataclasites rooting into a cataclastic detachment. Below the detachment, the middle unit is composed of metamorphic rocks with a sequence of alternating layers of dominantly marble and schist with some amphibolite and minor mafic and ultramafic boudins

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Fig. 2. Geological map of Naxos with fission-track sample sites (a) and ages (b). The line A– A0 marks the position of section onto which ages were plotted in Figure 3.

(Jansen & Schuiling 1976). The lower unit consists of migmatites dominated by diatexites with marble rafts that are exposed in the core of an elliptical structural dome (Vanderhaeghe 2004) (Fig. 2). The long axis of the dome is parallel to the NNE– SSW-trending stretching and mineral lineation of the metamorphic rocks. The dome pattern is outlined and emphasized by the pattern of the regional composite foliation and by the pattern of metamorphic isograds (Fig. 2a) (Jansen & Schuiling 1976; Buick & Holland 1989). It is noteworthy that the isograds are slightly discordant to the main transposition foliation with a general steeper dip toward the south. East –west brittle faults (Fig. 2) with both lateral and normal components cut across all lithological units (Siebenaller 2008). These faults possibly record the initial phases of a horst-like structure controlling the shape of the island and probably the trend of the Cyclades.

Petrological and geochronological constraints The metamorphic zoning of the middle and lower units is the result of two successive metamorphic events. Firstly, the HP –LT Eocene metamorphism

(M1) is attested by relics of blueschist facies minerals in the metamorphic rocks of the south of the island (Avigad & Garfunkel 1991). Argon thermochronology documents ages of about 45 + 5 Ma on micas associated with this metamorphism (Andriessen et al. 1979; Wijbrans & McDougall 1986, 1988). In the higher grade metamorphic rocks, no relics of a HP–LT metamorphism have been identified but metamorphic rims of zircon yield U –Pb ages ranging from 57 to 41 Ma interpreted to reflect zircon growth assisted by the circulation of prograde fluids during M1 (Martin et al. 2006, 2008). This HP –LT (M1) phase was later overprinted by a widespread Barrovian type (MP – MT) metamorphism during the Miocene (M2) grading from greenschist facies to amphibolite facies conditions reaching partial melting conditions as attested by the migmatites of the dome (670 8C and 5–7 kbar) (Jansen & Schuiling 1976; Buick & Holland 1989; Ducheˆne et al. 2006; Martin et al. 2006, 2008). The timing of the peak of this M2 metamorphism was first estimated to have been between 20 and 15 Ma based on argon thermochronology (Wijbrans & McDougall 1988) consistent with new Rb/Sr data (Ducheˆne et al. 2006). External

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

overgrowth rims of the zircons in the core migmatites have ages ranging from 19 to 14 Ma (Keay 1998; Keay et al. 2001; Martin et al. 2006). In the structurally higher metamorphic rocks, the age of MP –MT metamorphism is constrained to between 27 and 19.9 Ma by 40Ar/39Ar dating on phengite (Wjibrans & McDougall 1988). Two types of models have been proposed to account for this large range of ages and for the transition from M1 to M2. Buick & Holland (1989) invoke a progressive temperature increase during exhumation on the basis of thermobarometric data obtained on samples from different parts of the island. Alternatively, Wijbrans & McDougall (1988) argue for partial thermal resetting of the HP– LT paragenesis after their exhumation. This is further constrained by more recent Rb/Sr data correlated to metamorphic minerals such as micas and garnets (Ducheˆne et al. 2006). However, the lower bound for the temperature at which the HP –LT unit cooled before the M2 event is still ill-constrained. At about 12 Ma (Wijbrans & McDougall 1988; Keay et al. 2001), the western part of the island was intruded by a granodiorite (Fig. 2). This granodiorite is thought to be part of the late Miocene magmatic arc of the southward retreating Hellenic subduction zone and intruded synkinematically into the footwall of the Naxos extensional fault system (Buick 1991a, b). Jansen & Schuiling (1976) and Buick (1988) documented that the granodiorite was emplaced at 2 –3 kbar pressure at c. 700 8C. Gautier et al. (1993) refined the emplacement pressure to 1.5–2 kbar. The granodiorite has an intrusive contact with the marble and schist layers of the middle unit in the south where the metamorphic rocks had already cooled to temperatures of less than approximately 350 8C by this time as evidenced by Eocene 40Ar/39Ar ages on hornblende and biotite (Wijbrans & McDougall 1988; Andriessen et al. 1979, 1987). In contrast, it is in tectonic contact with the sedimentary sequence of the upper unit in the North. The sedimentary successions of the upper unit in the hanging wall of the detachment indicate that the marine to continental deposits record the erosion of the underlying ophiolitic nappes as well as lithologies that do not crop out on Naxos today. Early Miocene subsidence is marked with the change from reef deposits to deeper marine turbidites and an increasing number of pebbles from the marbles and schists (Vanderhaeghe et al. 2007). The marine deposits, turbidites, containing Aquitanian to early Burdigalian faunas (Angelier et al. 1978; c. 23–19 Ma) are well exposed in the centre of the island in a fault-bounded block between the metamorphic core and the granodiorite. Kuhlemann et al. (2004) using multiple methods, including fission-track analysis on detrital zircons, concluded

183

that these sediments were shed from an uplifting mountain region to the south of Naxos. The fissiontrack age spectrum ranged from 30 to 500 Ma. No high-grade pebbles from the metamorphic core complex were identified. However, rounded clasts of granodiorite and metamorphic rocks do occur in the upper fresh water conglomerates. This unit was originally assigned a late Pliocene age (Roesler 1972; Jansen 1973), but van Bo¨ger (1983) placed an upper Serravallian to Tortonian age (c. 12.5 – 7 Ma) on them. Other deposits of similar age include travertine and lacustrine sediments. Additionally, kilometre-scale landslides and olistostromic deposits overlie the sedimentary units between the granodiorite and the metamorphics rocks (Fig. 2) (Jansen 1973). These attest to large-scale erosion in this region and have been proposed to signify evidence for related strong seismic events recorded elsewhere on the island by pseudodtachylytes (Vanderhaeghe et al. 2007). Several old erosion surfaces that have no signs of tilting have been identified (Heijl et al. 2003) at 650, 500 and 300–400 m. A lower welldeveloped erosion surface, the Tragea plain, is also recognized in the western part of the island at a mean height of 230 m (Riedl 1982a, b; Heijl et al. 2003). Heijl et al. (2003) suggested that such palaeo-relief features have a maximum age of 8 Ma. The previous geochronological data sets discussed above support the structural evidence of exhumation of the footwall of Naxos to the SSW relative to the hanging wall along the base of the Naxos detachment fault (e.g. Andriessen et al. 1979, 1987; Lister et al. 1984; Wijbrans & McDougall 1986, 1988). Two of these studies were based on Rb/Sr, K –Ar or 40Ar/39Ar analyses representing temperatures of closure down to about 300 8C. Using this data, John & Howard (1995) calculated ductile slip rates for the detachment at 5 –8 mm a21, i.e. for rocks structurally below c. 330 8C at approximately 16 Ma. Lower temperature thermochronometric studies (300 to c. 40 8C) have been presented by Altherr et al. (1979), Heijl et al. (2003) and Brichau et al. (2006). Altherr et al. (1979) published one apatite fission-track age of c. 8 Ma for the granodiorite. Heijl et al. (2003) reported six apatite fission-track ages of between 13.4 +7.0 and 7.8 +1.6 (2s) Ma across the granodiorite and the metamorphic lithologies. They concluded that rapid cooling due to crustal extension took place between 12 and 8 Ma at a rate of 50–130 8C Ma21. No variability across the island was taken into account. We have not incorporated this data into our results because the ages were determined in a different manner to those in the present study [the population method and absolute fluence approach (Gleadow 1981) was used]. Brichau et al. (2006) presented zircon fission-track ages ranging between 11.8 + 0.8 and

184

D. SEWARD ET AL.

9.7 +0.8 (2s) and apatite fission-track ages from 11.2 + 1.6 to 8.2 +1.2 Ma. (U– Th)/He ages range from 10.4 + 0.4 to 9.2 + 0.3 and 10.7+1.0 to 8.9 + 0.6 Ma for zircon and apatite respectively. These data confirm the rapid cooling determined by Heijl et al. (2003) but more importantly they were used to determine slip rates of 6.4 (þ6.8/ 22.2) mm a21 to 13.2 (þ9.4/24.5) mm a21 over the temperature range from c. 300 8C to c. 40 8C for the brittle phase of the detachment. One additional single apatite (U –Th)/He age of 10.9 Ma was determined on one of our samples, NAX3 (Vermeesch et al. 2007).

New fission track data Methods All samples (except those of Thomson) were treated according to the procedure described in Seward (1989) whereby the apatite was etched in 7% HNO3 at 20 8C for 50 seconds. This etching condition yields equivalent lengths to that of the 5M HNO3 etch used by Thomson (Seward et al. 2000). Zircons (DS, PH, AZ) were etched in the eutectic melt at 210 8C for times ranging from 20 to 50 hours and those of Thomson at 225 8C. Samples were irradiated at the ANSTO facility, Australia and at the Oregon State University facility (the latter with an asterix in Tables 1 & 2). All ages were calibrated using the zeta method (Hurford & Green 1983). The central age (Galbraith & Laslett 1993) is reported with a 2 sigma error.

Results All results are presented in Tables 1 and 2 and Figure 2. Metamorphic rocks. The zircon fission-track ages (Table 1) within the metamorphic rocks young northwards from 25.2 Ma in the south to 11.9 Ma in the north (Figs 2b & 3). Ages projected onto a 0258 line, (line A –A0 Fig. 2b, direction of mineral and stretching lineation), reveal a marginally curved geographical distribution with oldest ages in the south, the youngest ages over the migmatite (the median part of the island), and a very slight increase northwards towards the detachment (Fig. 3). In order to shed light on the trend in zircon fission-track ages, the lengths of the confined horizontal zircon tracks were measured (Fig. 4). Such analysis is usually applied as a routine part of apatite fission-track analysis whereby the mean lengths and their associated statistics are used to reveal details of the thermal history of samples using modelling procedures (e.g. Ketcham 2005). Likewise, track lengths of zircons may yield

information about the rate of cooling and hidden events contained within an age. For rapidly cooled zircons the mean expected track length is c. 10.3 mm (Yamada et al. 1995). The older ages, in the south, have both shorter lengths and broader standard deviations relative to those farther to the north (Fig. 4) with a hint of bimodality increasing southwards. Apatite fission-track ages from the metamorphic rocks range from 13.0 + 5.4 Ma to 6.4 + 1.8 Ma (Table 2 and Fig. 2b). The large error bars are a result of the generally low uranium content of the samples. Measurement of track lengths was not carried out as the number of confined tracks was in general very low due again to young ages with the associated low uranium content. Despite these limitations, the following patterns are identified. Geographically (section A– A0 ), the apatite fissiontrack ages display an antiformal pattern with younger ages again occurring in the median part of the island (Fig. 3). The difference in ages from the median region to the north, toward the detachment, is slightly greater than that seen in the zircons. The youngest ages occur at the northern termination of the migmatites, the highest grade metamorphics. Granodiorite. The zircon fission-track ages for the granodiorite (Fig. 2 and Table 1) range from 13.7 + 2.2 to 12.2 + 1.4 Ma, all overlapping at the 2s level and with no geographical trend (Figs 2 & 5). Apatite fission-track ages range from 12.9 + 4.4 in the south to 9.0 + 2.6 Ma in the north, (Table 2 and Fig. 5) with overlapping between each neighbouring pair but with a trend in absolute numbers from south to north similar to that found by Brichau et al. (2006). Sedimentary sequences. The evolution of hanging wall basins yields information regarding the timing of extension. Gautier et al. (1993) noted the relationship between the extensional ductile shearing and the opening of the Mio-Pliocene basins. The sedimentary sequences represent the infilling of half grabens opened between major normal faults during extension. The sediments within the basins represent an archive of the eroded source rocks and thus single crystal ages on the detrital grains may yield information, either about the source rocks, or about the amount (if any) of heating undergone after deposition. Thus, ages on sediments were also completed in order to measure the exhumation rate of the granodiorite and metamorphic rocks at times prior to those recorded in the present surface samples as well as to estimate possible burial temperatures associated with the units after deposition. (1) The Pesulia Formation (23.5–15 Ma, van Bo¨ger 1983). Fission-track ages of detrital zircons

Table 1. Zircon fission-track ages from Naxos Sample Granodiorite H 1* H7 H 10*

Schists H4 H6 H 11 H 19 HPH2* Na 05-41** Na 05-49** NA 05-50** Migmatite H 17 H 21 NAX2* NAX6*

368590 06.900 0258230 06.900 378050 27.300 0258250 11.500 378050 02.000 0258200 26.000 378070 14.100 0258240 36.300

Alt (m)

Irradiation

No. grains counted

rd  105 cm22 (counted)

rs  105 cm22 (counted)

ri  105 cm22 (counted)

U (ppm)

P(x 2) Var %

Central age + 2s (Ma)

1

eth-251-3

20

4.07 (2219)

57.2 (1454)

117 (2967)

1118

0 (17)

12.2 + 1.4

80

eth-252-3

4

7.59 (3965)

53.75 (1036)

174 (807)

938

25 (4)

13.7 + 2.2

20

eth-252-5

20

7.33 (1681)

34.3 (1036)

1213 (3664)

645

0 (25)

12.4 + 1.8

70

eth-250-2

20

4.72 (2515)

52.8 (1627)

118 (3633)

976

56 (,1)

12.6 + 1.0

378010 26.600 0258290 03.6 368560 07.100 0258280 30.600 378040 53.100 0258340 51.200 378110 50.800 0258310 58.900 368570 22.800 0258310 49.800 378110 18.500 0258310 25.900 378100 52.100 0258300 39.800 378100 52.100 0258300 39.800

410

eth-253-2

20

4.98 (2628)

23.8 (1498)

42.8 (2697)

352

,1 (15)

16.2 + 1.8

5

eth-253-5

17

4.60 (2628)

25.4 (866)

33.1 (1131)

298

32 (9)

20.5 + 2.4

50

eth-252-7

13

6.54 (3965)

31.98 (740)

115 (2665)

774

40 (0)

10.6 + 1.2

180

eth-250-10

14

3.75 (2515)

16.9 (372)

32.1 (705)

382

91 (0)

11.9 + 2.4

0

eth-271-23

9

3.23 (2133)

56.5 (661)

43.3 (507)

536

85 (45)

25.2 + 3.8

378060 40.200 0258300 39.100 378050 26.200 0258260 52.200 378090 26.100 0258310 08.000 378050 07.500 0258280 29.800

260

YU-Z0 (U38Y)

10

2.735 (1707)

7.528 (70)

36.03 (335)

145

98 (0)

10.6 + 3.0

209

YU-Z0 (U38Y)

20

2.744 (1712)

28.19 (322)

148.6 (1697)

596

97 (0)

9.7 + 1.4

209

YU-Z0 (U38Y)

20

2.753 (1718)

17.20 (365)

83.46 (1771)

334

78 (1)

10.5 + 1.6

400

Eth-252-11

9

5.83 (3965)

26.6 (691)

88.2 (2293)

589

0 (25)

10.6 + 2.0

200

eth-252-14

20

5.30 (3965)

25.9 (843)

82.3 (2678)

640

92 (0)

9.7 + 1.0

723

eth-339A-2

9

48.1 (2161)

323 (449)

1152 (1602)

935

0 (40)

9.3 + 2.8

436

eth-339-4

20

47.1 (2161)

17.8 (979)

484 (2667)

401

0 (23)

10.5 + 1.4 185

(Continued)

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

H 24*

Coordinates

186

Table 1. Continued Sample

Alt (m)

Irradiation

No. grains counted

rd  105 cm22 (counted)

rs  105 cm22 (counted)

ri  105 cm22 (counted)

U (ppm)

P(x 2) Var %

Central age + 2s (Ma)

378050 3200 0258200 5100

2

eth-253-9

20

4.09 (2628)

57.3 (2081)

105 (3811)

1034

5 (9)

12.0 + 1.2

378080 36.200 0258270 13.300

1

eth-250-7

10

4.15 (2515)

71.6 (1464)

123 (2508)

1262

0 (21)

12.5 + 2.4

160

eth-250-4

19

4.45 (2515)

196 (3674)

41.4 (777)

371

0 (35)

123 + 23

378060 1800 0258250 3100 378070 27.100 0258240 10.400

1

eth-250-12

10

3.58 (2515)

97.6 (1306)

33.9 (454)

388

0 (34)

63.3 + 15.8

378060 31.000 0258220 26.600

2

eth-250-9

15

3.92 (2515)

59.0 (2003)

96.7 (3282)

961

0 (15)

12.3 + 1.6

rs and ri represent sample spontaneous and induced track densities; P(x 2) is the probability of x 2 for v degrees of freedom where v ¼ number of crystals 2 1. ld ¼ 1.55125  10210. Zircon grains for AZ were irradiated at the ANSTO facility, Australia; those for DS (*) and ST (**) at the Oregon State University, Oregon. Zircons of AZ and DS were counted at 1600 (oil); for ST at 1250 (dry). All other counting (apatite and mica) was at 1250 (dry). Ages were calculated using the zeta approach (Hurford & Green 1983). zCN1/zircon ¼ 117 + 3 (AZ); *120 + 5 (DS); ** zCN5/zircon ¼ 371.5 + 14.0 (ST). All ages are reported as central ages (Galbraith & Laslett 1993) with a 2s error.

D. SEWARD ET AL.

Sediment H9 Eremonisia (sandstone) H 20 Eremonisia (Granodiorite boulder) H 23 Pesulia (sandstone) H 25 Pesulia (sandstone) H 26 Eremonisia (Granodiorite boulder)

Coordinates

Table 2. Apatite fission-track ages from Naxos Sample number Granodiorite H1 H7 H 10

Schist H6 H 19 HPH2* NA 05-41** NA 05-43** NA 05-48** NA 05-49** NA 05-50** NAX12* NAX13* Migmatite H 18 NAX1*

368590 06.900 0258230 06.900 378050 27.300 0258250 11.500 378050 0200 0258200 2600 378070 14.100 0258240 36.300 368560 07.100 0258280 30.600 378110 50.800 0258310 58.900 368570 22.800 0258310 49.800 378110 18.500 0258310 25.900 378090 28.000 0258320 48.900 378100 54.800 0258300 24.500 378100 52.100 0258300 39.800 378100 52.100 0258300 39.800 378050 57.600 0258260 24.300 378050 57.600 0258260 24.300 378090 25.900 0258300 27.100 378090 21.500 0258310 30.800

No. grains counted

rd  105 cm22 (counted)

rs  105 cm22 (counted)

eth-245-3

20

14.38 (5127)

1.168 (79)

21.47 (1452)

18

80

eth-245-12/13

17

9.47 (5127)

1.713 (65)

22.80 (865)

34

100 (0)

11.5 + 3.6

20

eth-246-2

17

11.56 (4527)

1.161 (17)

20.54 (71)

23

62 (5)

10.6 + 3.4

70

eth-248-3

20

11.33 (4500)

24.3 (2354)

26

36 (15)

9.0 + 2.6

5

eth-245-9

20

11.28 (5127)

1.49 (213)

3

94 (0)

12.0 + 7.0

eth-247-8/9

20

9.31 (4527)

1.43 (174)

19.8 (2417)

25

eth-260-2

20

9.3 (2424)

0.95 (75)

19.2 (1518)

26

67 (1)

8.1 + 2.0

260

YU-A0 (U34B)

20

12.17 (3797)

0.592 (25)

9.87 (417)

9

99 (0)

13.0 + 5.4

169

YU-A0 (U34B)

20

12.06 (3764)

0.078 (4)

2.53 (130)

2

49 (74)

6.5 + 7.0

190

YU-A0 (U34B)

20

11.96 (3731)

0.4078 (29)

7.57 (539)

7

99 (0)

11.5 + 4.4

209

YU-A0 (U34B)

20

11.85 (3698)

0.151 (11)

2.51 (183)

2

99 (0)

12.7 + 8.0

209

YU-A0 (U34B)

20

11.74 (3665)

0.344 (18)

7.33 (384)

7

76 (1)

9.8 + 4.8

20

eth-363-13

20

12.06 (2994)

0.056 (6)

12.3 (132)

1

100 (0)

9.7 + 8.2

20

eth-363-14

40

11.93 (2994)

0.339 (83)

5.60 (1370)

6

99 (0)

12.8 + 3.0

350

eth-247-5/6

20

9.99 (4237)

0.99 (109)

24.8 (2743)

27

26 (0)

6.4 + 1.8

337

eth-313-7

10

10.47 (2597)

3.42 (191)

64.2 (3481)

77

Alt. (m)

1

180 1

Irradiation

1.17 (113)

0.098 (14)

ri  105 cm22 U (ppm) P(x 2) Var % (counted)

Central age + 2s (Ma)

10 (33) 12.9 + 4.4

5 (24) 11.6 + 3.4

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

H 24

Coordinates

5 (20) 10.1 + 2.0 187

(Continued)

188

Table 2. Continued No. grains counted

rd  105 cm22 (counted)

eth-313-8

20

10.34 (2597)

1.37 (162)

25.9 (3076)

31

194

eth-297-13

20

11.1 (2579)

1.12 (123)

23.1 (2543)

26

99 (0)

9.5 + 1.8

243

eth-313-9

20

10.2 (2579)

0.58 (73)

13.61 (1700)

17

96 (0)

7.8 + 1.8

256

eth-339B-32

20

11.51 (3509)

2.88 (265)

62.9 (5790)

219

36 (10)

9.2 + 1.2

eth-246-8/9/10

20

9.87 (4527)

0.21 (18)

1.96 (160)

4

65 (5)

18.0 + 9.6

0

eth-246-11/12/13

20

9.15 (4527)

1.72 (208)

12.9 (1557)

20

75 (0)

19.8 + 5.0

1

eth-247-10/11

20

8.86 (4237)

0.89 (108)

15.74 (1918)

24

160

eth-248-2

16

11.6 (4500)

1.45 (63)

20.2 (878)

22

71 (9)

13.5 + 4.6

378070 27.100 0258240 10.400

1

eth-248-5

20

10.9 (4500)

1.35 (72)

16.0 (849)

19

58 (4)

15.0 + 4.8

378060 31.000 0258220 26.600

2

eth-248-6/7

17

10.6 (4500)

1.51 (106)

30.7 (2151)

35

eth-363 11

40

12.33 (2994)

0.07 (18)

1.26 (323)

Coordinates

NAX2*

378090 26.100 0258310 08.000 378090 50.900 0258290 51.600 378080 34.000 0258290 38.200 378030 34.100 0258270 28.800

723

378040 46.100 0258350 40.700

20

378050 09.900 0258350 38.000 378080 36.200 0258270 13.300

NAX3* NAX5* NAX7* Sediment H14 Eremonisia Sandstone H15 Eremonisia Sandstone H 20 Eremonisia (Granodiorite boulder) H 23 Pesulia (sandstone) H 25 Pesulia (sandstone) H 26 Eremonisia (Granodiorite boulder) NAX9* Weathered granite within landslide

378060 1800 0258250 3100

378050 57.600 258260 24.300

Alt. (m)

140

Irradiation

rs  105 cm22 (counted)

ri  105 cm22 U (ppm) P(x 2) (counted) Var %

1.3

Central age + 2s (Ma)

0 (40) 10.4 + 2.6

6 (29)

1 (70)

100 (0)

8.5 + 2.6

9.0 + 3.0

12.2 + 6.0

rs and ri represent sample spontaneous and induced track densities; P(x 2) is the probability of x 2 for v degrees of freedom where v ¼ number of crystals – 1. ld ¼ 1.55125  10210. Apatite grains for PH were irradiated at the ANSTO facility, Australia; those for DS (*) and ST (**) at the Oregon State University, Oregon. Apatite and mica of PH, DS, and ST were counted at 1250. Ages calculated using the zeta approach (Hurford & Green 1983). zCN5 ¼ 324 + 32 (PH); *zCN5/apatite 355 + 5 (DS); **zCN5/apatite ¼ 342.5 + 3.8 (ST). All ages are reported as central ages (Galbraith & Laslett 1993) with a 2s error.

D. SEWARD ET AL.

Sample number

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

189

Fig. 3. Distribution of zircon and apatite fission-track ages in the metamorphic core projected onto a line 0258 from south to north, A–A0 (Fig. 2).

were obtained from two sandstone samples in the west, samples H23 and H25 (Table 1 and Fig. 2a). The central ages of these are 123 +23 and 63.3 + 15.8 Ma respectively. The error margins are large as the central ages are from detrital grains with a large age spread. The single-grain ages range from 200 to 39 Ma (Fig. 6a) with an isolated 501 Ma age. There is no single clear mode and no ages are younger than the stratigraphic age. Apatite fissiontrack ages (Table 2) from this formation have central ages of 13.5 + 4.6 and 15.0 + 4.8 Ma with a mode of 15– 20 Ma (Fig. 6b) with single grain ages ranging from 35 to less than 5 Ma for the grouped data. (2) The Eromonisia Formation (11 –7.0 Ma, van Bo¨ger 1983). Samples H20 and H26 (Fig. 2a) were each from a boulder of granodiorite. Central zircon fission track ages are 12.5 + 2.4 and 12.3 + 1.6 Ma (Table 1) which is slightly older than the apatite fission-track ages of 8.5 + 5.2 and 9.0 + 3.0 Ma respectively (Table 2). Sample H9 is from a sandstone and has a zircon fission track age of 12.0 + 1.2 Ma in agreement with the granodiorite boulder ages. This sample has a chi-square value that predicts that all grains most likely belong to one population and hence were derived from a single source. On the eastern margin of the island, apatite fission-track ages are 18.0 + 9.6 and 19.8 + 5.0 Ma (H14, H15, Table 2 and Fig. 2) with a very large spread from 80 –75 Ma down to less than 5 Ma. Any interpretation from such data sets must be done with caution remembering also that the number of grains suitable for counting was probably statistically insufficient (Vermeesch 2005).

A single sample of a weathered granite from a landslide (NAX9) also yielded a very imprecise apatite fission-track age of 12.2 + 6.0 Ma, (Fig. 2 and Table 1). This age is compatible with the evolution of the metamorphic sequences in the proximity from which the sample was most likely sourced.

Interpretation and discussion The new fission-track ages presented above were obtained from samples of all lithotectonic units exposed in Naxos and thus provide an extensive data base to constrain the thermal history of these rocks in the 300–60 8C temperature range. These data complement results from previous thermochronological studies and are discussed taking into account the tectonic history of Naxos Island based on published structural and metamorphic data. This discussion will also reassess previously published thermal and tectonic models accounting for thermochronometric data obtained previously.

Metamorphic sequences The 25.2–6.4 Ma age range obtained from Naxos metamorphic rocks from fission-track analysis partly covers the age ranges obtained by previous studies based on Rb/Sr, Ar/Ar, fission-track and U–Th/He thermochronology. No simple, single vertical or near-vertical sections were sampled and hence the relationship of age with altitude could not be used to determine confidently rates of exhumation.

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Fig. 5. Distribution of zircon and apatite fission-track ages from south to north across the granodiorite.

Fig. 4. Zircon length distributions for three samples from the southern region to the migmatites.

The oldest zircon fission-track ages within the metamorphic rocks are recorded to the south of the island. The zircon length distributions (Fig. 4) indicate that the ages are mixed whereby the southernmost sample (H6, Fig. 4) has a slightly bimodal distribution which is also slightly expressed in sample H4, to the north of H6. This indicates that these samples may have either remained in the zircon fission-track partial annealing zone in the early stages of cooling for some time and then cooled more rapidly or they may represent a record of a more complex thermal history with a first cooling event followed by partial resetting of the fission-track thermochronometer during a subsequent heating event and then final cooling. Argon thermochronology also indicates a northward younging of ages (Andriessen et al. 1979, 1987; Wijbrans & McDougall 1986, 1988) but does not fully discriminate between these two models. In

Fig. 6. Apatite and zircon single grain age distributions for the samples from the Pesulia Formation.

contrast, the analysis of metamorphic assemblages, the difference in composition of white micas from north to south suggests partial resetting of a HP –LT (M1) metamorphism during a subsequent MT metamorphism (M2) (Wijbrans & McDougall 1986; Avigad 1998; Ducheˆne et al. 2006). Partial resetting is furthermore consistent with the various Rb/Sr isochrons obtained on metamorphic rocks and minerals therein such as biotite, muscovite and garnet indicating the impact of metamorphic reactions on radiochronometric systems (Ducheˆne et al. 2006). The tight clustering of the track-length distribution for sample H17, located within the

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

migmatite with a mean length of 10.67 + 0.13 mm implies rapid cooling through the partial annealing zone of the zircons from approximately 330– 170 8C, (Yamada et al. 1995). Accordingly, this pattern of track-length distribution characterized by a slight bimodal distribution in the south and a tight clustering in the north, associated with a northward increase in the mean track length and decrease in the standard deviation from sample H6 to H17 is best interpreted in terms of partial resetting of older cooling ages during subsequent reheating followed by a renewed phase of cooling. The new zircon fission-track ages presented in this paper thus suggest that following HP –LT metamorphism, the schists exposed in the south of Naxos were exhumed and cooled below c. 300 8C. These rocks were then reheated to temperatures corresponding to the partial annealing zone for zircon fission tracks (330 –170 8C) during the subsequent M2 metamorphism. The efficiency of this reheating is a function of the structural position as samples in the south of the island best preserve the relics of the M1 HP–LT metamorphism. Because closure temperatures are related to cooling rates, (Dodson 1973; Reiners & Brandon 2006) the zircon fission-track thermochronometer probably had a closure temperature closer to 300 8C (Yamada et al. 1995) a little less than the one for argon thermochronology on muscovite and hornblende. The spatial distribution displayed by ages obtained by argon thermochronology on hornblende and mica (closure temperature ranging from c. 525 to 300 8C, Reiners & Brandon 2006) reveal a linear younging trend from the south to the north

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towards the detachment (Fig. 7). This suggests that these metamorphic rocks cooled through the 525 to c. 300 8C isotherms as they were progressively exhumed. Ages obtained with thermochronometers with different closure temperatures display similar slopes in the graph representing ages as a function of distance of the sample to the detachment suggesting that the cooling rate was constant for all samples irrespective of their position. The northward younging indicates that rocks in the south were exhumed through the different closure isotherms before those in the north and both the linear younging trend and the similar slopes for distinct thermochronometers suggest a steady-state cooling that is best interpreted in terms of steady-state tectonic exhumation according to a top-to-the-north detachment as already proposed (John & Howard 1995). In contrast to the higher-temperature thermochronometers, the zircon and apatite fission-track ages do not show this simple spatial distribution. In the southern part of the island, we proposed above that the older zircon fission-track ages are related to partial resetting of the fission-track thermochronometer. Insufficient track length data precludes such an analysis for the apatite fission track ages but all ages are 13 Ma and no remnants of the M1 Eocene metamorphic event are identified. Making abstraction to these older ages, zircon and apatite fission-track ages display a significant northward younging trend (Fig. 3) over the median core region but then older towards the detachment. This is at odds with the simple linear northward younging trend proposed on the basis of a less extensive database by Brichau et al. (2006).

Fig. 7. K/Ar, Rb/Sr and 40Ar/39Ar ages on micas and hornblende from Wijbrans & McDougall (1986, 1988) and Andriessen et al. (1979, 1987), plotted onto line A– A0 .

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A plausible explanation for the slight increase in ages towards the detachment is the more efficient cooling at the vicinity of the hanging wall as demonstrated by thermal modelling of the footwall of normal faults (e.g. Grasemann & Mancktelow 1993; Ehlers et al. 2001). In these models, the warmer footwall is exhumed and juxtaposed against the colder hanging wall, which results in downward deflection of the isotherms towards the hanging wall. This effect is stronger for lowtemperature isotherms and will thus be best recorded by thermochronometers with lower closure temperatures. The rocks proximal to the fault are exhumed through the isotherms corresponding to their closure temperatures earlier than those more distal. They may also be following a non-vertical path of cooling especially those being exhumed along a low angle detachment. With refined sampling, the new fission-track data attest to the downwarping of the isotherms of the metamorphic rocks close to cold hanging-wall sediments whereby rocks closest to the detachment reach the closure temperatures prior to those more distal. As noted above, the higher temperature chronometric data of Andriessen et al. (1979, 1987), Wijbrans & McDougall (1986, 1988) and John & Howard (1995) (Fig. 7) do not display this effect. At the temperature and associated depth regime of 525 – 300 8C, the impact of the fault and associated hanging wall should be minimal and the isotherms not be bent down sufficiently in the footwall to yield a difference in ages. Within the migmatites there is a predominance of both apatite and zircon fission-track ages at 10–9 Ma attesting to a period of rapid cooling from c. 300 to 60 8C at this time. Additionally NAX3 has an apatite fission-track age of 9.5 + 1.8 Ma and an apatite (U –Th)/He age of 10.9 Ma (Vermeesch et al. 2007) extending this rapid cooling rate to within a temperature range of about 40 8C at this time. Sample NAX13 (Fig. 2) has a slightly older apatite fission-track age (12.8 + 1.8 Ma) than those along the same ‘structural’ latitude: it falls away from the curve. However, this sample lies in the hanging wall of a north –south normal fault and its age is compatible with the age it would have had prior to faulting, i.e. it is older than the ages in its associated footwall. If this is correct, then the faulting is younger than 12.8 Ma. Recent mapping (Siebenaller 2008) has revealed a series of east –west oblique strike-slip to normal faults that cross the northern region in particular. Inspection of the fission-track ages, however, reveals a certain randomness to the relationship of these ages to the faults, i.e. there is no consistent pattern with older ages for example in the hanging wall of normal faults as would be expected if the

ages were related at all to these younger structures. One concludes that vertical displacement has been insufficient to affect the ages on either side of these younger cross-cutting faults. Neither is any significant vertical displacement obvious from the geological map.

Granodiorite The zircon fission-track ages are identical for the granodiorite, at a one sigma level, to the U –Pb cooling ages previously determined by Keay et al. (2001). Shallow intrusion is supported by the emplacement pressure of 1.5–2 kbar using data from Holdaway (1971) for the aluminium silicate triple point (Gautier et al. 1993), as well as the fact that the granodiorite has an intrusive contact with the marble and schists in the south where the metamorphic rocks had already cooled to temperatures of less than 350 8C. Such results suggest very rapid cooling to at least 300 8C. In the south this rapid cooling took place to within the temperature regime of at least 110 8C as evidenced by the coincidence of the apatite and zircon fission-track ages. Northwards, however, zircon and apatite fission track ages are different (Fig. 5). Near the current contact with the detachment in the north, the zircon fission-track age is 12 Ma while the apatite fission-track age is as low as 9 Ma. The range in ages yields a variation in cooling rates from 100 8C Ma21 to c. 50 8C Ma21. This discrepancy suggests that the host rock of the granodiorite may have been cooler in the south compared to the north at the time of emplacement. Considering a pre-intrusion homogeneous distribution of isotherms with depth, this implies that the granodiorite exposed in the south was emplaced at a shallower depth than the one exposed in the north. In contrast, if isotherms in the host rock were not horizontal at the time of granodiorite emplacement then this fission-track age distribution reflects granodiorite emplacement across isotherms. Metamorphic isograds are currently dipping towards the south which might be a record of dipping isotherms at the time of granodiorite emplacement. Such a dip is also consistent with tilting of the footwall of the northward dipping detachment owing to isostatic rebound. These models are not mutually exclusive and we propose that one solution for the fissiontrack age pattern of the granodiorite is the reflection of the combined effects of both a different emplacement depth and granodiorite emplacement across isotherms. The greater exhumation of the northern tip of the granodiorite is potentially related to a switch from ductile shearing along shallow-dipping shear planes expressed by C–S fabrics to cataclastic faulting dissecting the northern part of the pluton. This switch in deformation style corresponds to

NAXOS TECTONICS: FISSION-TRACK ANALYSIS

the ductile –brittle transition reached in the temperature range 300 –200 8C that was reached soon after emplacement of the granodioritic magma in a host rock cooler than c. 330 8C. Another possibility for the northward younging trend of the apatite fission-track ages (from 12 to 9 Ma) is that they are purely due to tectonic denudation. If exhumation is solely due to progressive cooling during exhumation along the northern detachment, a slip rate of 4 mm a21 is obtained (Fig. 5). This is slightly lower than that calculated by John & Howard (1995) who estimated 5–8 mm a21 for the metamorphic rocks when they were at temperature greater than c. 330 8C. Brichau et al. (2006) made estimates of slip rates, using four thermochronometers, ranging from 13.2 to 9.1 mm a21 over the time period 12–9 Ma. The rate is more than twice that observed in this study and the discrepancy most likely hinges on the fact that the new fission-track ages on zircon are a little older, though within error limits, in the southern region. The slip rate of Brichau et al. (2006) may have been inappropriately calculated by combining ages from the metamorphic rocks and the granodiorite. As observed above, these two lithological units recorded different cooling histories.

Sedimentary sequences In the Pesulia Formation (23.5– 15 Ma, van Bo¨ger 1983), the lack of zircon fission-track ages less than the stratigraphic age supports the notion that this formation has not been exposed to temperatures greater than 170 –200 8C (the lower bound of the zircon fission-track partial annealing zone) and certainly not greater than 300 8C since deposition. The source of these sediments is unclear from this data set, but, they may have travelled long distances. The age spectrum is in accordance with that determined by Kuhlemann et al. (2004) who suggested a source for equivalent samples to the south. The apatite fission-track ages of samples, H23 and H25 from the Pesulia Formation are 13.5 + 4.6 and 15.0 + 4.8 Ma respectively (Table 1 and Fig. 2). These ages are much younger than the associated zircon fission-track ages and are within the age span of sedimentation. The grouped data sets reveal (n ¼ 52) a mode between 15 and 20 Ma with 25% of the ages less than this (Fig. 6). On the supposition that the apatites are probably from the same source as the zircons, implications are that the sediments have been heated into the apatite partial annealing zone since deposition (60–110 8C), a conclusion strongly supported by the fact that some ages are younger than the supposed age of sedimentation. Unfortunately track lengths that might have allowed some thermal modelling are very rare, due to a very low spontaneous

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track density as a response to the low uranium concentration. Both of these samples were deposited prior to the intrusion of the granodiorite and now are horizontal distances of less than 1.2 km away from the exposed contacts, i.e. they were potentially heated by the granodiorite at the time of its emplacement. The apparent ages are slightly older than the granodiorite intrusion time. Thus the source of heating could be the proximity of the granodiorite. Alternatively, heating of the sediments could reflect burial followed by exhumation together with the granodiorite and the metamorphics at 12 –9 Ma. The amount of burial would have had to be of the order of ,6 km but .3 km. The current thickness of the sedimentary sequence on Naxos and neighbouring islands is of the order of less than a kilometre. If a sedimentary sequence of 3 –6 km had been deposited, it would require erosion of more than 2 km in 2 –3 Ma. Ages of granodiorite boulders in the Eremonisia Formation on the northern hanging wall to the granodiorite, are similar to those presently cropping out in the same region. Furthermore, zircon and apatite fission-track ages obtained on the granodiorite pluton and on the boulders are similar within error range. These data suggest that the boulders are coming from the local erosion of the pluton. The youngest apatite fission track ages of 8.5 + 2.6 and 9.0 + 3.0 Ma indicate that they had crossed the c. 110 8C isotherm at that time. The Eremonisia Formation is estimated to be no younger than 7.0 Ma: the time of sedimentation (van Bo¨ger 1983). This yields a lag time of about 1.5– 2 Ma implying rapid exhumation of the source rocks. The source rocks of this age are still currently cropping out, which implies that there has been little erosion of the granodiorite since its deposition into the Eremonisia Formation. This scenario is further supported by the preservation of paleo-erosional surfaces in the centre of Naxos. Riedl (1984a, b) proposed an Upper Miocene to Lower Pliocene age for this paleo-erosional surfaces based on climatic and geomorphological reasoning that suggests that there has been limited erosion since that period.

Conclusion This study has shown the value a multi-facetted fission-track approach on a complex geological target. In particular, this builds not only upon existing geological data, but draws its conclusions from a full range of new fission-track analyses on all exposed geological units of the island of Naxos. The study has also emphasized the need for tighter sampling in such complex tectonic settings where rocks might have recorded very distinct thermal histories.

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On this basis the new ages on different rock types has allowed the determination of the following. (1) Cooling–reheating of metamorphic rocks during the metamorphic history (HP –LT followed by MT) followed by cooling due to exhumation during extension along low-angle shear zones. The impact of the upwelling footwall against a colder hanging wall is emphasized in the cooling ages whereby the youngest ages are not at the detachment fault but some distance away that is interpreted to be due to the effects of downwarping of the isotherms in the proximity of the fault. The closeness of sampling in this study has revealed the slight variations of cooling that was not recognized in previous studies. The spacing confirms that slip rates on detachment faults determined in other studies may be inaccurate. Further, slip rates determined by Brichau et al. (2006) were made across the detachment from the granodiorite to the metamorphic sequences. This may not be appropriate as our data shows that the unroofing may not have been as simple as one detachment controlling the exhumation. (2) Rapid cooling of a magma emplaced into a colder host rock thus constraining the temperature both of the host rock as well as the emplacement of the magma to a temperature of c. 300 8C. Differential cooling of the pluton is related to either to variable depth of emplacement and/or to progressive tectonic exhumation accommodated first by ductile shallow-shear zone and then by high-angle cataclastic faults. (3) Cooling ages of the detrital grains in order to trace their source, as well as the thermal history of the sediments after deposition, which has added information about the rates of erosion and associated cooling in the source region. We thank J.-P. Burg and B. den Brock for supervision with the field collections of A. Zingg and P. Holzner as well as collecting later samples during student field excursions as the project progressed. Brian Wernicke and Steven Kidder are thanked for their helpful reviews.

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B RANDON , M. T., L ISTER , G. S. & W ILLETT , S. (eds) Exhumation Processes. Geological Society, London, Special Publications, 154, 509–522. V EERMESCH , P. 2005. How many grains are needed for a provenance study? Earth and Planetary Science Letters, 224, 441–451. V ERMEESCH , P., S EWARD , D., L ATKOCZY , C., W IPF , M., G U¨ NTHER , D. & B AUR , H. 2007. a-Emitting mineral inclusions in apatite, their effect on (U–Th)/He ages, and how to reduce it. Geochimica et Cosmochimica Acta, 71, 1737– 1746. VAN B O¨ RGER , H. 1983. Stratigraphische und tektonische verknu¨pfungen kontinentaler Sedimente ¨ ga¨is-Raum. Geologische Rundschau, des Neogens im A 72, 771–814. V ANDERHAEGHE , O. 2004. Structural development of the Naxos migmatite dome. In: W HITNEY , D. L., T EYSSIER , C. & S IDDOWAY , C. S. (eds) Gneiss domes in orogeny. Geological Society of America Special Paper, 380, 211– 227.

V ANDERHAEGHE , O., H IBSCH , C. ET AL . 2007. Penrose Conference – extending a continent – Naxos Field Guide. In: L ISTER , G., F ORSTER , M. & R ING , U. (eds) Inside the Aegean metamorphic core complexes. Journal of Virtual Explorer, Electronic Edition, ISSN 1441-8142, 27, paper 4. W IJBRANS , J. R. & M C D OUGALL , I. 1986. 40Ar/39Ar dating of white micas from an Alpine high-pressure metamorphic belt on Naxos (Greece): the resetting of the argon isotopic system. Contributions to Mineralogy and Petrology, 93, 187– 194. W IJBRANS , J. R. & M C D OUGALL , I. 1988. Metamorphic evolution of the Attic Cycladic metamorphic Belt on Naxos (Cyclades, Greece) utilizing 40Ar/39Ar ages spectrum measurements. Journal of Metamorphic Petrology, 6, 571– 594. Y AMADA , R., T AGAMI , T., N ISHIMURA , S. & I TO , H. 1995. Annealing kinetics of fission tracks in zircon: an experimental study. Chemical Geology, 122, 249–258.

Syn-extensional granitoids in the Menderes core complex and the late Cenozoic extensional tectonics of the Aegean province ¨ NER1 YILDIRIM DILEK1*, S¸AFAK ALTUNKAYNAK2 & ZEYNEP O 1

Department of Geology, 116 Shideler Hall, Miami University, Oxford, OH 45056, USA 2

Department of Geology Engineering, Istanbul Technical University, Maslak 80626, Istanbul, Turkey *Corresponding author (e-mail: [email protected]) Abstract: The Miocene granitoid plutons exposed in the footwalls of major detachment faults in the Menderes core complex in western Anatolia represent syn-extensional intrusions, providing important geochronological and geochemical constraints on the nature of the late Cenozoic magmatism associated with crustal extension in the Aegean province. Ranging in composition from granite, granodiorite to monzonite, these plutons crosscut the extensional deformation fabrics in their metamorphic host rocks but are foliated, mylonitized and cataclastically deformed in shear zones along the detachment faults structurally upward near the surface. Crystallization and cooling ages of the granitoid rocks are nearly coeval with the documented ages of metamorphism and deformation dating back to the latest Oligocene– early Miocene that record tectonic extension and exhumation in the Menderes massif. The Menderes granitoids (MEG) are represented by mainly metaluminous-slightly peraluminous, high-K calc-alkaline and partly shoshonitic rocks with their silica contents ranging from 62.5 to 78.2 wt%. They display similar major and trace element characteristics and overlapping inter-element ratios (Zr/Nb, La/Nb, Rb/Nb, Ce/Y) suggesting common melt sources. Their enrichment in LILE, strong negative anomalies in Ba, Ta, Nb, Sr and Ti and high incompatible element abundances are consistent with derivation of their magmas from a subduction-metasomatized, heterogeneous sub-continental lithospheric mantle source. Fractional crystalization processes and lower to middle crustal contamination also affected the evolution of the MEG magmas. These geochemical characteristics of the MEG are similar to those of the granitoids in the Cyclades to the west and the Rhodope massif to the north. Partial melting of the subduction-metasomatized lithospheric mantle and the overlying lower-middle crust produced the MEG magmas starting in the late Oligocene–early Miocene. The heat and the basaltic material to induce this partial melting were provided by asthenospheric upwelling caused by lithospheric delamination. Rapid slab rollback of the post-Eocene Hellenic subduction zone may have peeled off the base of the subcontinental lithosphere, triggering the inferred lithospheric delamination. Both slab retreat-generated upper plate deformation and magmatically induced crustal weakening led to the onset of the Aegean extension, which has migrated southward through time.

The Aegean extensional province is a rapidly deforming and seismically active domain in the Africa–Eurasia convergent zone in the eastern Mediterranean region and is considered to have evolved as a backarc tectonic environment above the north-dipping Hellenic subduction zone (Fig. 1; Le Pichon & Angelier 1979; Jolivet 2001; Faccenna et al. 2003; van Hinsbergen et al. 2005; Jolivet & Brun 2008). Southward retreat of the subducting African lithosphere along the Hellenic trench and the faster SW motion of the southern part of the Anatolian block in the upper plate have resulted in approximately north–south extension in the Aegean region since the Oligo-Miocene (Jolivet et al. 1994; Jolivet & Faccenna 2000; Ring & Layer 2003). The thrust front associated with this subduction zone and its slab retreat has also migrated from the Hellenic trench to the

south of the Mediterranean Ridge since then (Le Pichon et al. 2003). These observations suggest that the driving forces for regional extensional tectonics in the broader Aegean region reside mainly within the retreating lithospheric slab. Subduction of the Tethyan mantle lithosphere northward beneath Eurasia was nearly continuous since the latest Cretaceous, only temporarily punctuated by the collisional accretion of several ribbon continents (i.e. Pelagonia, Sakarya, Anatolide –Tauride) to the southern margin of Eurasia and related slab breakoff events (Rosenbaum et al. 2002; van Hinsbergen et al. 2005; Dilek & Altunkaynak 2007). Exhumation of middle to lower crustal rocks and the formation of extensional metamorphic domes occurred in the backarc region of this progressively southward migrated trench and the Tethyan slab throughout the Cenozoic.

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 197–223. DOI: 10.1144/SP321.10 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Tectonic map of the Aegean and eastern Mediterranean region, showing the main plate boundaries, major suture zones and fault systems. Thick, white arrows depict the direction and magnitude (mm a21) of plate convergence; grey arrows mark the direction of extension (Miocene–Recent). Orange and purple colours delineate Eurasian and African plate affinities, respectively. Key to lettering: BF, Burdur fault; CACC, Central Anatolian crystalline complex; DKF, Datc¸a– Kale fault (part of the SW Anatolian Shear Zone); EAFZ, East Anatolian fault zone; EF, Ecemis fault; EKP, Erzurum–Kars Plateau; IASZ, Izmir–Ankara suture zone; IPS, Intra-Pontide suture zone; ITS, Inner-Tauride suture; KF, Kefalonia fault; KOTJ, Karliova triple junction; MM, Menderes massif; MS, Marmara Sea; MTR, Maras triple junction; NAFZ, North Anatolian fault zone; OF, Ovacik fault; PSF, Pampak– Sevan fault; TF, Tutak fault; TGF, Tuzgo¨lu¨ fault; TIP, Turkish-Iranian plateau (modified from Dilek 2006).

Recent geochronological data from several core complexes (i.e. Menderes, Kazdag) and the synextensional granitoid plutons in them (Fig. 2) indicate that both extension and attendant magmatism in the Aegean region date back to the at least latest Oligocene – early Miocene (Bozkurt & Satir 2000; Okay & Satir 2000; Ring et al. 2003; Is¸ik et al. 2004; Thomson & Ring 2006). These findings suggest that tectonic extension and magmatism were synchronous events starting around 25– 24 Ma. The Aegean extension thus appears to have started c. 25 Ma, long before the onset of the Arabia– Eurasia collision-driven SW escape of the Anatolian microplate in the late Miocene (Barka & Reilinger 1997; Jolivet & Faccenna 2000). Syn-extensional magmatism during the early and middle Miocene produced widespread volcanic rocks and plutons in the Cyclades and western Anatolia (Altherr et al. 1988; Altherr & Siebel 2002; Pe-Piper & Piper 2002, 2006; Pe-Piper et al. 2002; Akay & Erdog˘an 2004; Bozkurt 2004; Gessner et al. 2004; Is¸ik et al. 2004; Ko¨pru¨bas¸i & Aldanmaz 2004; Innocenti et al. 2005; Ring & Collins 2005; Agostini et al. 2007; Dilek & Altunkaynak 2007; Akay 2008). The Oligo-Miocene volcanic rocks

have medium- to high-K calcalkaline compositions (Pe-Piper & Piper 2006; Altunkaynak & Genc¸ 2008), and their trace-element and isotope geochemistry characteristics indicate that parental magmas were derived from partial melting of an enriched lithospheric mantle (Aldanmaz et al. 2006; Altunkaynak & Genc¸ 2008) and that they underwent decreasing subduction influence and increasing crustal contamination through time (Altunkaynak & Dilek 2006; Dilek & Altunkaynak 2007, and references therein). Mildly alkaline bimodal volcanic products that were erupted during the middle Miocene (16– 14 Ma), on the other hand, show a decreasing amount of crustal contamination and subduction influence through time (Altunkaynak & Dilek 2006; Dilek & Altunkaynak 2007). Melting of a subduction-modified continental lithospheric mantle and asthenospheric mantlederived melt contribution both appear to have played a major role in the generation of the magmas of these middle Miocene volcanic rocks. The Oligo-Miocene and middle Miocene plutons in the region are represented by mainly I-type granitoids, whose chemical compositions display significant differences throughout the region

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Fig. 2. Geological map of western Anatolia, showing the distribution of Menderes and Kazdag metamorphic massifs, Neotethyan ophiolites, Cenozoic sedimentary basins and igneous provinces, salient fault systems. Major tectonic blocks and suture zones are also depicted. AF, Acigo¨l fault; BFZ, Burdur fault zone; DF, Datc¸a fault; IASZ, Izmir–Ankara suture zone; IPSZ, Intra-Pontide suture zone; KDM, Kazdag massif; KF, Kale fault; NAFZ, North Anatolian fault zone; SWASZ, Southwest Anatolian shear zone. Key to lettering for the detachment faults: AD, Alasehir detachment; GD, Gu¨ney detachment; SD, Simav detachment. Key to lettering for the granitoid plutons: AG, Alasehir; BG, Baklan; CGD, C ¸ ataldag; EGP, Egrigo¨z; EP, Eybek; EVG, Evciler; GBG, Go¨ynu¨kbelen; GYG, Gu¨rgenyayla; IGD, Ilica; KBG, Karabiga; KG, Kozak; KCG, Kusc¸ayiri; KOP, Koyunoba; KSG, Kestanbol; KTG, Katrandag; OGD, Orhaneli; SG, Salihli; SVG, Sevketiye; TG,Turgutlu; TGD, Topuk; YCG, Yenice.

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(Altherr & Siebel 2002). However, similar to their volcanic counterparts (Altunkaynak & Genc¸ 2008), these granitoids do not exhibit geochemical signatures of an active subduction zone magmatism for the origin of their magmas. The melt sources, the magmatic evolution and the nature of the heat supply for the production of syn-extensional granitoids in the Aegean province remain, therefore, some of the most signifcant questions in the late Cenozoic geodynamic evolution of this region. Granitic plutons commonly exposed in the footwalls of major detachment faults of the metamorphic core complexes in the region provide critical information to address some of these questions. In this paper we present new geochemical and isotope data as well as evaluating the extant geochemical data from the Miocene granitoid plutons intruded into the Menderes metamorphic core complex in western Anatolia in order to constrain their magma sources and tectonomagmatic evolution. We also present new field observations from the Menderes granitoids to document the synextensional nature of these intrusions. We compare the Menderes granitoid data with the available geochemical data from the granitoids emplaced into the other core complexes in the Central Cyclades and the northern Aegean in order to understand better the melt evolution of syn-extensional granitoids in western Anatolia. We then discuss the role of the late Cenozoic magmatism in the onset and the spatial and temporal progression of orogen-wide extension in the Aegean province within the framework of a regional tectonic model.

Basement geology of Cenozoic granitoids in western Anatolia The Cenozoic granitoid plutons in western Anatolia are intrusive into the crystalline basement rocks of the Sakarya continent and the Anatolide –Tauride block. They also crosscut the Tethyan ophiolites (c. .92– 90 Ma) and blueschist assemblages occurring along the Izmir–Ankara suture zone (IASZ) between the Sakarya and Anatolide– Tauride conti¨ nen & Hall 1993; Okay et al. nental blocks (Fig. 2; O ¨ nen 2003). The 40Ar/39Ar cooling ages 1998; O (phengite crystallization during exhumation) of 79.7 + 1.6– 82.8 + 1.7 Ma (Sherlock et al. 1999) from the blueschist rocks along the IASZ indicate a latest Cretaceous timing of the HP–LT metamorphism in the region. This event was followed by the collision of the Sakarya and Anatolide – Tauride blocks in the late Palaeocene –early Eocene (Harris et al. 1994; Dilek & Altunkaynak 2007). The Lycian ophiolite nappes structurally overlying the platform carbonates of the Tauride block farther south (Fig. 2; Collins & Robertson 2003; Ring & Layer 2003) represent the tectonic

outliers of the Cretaceous oceanic crust derived from the IASZ. These ophiolitic thrust sheets are inferred to have once covered the Menderes metamorphic massif and then to have been removed as a result of the tectonic uplift and erosion associated with the exhumation of the Menderes core complex during the late Cenozoic (Dilek & Whitney 2000; Ring & Layer 2003; Thomson & Ring 2006). The Sakarya continent consists of a Palaeozoic crystalline basement with its Permo-Carboniferous sedimentary cover and Permo-Triassic rift or accretionary-type melange units (Karakaya complex) that collectively form a composite continental block (Tekeli 1981; Okay et al. 1996). The Carboniferous felsic gneisses, amphibolites, marbles and meta-ophiolitic units that are tectonically intercalated with sillimanite-bearing gneisses and migmatites in the western part of the Sakarya continent constitute the Kazdag metamorphic massif exposed in the Biga Peninsula (KDM in Fig. 2; Okay et al. 1991; Duru et al. 2004). The NE –SW-trending Kazdag massif forms a structural dome and represents a metamorphic core complex. The highgrade metamorphic basement rocks in the massif are separated by the overlying unmetamorphosed middle Cretaceous accretionary melange (C ¸ etmi melange) along a mylonitic shear zone (Alakec¸i shear zone) that defines a detachment surface (Bonev & Beccaletto 2007). The metamorphic assemblages in the footwall of this shear zone record amphibolite-facies metamorphic conditions at 5 kbar and 640 8C that were reached around 24 Ma as constrained by Rb– Sr mica ages (Okay & Satir 2000). Kinematic indicators and deformation patterns in the mylonitic rocks suggest top-tothe-NNW (in general) normal shear sense and deformation mechanisms grading from ductile shear flow to brittle fracturing and cataclasis toward the top of the detachment zone (Bonev & Beccaletto 2007). The Kazdag core complex in the footwall of this detachment zone is inferred to have been exhumed starting at c. 24 Ma from a depth of c. 14 km along the north-dipping Alakec¸i mylonitic shear zone (Okay & Satir 2000). The Sakarya continental rocks and the ophiolitic units of the IASZ are intruded by a series of east – west trending Eocene and Oligo-Miocene plutons that are represented by I-type calc-alkaline granitoids (Fig. 2; Harris et al. 1994; Genc¸ 1998; Altunkaynak & Yılmaz 1998; Ko¨pru¨bas¸i & Aldanmaz 2004; Altunkaynak 2007). The plutons straddling the IASZ (the suture zone granitoids, SZG, of Altunkaynak 2007) range in composition from diorite, quartz diorite and granodiorite to syenite (Orhaneli, Topuk, Gu¨rgenyayla and Go¨ynu¨kbelen plutons) and have ages around 54 –48 Ma (Ataman 1972; Bingo¨l et al. 1982, 1994; Harris et al. 1994; Delaloye & Bingo¨l 2000; Yılmaz et al. 2001). The plutons farther north along the Marmara Sea (Marmara

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granitoids, MG, of Altunkaynak 2007) are composed of monzogranite, granodiorite and granite (Armutlu, Lapseki and Kapidag plutons) and have ages between 48 and 34 Ma (Ercan et al. 1985; Bingo¨l et al. 1994; Harris et al. 1994; Genc¸ & Yılmaz 1997; Delaloye & Bingo¨l 2000; Ko¨pru¨bas¸i et al. 2000; Ko¨pru¨bas¸i & Aldanmaz 2004). The granitoid plutons intruded into the Kazdag core complex farther west are Oligo-Miocene in age and crosscut the extensional foliation and lineation in the Kazdag massif, providing an upper age constraint for the timing of the extensional deformation in NW Anatolia. These plutons are composed of granite, granodiorite, quartz diorite and monzonite. They are metaluminous and subalkaline in nature and have medium- to high-K calc-alkaline compositions (Okay & Satir 2000; Genc¸ & Altunkaynak 2007). Rb –Sr biotite dating of the Evciler pluton (Fig. 2) revealed cooling ages of 20.7 + 0.2 Ma– 20.5 + 0.2 Ma, nearly identical to the Rb –Sr biotite ages of 18 –20 Ma from the gneissic rocks of the host Kazdag massif (Okay & Satir 2000). Birkle & Satir (1995) have also reported a Rb –Sr biotite age of 25 + 0.3 Ma from the northeastern part of the Evciler pluton. These cooling ages of the Evciler pluton are nearly coeval with a single-grain zircon SHRIMP age of 23.94 + 0.31 Ma from the Eybek pluton (Fig. 2; Altunkaynak and Dilek, unpublished data). These structural and temporal relations indicate that the syn-extensional granitoids in the Kazdag core complex were emplaced during the latest Oligocene – early Miocene, shortly after the peak deformation and metamorphism. The Menderes metamorphic massif farther south in the Anatolide block is a NE– SW-oriented, subelliptical dome divided into northern, central and southern sections that are separated by nearly east –west-trending structural grabens (Fig. 2). It consists of a Precambrian ‘core’ and Palaeozoic– Cenozoic ‘cover’ (Satir & Friedrichsen 1986; Bozkurt & Park 1994; Bozkurt & Oberha¨nsli 2001 and references therein). The core sequence includes augen gneisses, metagranites, high-grade schists and eclogitic metagabbros with metamorphic ages older than 50 Ma (Candan et al. 2000; Bozkurt & Oberha¨nsli 2001). The cover sequence consists of various schist types and metamorphosed carbonates and the protoliths of the cover sequences range in age from Palaeozoic to the early Eocene (Loos & Rischmann 1999; Bozkurt & Oberha¨nsli 2001; Rimmele´ et al. 2003). The core and cover sequences of the Menderes massif collectively comprise several nappe systems that were assembled mainly during the late Mesozoic –early Cenozoic collisional events in the region (Gessner et al. 2001; Ring et al. 2001; Reignier et al. 2007). The main episode of metamorphism in the Menderes massif is inferred to have resulted from the burial regime associated with the emplacement of

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the Lycian nappes and ophiolitic thrust sheets (Dilek & Whitney 2000; Yılmaz 2002). Imbricate stacking of the Menderes nappes beneath the Lycian nappes and ophiolitic thrust sheets appears to have migrated southwards throughout the ¨ zer et al. 2001; Palaeocene – middle Eocene (O Candan et al. 2005). The unroofing and exhumation of the Menderes massif may have started as early as the latest Oligocene–early Miocene (25–21 Ma) as constrained by the cooling ages of the syn-extensional granitoid intrusions crosscutting the metamorphic rocks (Bozkurt & Satir 2000; Catlos et al. 2002; Is¸ik et al. 2003; Ring & Collins 2005; Thomson & Ring 2006). The Simav detachment along the northern edge of the northern submassif (Is¸ik et al. 2003; Ring & Collins 2005) and the Alasehir and Bu¨yu¨k Menderes (or Gu¨ney) detachments along the northern and southern edges (respectively) of the central submassif (Fig. 2; Gessner et al. 2001) played a major role in the exhumation of the Menderes massif as a core complex.

Cenozoic granitoids in the Menderes Massif We have examined the occurrence, structure and geochemistry of the Egrigo¨z, Koyunoba and Baklan granitoids in the northern submassif and the Salihli, Turgutlu and Alasehir granitoids in the central submassif (Fig. 2) in order to document the nature and petrogenesis of syn-extensional magmatism during the exhumation of the Menderes core complex. Granitoid plutons in the Menderes massif are mainly exposed in the footwalls of major detachment faults and show structures and textures associated with extensional deformation. Collectively, we group these plutons under a descriptive name of Menderes Granitoids (MEG).

Field relations and structure of the MEG The NNE-trending Egrigo¨z and Koyunoba granitoids are intruded into the Precambrian– Palaeozoic gneiss, schist, amphibolite and gneissic mylonite in the footwall of the Simav detachment in the northern submassif and are overlain by Neogene volcanosedimentary rocks of the Akdere basin (Is¸ik et al. 2004; Ring & Collins 2005). They are composed of granite, granodiorite and monzonite, with less common diorite and monzodiorite, and are cut by fine-grained dykes of similar compositions. The majority of the plutonic rocks are medium- to coarse-grained and isotropic. They become foliated and mylonitic toward the Simav detachment fault to the north and west, with a preferred alignment of biotite, quartz and feldspar defining a hightemperature foliation plane in a ductile shear zone. S-C fabrics, biotite fish structures, asymmetric

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porphyroclasts and oblique quartz grain-shape foliation all consistently suggest top-to-the-NW sense of shear, indicating extensional deformation parallel to the Simav detachment surface (Is¸ik et al. 2003; Ring & Collins 2005; Thomson & Ring 2006). The timing of the mylonitic shear zone development along the northern edge of these plutons and their intrusion into the northern Menderes massif are nearly coeval as constrained by the recent geochronological studies. 40Ar/39Ar dating of metamorphic muscovite grains from a mylonitic gneiss in the extensional shear zone along the Koyunoba pluton suggests that mylonitic deformation of the host rock occurred around 22.86 + 0.47 Ma (Is¸ik et al. 2004). Ring & Collins (2005) obtained secondary ion mass spectrometry (SIMS) U– Th–Pb zircon ages of 20.7 + 0.6 Ma and 21.0 + 0.2 Ma from the Egrigo¨z and Koyunoba granitoids, respectively. These intrusion ages are within error of the 40 Ar/39Ar biotite cooling ages of 20.19 + 0.28 Ma from the Egrigo¨z granitoid (Is¸ik et al. 2004). These geochronological data, combined with the structural observations, demonstrate the syn-extensional nature of the plutons and possibly their intrusion at shallow depths resulting in rapid cooling during exhumation (Ring & Collins 2005; Thomson & Ring 2006). The c. 15 Ma volcanic interlayers in the Akdere basin sedimentary srata, which rest on the mylonitic gneiss and granitoids along the Simav detachment (Ercan et al. 1997), suggest that the northern Menderes massif and the syn-extensional Egrigo¨z and Koyunoba plutons were exhumed to the surface by the early–middle Miocene. The Baklan granitoid to the SE of the Egrigo¨z and Koyunoba plutons (Fig. 2) is intrusive into the Palaeozoic marble, schist and quartzite of the Menderes massif and the tectonically overlying Muratdagi ophiolitic melange (Aydogan et al. 2008). The lower Miocene Ku¨rtko¨yu¨ Formation consisting of sandstone and gravel deposits rest unconformably on all these units, including the Baklan granitoid, indicating that the pluton was exposed at the surface by the end of the early Miocene. The majority of the Baklan pluton rocks are made of medium- to coarse-grained granodiorite composed of K-feldspar, plagioclase, quartz, biotite and hornblende. Monzodioritic to monzogabbroic microgranular enclaves are common in the granodioritic host rocks. K –Ar whole-rock dating of the granodiorite has revealed ages between 17.8 + 0.7 Ma and 19.4 + 0.9 Ma (Aydogan et al. 2008). The Turgutlu, Salihli and Alasehir granitoids farther south occur along an approximately east – west-trending belt in the footwall of the Alasehir detachment in the central submassif and are intrusive into the metasedimentary rocks and schists of the Menderes massif (Fig. 3). The shear zone beneath the detachment surface includes mylonitized

metamorphic and granitoid rocks that display welldeveloped foliation and stretching lineation in the lower sections and microbreaccia, breccia, cataclasite, foliated cataclasite and pseudotachylite ¨ ner & Dilek toward the top (Is¸ik et al. 2003; O 2007). The nearly 100 m thick cataclastic shear zone beneath the detachment surface contains S-C fabrics, microfaults, Riedel shears and shear bands, all consistently indicating top-to-the-north-NE ¨ ner & Dilek 2007). The majority shearing (Fig. 4; O of all three granitoids are composed of isotropic granodiorites, which become increasingly mylonitic upward into the detachment shear zone. The mylonitic foliation in these deformed granitoids is defined by the alignment of biotite and feldspar porphyclasts (Fig. 5c), whereas the lineation is marked by stretched quartz and preferred orientations of feldspar and biotite grains plunging to the north–NE at ¨ ner & Dilek 2007). 12 –148 (Is¸ik et al. 2003; O Kinematic indicators in the mylonitic granitoids and their host metamorphic rocks include S-C fabrics, asymmetric porphyroclasts, biotite fish, fractured and displaced grains and asymmetric enclaves that consistently show top-to-the-north– NE normal sense of shearing (Is¸ik et al. 2003; ¨ ner & Dilek 2007). The similar orientation of O the mylonitic foliation and stretching lineation (Fig. 4), the same top-to-the-north– NE normal shear sense and a corresponding retrograde greenschist-facies metamorphic overprint in both the deformed granitoid and its host metamorphic rocks indicate that the plutons and the Menderes metamorphic rocks were affected by the same extensional deformation. The progression from relatively undeformed isotropic granodiorite at depth to mylonitic-ultramylonitic plutonic rocks within the detachment shear zone at the surface (Fig. 4) further show that these plutons are syn-extensional intrusions (Hetzel et al. 1995; Is¸ik et al. 2003; ¨ ner & Dilek 2007). O The oldest sedimentary units overlying the Alasehir detachment surface and the granitoids are the lower –middle Miocene shale and limestone of lacustrine and fan-delta facies (Gerentas and Kaypaktepe Units). Stratigraphically upward these rocks are overlain by the upper Miocene Acidere Formation and the Plio-Pleistocene Go¨bekli, Yenipazar, Asartepe and Erendali Formations ¨ ner & Dilek 2007). Several major (Figs 4 & 5; O unconformities exist in these Miocene– Pleistocene strata that likely developed as a result of extensional tilt-block faulting during the evolution of the Alasehir supradetachment basin.

Geochronology of the MEG The geochronology of both the Salihli and Turgutlu granitoids and their metamorphic host rocks provide

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Fig. 3. Geological map of the Menderes massif and its environs in western Anatolia. The Menderes massif consists of several nappe systems (Bozdag, C¸ine and Bayindir) that are stacked up along north-directed thrust sheets in the field area. The Alasehir granitoid (AG), Salihli granitoid (SG) and Turgutlu granitoid (TG) are intrusive into different basement units in the Menderes massif.

significant constraints on the timing of the extensional deformation and the synchronous magmatism in the footwall of the central Menderes massif. Igneous biotites from the Turgutlu and Salihli granodiorites yielded 40Ar/39Ar cooling ages of 13.2 + 0.2 Ma and 12.2 + 0.4 Ma, respectively (Hetzel et al. 1995). U –Pb monazite ages of 16.1 + 0.2 Ma from the Turgutlu and U –Pb allanite ages of 15.0 + 0.3 Ma from the Salihli granodiorite date the crystallization ages of these two granitoid bodies as the early –middle Miocene (Glodny & Hetzel 2007). These crystallization ages are close to the matrix monazite ages of 17 + 5 Ma (Catlos & C ¸ emen 2005) from metamorphic rocks in the eastern part of the Alasehir detachment that are interpreted to record tectonic extension. More recently, Catlos et al. (2008) reported in situ Th –Pb ion microprobe monazite ages of 21.7 + 4.5 Ma to 9.6 + 1.6 Ma (+1s) from the

Salihli and Turgutlu granitoids and 31.5 + 2.7 Ma to 22.8 + 2.4 Ma (+1s) from garnet-bearing schists from the Bozdag nappe. These new ages demonstrate an older exhumation history of the middle crustal rocks and their syn-extensional plutons in the central Menderes massif, going back to the early Miocene and possibly to the Oligocene.

Geochemistry We present new major and trace element analyses and isotope data from the Salihli granitoid occurring on the southern shoulder of the Alasehir graben (Figs 2 & 3). These new data, combined with the previously published geochemical data from the Salihli, Turgutlu, Egrigoz and Baklan plutons, are used to constrain the magma sources and tectonomagmatic evolution processes of the MEG. We

204 Y. DILEK ET AL. Fig. 4. Geological cross-section from the northern part of the central Menderes core complex showing the spatial relations between the high-grade metamorphic rocks, the Salihli pluton, the cataclastic zone of the detachment fault and the Upper Miocene-Pleistocene supradetachment sedimentary strata (tilted southward into the detachment fault). The cataclastic zone associated with the Alasehir detachment is up to 150 m in thickness and encompasses the exposed upper surface of the Salihli granitoid. Note the decreasing strain structurally down-section into the pluton. Outcrop photos 1 and 2 show the internal fabric in a relatively less-deformed and a highly-deformed granodiorite from two different elevations in the pluton. Both the cataclastic zone and the Salihli granitoid rocks display similar extensional fault and foliation geometries. South-dipping foliation in both units represents the extensional fabric in south-tilted fault blocks.

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205

Fig. 5. Field photos of the Salihli granitoid in the Menderes core complex. (a) Gently north-dipping slope in the background represents the Alasehir detachment surface, overlain by the Upper Miocene Acidere Formation. The surface in the foreground shows the cataclastic zone of the detachment surface. View to the East. (b) The Alasehir detachment dips gently to the North and overlain by the red fluvial clastic rocks of the Upper Miocene Acidere Formation. Deformed plutonic rocks of the Salihli granitoid are seen along the detachment surface in the foreground. View to the NW. (c) North-dipping mylonitic foliation and banding in the Salihli plutonic rocks beneath the cataclastic zone.

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also compare the geochemical features of the MEG with those granitoids that are emplaced into other metamorphic core complexes in the northern (Rhodope massif) and central (Cyclades) Aegean Province.

Analytical techniques Multi-element concentration was determined by using polarized energy dispersive XRF. The spectrometer used in this study is the Spectro XLAB 2000 PEDXRF that is equipped with an Rh anode X-ray tube and 0.5 mm Be side window and housed in the Department of Geological Engineering at the University of Ankara (Turkey). The detector of spectrometer is Si (Li), cooled by liquid N2, with a resolution of ,150 eV at Mn Ka, 5000 cps. The spectrometer was calibrated with two standard rocks, G01-MA-N and K03-MRG-1, Canadian Certified Reference Materials and Centre de Recherches Pe´trographiques et Ge´ochimiques (CRPG) of France, based on the certificated concentrations of the elements under investigation. The samples were ground into fine powder in an agate mortar that was sieved to pass through of 200 mm and then pressed into thick pellets of 32 mm diameter using wach as blinder. USGS standards, GEOL, GBW 7109 and GBW-7309 Sediment were equally pressed into pellets in a similar manner as the samples and these standards were used for quality assurance (La Tour 1989; Johnson et al. 1999). Total analysis time for each additional element was 30 minutes. Sr and Nd isotope ratios were determined by thermal ionization mass spectrometry (TIMS) (Thermo Quest Finnigan MAT 261) at the University of Texas at Dallas.

Major and trace element characteristics Major and trace element compositions of the MEG are given in Table 1. Different intrusive units of the MEG (Salihli, SG; Turgutlu, TG; Egrigo¨z, EG; Baklan granitoids, BG) show moderate variations of major element compositions, with their silica content ranging from 62.5 to 78.2 wt%. The MEG are composed mainly of tonalite, granodiorite and granite (de la Roche et al. 1980; Fig. 6), consistent with their modal mineralogy. The A/CNK [Al2O3/(CaO þ Na2O þ K2O) molecular ratio] values range between 0.81 and 1.27. The vast majority of the MEG are metaluminous and slightly peraluminous; rocks having peralkaline compositions are absent among the analysed samples (Fig. 7). The least-evolved members are predominantly metaluminous (tonalite and hornblende– biotite granodiorite with A/CNK ,1), whereas some more-evolved biotite granodiorite and two-mica granodiorites exhibit slightly to mildly

peraluminous signatures (A/CNK ¼ 1.1 –1.27). However, the metamorphic basement rocks into which these granitoids were emplaced are strongly peraluminous (A/CNK ¼ 1.45–1.91) (Fig. 7). The MEG are all subalkaline in nature and display a calc-alkaline trend in an AFM ternary diagram (Irvine & Baragar 1971; Fig. 8a). The majority of the MEG belong to the high-K calc-alkaline series, although a few samples belong to the shoshonitic and medium-K series (Fig. 8b). Their K2O/Na2O ratio ranges between 0.68 and 2.67 and they have moderate Mg-numbers (average: 35–57) (Fig. 9). The major element variations of the MEG define linear trends as seen in Figure 9. Their TiO2, MgO, Al2O3, FeO, CaO and P2O5 contents decrease and K2O contents and A/CNK, K2O/Na2O ratios increase with increasing SiO2 contents (Fig. 9). The trace element compositions also show moderate variations, some of which (e.g. Sc, Zr, Rb and Sr) correlate well with the SiO2 contents (Table 1; Fig. 10). Samples with SiO2 . 70 wt% mostly have 102–250 ppm Rb, 241–279 ppm Sr and 103–154 ppm Zr, whereas samples with SiO2 , 70 wt% have 95 –144 ppm Rb, 216– 372 ppm Sr and 108–239 ppm Zr. Other trace elements display more scattered variations (e.g. Y, La, Ce ) with SiO2 (not shown here). There is a considerable overlap among all four groups of the plutons (SG, TG, EG, BG) for inter-element ratios, such as Zr/Nb, La/Nb, Rb/Nb and Ce/Y (Fig. 9). This similarity is supported by the abundance and the shape of trace element patterns on the PM-normalized and chondrite-normalized REE diagrams (Figs 10a & 11a). Primitive mantle (PM)-normalized trace element patterns of representative rock samples from the MEG, Cyclades and Rhodope granitoids and metamorphic basement rocks of the Menderes massif are shown in Figure 10. Elemental patterns of upper (UC), middle (MC) and lower continental crust (LC) are also plotted in Figure 10a for the purpose of comparison (UC, MC, LC values are taken from Taylor & Mc Lennan 1985). The MEG display enrichment in LILE (K, Rb,Cs), in some HFSE (Th, U) and in Pb over LREE and MREE and show strong negative anomalies in Ba, Nb, Sr, P and Ti. These depletions are more pronounced in two-mica granitoids. Trace element patterns of the MEG are similar to the trends displayed by the granitoids in the Cyclades, such as the plutons on Mikanos, Naxos, Delos and Ikeria (Pe-Piper et al. 1997, 2002; Pe-Piper & Piper 2001; Alther & Siebel 2002) and in the Rhodope massif (Fig. 10b; Christofides et al. 1998). These patterns are also partly similar to those of the metamorphic basement rocks of the Menderes massif (Fig. 10c; Catlos et al. 2008). The REE distributions of the MEG show

Table 1. Major and trace element compositions and Sr and Nd isotopes of selected samples from the Salihli granitoid in the Menderes metamorphic massif, western Turkey 01DEG07 02DEG07 03DEG07 04DEG07 05DEG07 06DEG07 07DEG07 08DEG07 09DEG07 10DEG07 11DEG07 12DEG07 13DEG07 14DEG07 15DEG07 16DEG07 17DEG07 18DEG07 19DEG07 63.92 19.11 3.89 0.07 1.71 3.94 1.75 4.66 0.67 0.15 0.58 100.44

Ba 578.9 Sr 216 Y 19.2 Zr 220.8 Co 24 Zn 43 Ga 26.4 Ge 1.1 Rb 162.2 Nb 14.7 Sn 8.2 Cs 9 La 25.7 Ce 55 Hf 2.5 Ta 3.6 Tl 1.4 Pb 56.4 Bi 1 Th 12 U 7.6 87 Sr/86Sr 0.71096 143 144 Nd/ Nd 0.51223

70.25 14.66 3.1 0.06 1.14 2.77 3.04 3.46 0.47 0.16 0.82 99.93

68.43 14.9 3.26 0.06 1.57 3.77 3.58 3.13 0.48 0.16 0.57 99.92

67.45 14.49 4.18 0.07 2.09 3.93 3.04 3.21 0.61 0.17 0.64 99.87

69.17 14.76 3.29 0.06 1.62 2.91 4.27 3.09 0.46 0.15 0.15 99.93

67.77 15.21 3.63 0.07 2.05 4 3.18 2.78 0.53 0.18 0.53 99.92

67.87 14.02 3.96 0.07 2.24 3.84 4.33 2.91 0.59 0.15 0.63 100.61

625.8 243.7 21.2 197.3 44.9 57.6 20.9 1.8 139.9 15.2 4.5 22.3 78.3 130.2 2.8 4.2 2.3 64.7 1.7 14.2 10.8 n.d. n.d.

633.9 326.8 17.5 193.2 37.7 51.9 18.6 1.3 110 14.6 5.4 14.3 53.3 92.6 2.9 10.7 2.7 50 0.9 13.7 8.2 n.d. n.d.

511.5 297.7 24.5 205.3 41.1 61.7 20.7 1.33 121.1 19.9 5.6 17.7 60 101.8 4.2 9.1 2.5 38 0.9 15.3 7.7 n.d. n.d.

556.1 302.4 16.5 203.3 59.2 63.7 19.8 2.4 112.7 14.5 4.6 21.7 73 113.2 2.9 4.2 1.4 45.7 1.1 17.7 6.8 n.d. n.d.

531.3 350.7 16.9 203.8 63.1 55.4 18.5 1.93 92.2 16.2 4.2 16.5 80.7 119.3 3 4.9 2.4 39.3 1.1 14.1 7.7 n.d. n.d.

396.1 306.3 17.3 170.5 44.9 60.6 20.3 2.6 122.3 13.6 6 22.8 43.1 63.4 3.1 4.5 1.3 44.2 0.9 9 18.3 n.d. n.d.

68.58 15.33 2.98 0.06 1.24 2.56 4.05 3.8 0.45 0.13 0.75 99.93 607.6 302.4 19.3 154.4 50.4 52.4 20.5 1.3 120.4 13.9 6.5 17.3 37.7 70.3 3 4.2 1.3 54 0.7 11.4 7.3 0.71141 nd.

69.08 14.31 3.18 0.06 1.48 3.43 3.67 3.15 0.5 0.19 0.82 99.87

65.85 15.26 3.98 0.07 1.99 4.23 3.4 3.71 0.55 0.19 0.86 100.08

65.08 15.72 4.31 0.07 2.39 4.11 3.5 2.99 0.64 0.22 0.82 99.85

597.5 285.7 13.6 202.9 41.9 55.6 19.5 2.4 95.6 15.1 1.2 22.5 74.9 129.8 2.8 3.7 1.9 44 0.7 13.4 6.8 n.d. n.d.

474.5 349.4 17.8 218 50 55.2 21.4 2.4 110.5 16.5 5 7.8 41.6 78.9 3.2 4.3 1.1 40.9 0.7 15.4 15.6 n.d. n.d.

477.6 355.7 18.6 227.5 42.6 64.1 22 1.32 105 16.8 3.5 3.52 18.9 35.5 3.3 4.6 1.7 41.9 1.2 14.2 15.3 n.d. n.d.

63.92 14.94 5.47 0.13 3.04 5.02 2.83 2.95 0.58 0.16 0.85 99.89 472.8 319.4 34.4 173.6 47.6 79.7 18.8 1.3 111.7 14.7 5.4 16.3 42.8 80.6 4.7 9.9 1.4 38.8 1.2 9.5 8 0.7109 0.51222

66.51 15.07 4.18 0.08 2.39 4.16 3.01 2.99 0.57 0.2 0.72 99.88

67.66 14.73 3.88 0.07 1.85 3.46 3.14 3.48 0.57 0.17 0.85 99.87

68.29 14.35 3.81 0.07 1.63 3.97 2.69 4.1 0.57 0.18 0.86 100.51

67.06 14.73 3.97 0.07 2.3 4.03 3.01 3.24 0.55 0.19 0.73 99.89

67.67 14.62 3.77 0.06 1.64 3.59 3.2 3.59 0.53 0.17 0.95 99.8

66.62 15.25 4.05 0.07 2.02 4.13 3.42 2.78 0.62 0.2 0.71 99.88

69.58 14.33 3.2 0.06 1.86 3.04 3.01 3.71 0.46 0.18 0.93 100.35

608.6 346.7 14.7 210.5 46.1 60.1 20.9 1.3 101.5 11.9 3.4 29.1 96.5 150 3.5 3.8 1.7 42.6 0.5 13 9.5 n.d. n.d.

624.5 292.4 16.8 184.8 37.2 67.9 21.7 2.3 135.4 18.7 3.1 26.1 61 107.6 4.6 5.7 1.4 48.7 0.4 10.4 8.2 n.d. n.d.

639.5 291.9 16.3 214.3 41 60.7 20.7 1.4 139.2 15.9 4.1 20.5 39 81 3.8 4 1.1 53.3 0.7 10.8 8.3 n.d. n.d.

581.9 350.9 16.4 212.8 44.3 54.3 20.4 1.3 97.9 12 3.9 19.2 78.5 122.5 2.5 3.9 2.4 38.1 1.5 12.9 9.2 n.d. n.d.

686.9 323.8 18.1 191.3 30.6 46.3 19.1 1.3 95.8 13.4 3.2 23.1 117.5 191.4 3.1 3.3 1.7 42.1 0.8 10.3 7.5 n.d. n.d.

481.9 312.8 23.5 217.4 64.7 61.2 21 1.3 91 18.3 4.1 16.9 65.7 111.7 3.1 4.5 1.9 38.2 1.6 15.6 7.8 n.d. n.d.

995.5 296.4 11.5 207.2 46 52.4 21.5 1.3 113.3 16 0.9 23.6 79.3 126.7 3 4.2 1.4 53.7 0.6 12.1 8 n.d. n.d.

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SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total

(Continued ) 207

208

Table 1. Continued 20DEG07 21DEG07 22DEG07 23DEG07 24DEG07 25DEG07 26DEG07 27DEG07 28DEG07 29DEG07 30DEG07 31DEG07 32DEG07 33DEG07 34DEG07 35DEG07 36DEG07

Ba Sr Y Zr Co Zn Ga Ge Rb Nb Sn Cs La Ce Hf Ta Tl Pb Bi Th U 87 Sr/86Sr 143 Nd/144Nd

69.86 14.2 3.23 0.06 1.39 3.12 3.21 3.17 0.48 0.19 0.94 99.85

69.44 14.22 3.82 0.06 1.35 3.14 3.07 3.98 0.42 0.15 0.91 100.57

68.4 14.21 3.94 0.08 1.88 3.79 3.06 2.83 0.56 0.17 0.94 99.87

69.47 14.34 3.56 0.06 1.63 3.62 2.79 2.94 0.56 0.16 0.79 99.93

555.9 282.4 17.9 222.7 45 58.4 21.9 1.1 104.8 20.2 2 3.9 8.5 135.4 2.9 4 1.2 50.6 1 10.3 10.6 n.d. n.d.

652.4 274 15.7 172.6 29.3 50.6 20.8 0.3 144.2 14.8 4.3 17.4 62.2 108.1 3.2 4.3 2.3 58 0.9 14.4 7 n.d. n.d.

454.4 282 16.9 196.6 32.5 68.9 21.3 1.9 120.1 17.1 10.1 17.4 44.1 80.3 3.3 3.8 1.6 48.4 1 13.5 6.7 n.d. n.d.

487.2 292.5 20.9 206.3 39.4 55.8 18.1 1.34 121.1 18.1 3.9 21.2 39.6 67.2 3.1 7.6 1.7 42.9 1.1 8.9 7.3 n.d. n.d.

71.52 14.25 2.21 0.05 0.97 2.97 3.25 3.69 0.3 0.11 0.82 100.14 534 279.8 12.7 131.4 52 41.7 19.6 0.8 102 8 3.5 13.6 42.1 64 3.1 4.3 1.6 60.3 2 8.1 9 0.71132 0.5122

66.48 15.49 4.12 0.07 2.04 3.95 3.28 2.98 0.64 0.22 0.62 99.89

67.48 14.95 3.73 0.06 1.93 3.68 3.2 3.36 0.55 0.2 0.76 99.91

66.25 15.34 4.28 0.07 2.18 4.19 3.27 2.7 0.61 0.24 0.69 99.82

609.9 372.8 18.5 222.2 45.2 61.6 21.6 2 101 15.4 4.1 15.3 60.9 111.9 3.3 3.7 1 40.7 1 13.8 6.6 n.d. n.d.

777.9 352.6 17.2 202.6 41.8 56.4 20.3 0.34 102.3 14.8 1.7 18.5 73.7 113.4 3.3 4 1.8 44.6 1 12.6 7.2 n.d. n.d.

511.1 370 19.9 218 12 58.2 21.4 1.8 93.7 16.8 2.9 3.54 10.9 88.7 4.5 4.2 1.1 39.4 1.4 13 20.8 n.d. n.d.

65.92 15.23 4.68 0.08 2.31 4.54 2.85 2.38 0.7 0.24 0.82 99.76 620 384 17.8 239.4 28.7 60.9 19.4 1.32 70.4 15.5 5 9.4 64.3 108.9 3.2 4.8 1.1 41.4 1.3 12.5 7.3 0.71104 0.51223

66.96 15.25 3.96 0.07 2.04 4.1 3.34 2.77 0.59 0.21 0.62 99.91

71.66 13.44 2.89 0.06 1.18 2.69 2.86 3.91 0.43 0.17 0.65 99.94

67.34 15.03 3.47 0.07 1.84 3.88 3.34 3.3 0.54 0.18 0.9 99.89

607.3 369.7 21.2 231.7 41.7 58.1 19.6 0.63 86.2 15.7 4.7 3.21 11.3 104.6 3 4 1.6 39.5 1.1 15.1 10.6 n.d. n.d.

564.8 256.8 21.9 154.6 58.7 50.6 20.2 1.3 132.7 12.8 4 22.4 59.5 94.6 3 3.4 1.2 50.7 1.7 12.3 8 n.d. n.d.

731 337.8 19.6 196.9 59 50.2 19.8 1.7 103.7 12.3 3.2 20.7 68.4 95.5 3.1 4.1 1.5 44.8 1 10.5 10.7 n.d. n.d.

67.1 15.32 3.81 0.07 1.5 3.76 4.22 3.27 0.47 0.16 0.72 100.4 601 319 22.5 188.4 42.8 57.6 19.5 1.1 108.9 23.5 6.9 3.11 20.3 32.5 3.2 8.1 2.2 45.4 1.1 16.2 16.1 0.71106 0.51223

66.21 15.22 4.18 0.07 2.53 3.66 3.43 2.99 0.68 0.2 0.68 99.85

66.48 15.71 3.83 0.06 2.13 3.41 3.51 3.38 0.57 0.19 0.62 99.89

460.8 338.3 24 215.8 45.4 61 20 2.3 108 17.9 7.6 12.3 47.1 85 4.3 5.2 1.8 41.8 1.2 12.3 10.1 n.d. n.d.

618.3 345.7 23.5 193.3 56.7 56.9 21.1 0.67 116 16.4 1 33.4 103 166.2 3 4 1.5 43.4 0.8 11.6 6.9 n.d. n.d.

73.95 13.1 1.77 0.03 0.97 2.96 2.75 4.1 0.28 0.09 0.34 100.35 376.3 269.2 12.4 103.4 46.6 29.9 15.8 1.78 85.6 12.5 1 3.5 7.4 12.2 2.8 4.1 1.5 54.2 0.8 14.9 20.6 0.71103 0.51224

67.17 15.06 3.8 0.07 1.87 4.01 3.78 2.82 0.57 0.19 0.54 99.88 481.8 326.2 21.9 223.9 39.7 57.4 21.2 0.32 95.3 17.3 4.4 16.8 69.1 124.8 3.3 4.9 1.5 42.6 1.2 15.5 10.3 n.d. n.d.

Y. DILEK ET AL.

SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total

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209

Fig. 6. Chemical variation in the MEG shown on a plot by de la Roche et al. (1980). Data sources include Salihli granitoid: this study and Catlos et al. (2008); Turgutlu granitoid and metamorphic basement rocks: Catlos et al. (2008); Baklan granitoid: Aydogan et al. (2008); Egrigo¨z ¨ zgenc¸ & ˙Ilbeyli (2008), Is¸ik granitoid: Akay (2008), O et al. (2004) and Delaloye & Bingo¨l (2000).

moderately fractionated REE patterns with LaN/ YbN ratios of 4.2–23.3 (average: 7– 14), regardless of the rock types (Fig. 11a). Samples from the Salihli pluton have minor Eu anomalies with Eu*/Eu values of 0.68 and 0.88, whereas others (TG, EG, BG) have medium to strong negative Eu anomalies ranging from 0.22 to 0.55 that show similar Eu*/Eu ratios to those of the Cyclades granitoids (Fig. 11a, b). The magnitude

Fig. 8. (a) AFM diagram diagram of Irvine & Baragar (1971); (b) K2O v. Na2O diagram of the MEG using the classification scheme of Peccerillo & Taylor (1976). See Figure 6 for the data sources.

of the negative Eu anomalies of the MEG samples increases with the increasing SiO2 content.

Petrogenesis of the Menderes granitoids Previous studies and interpretations

Fig. 7. The Shand’s index diagram for the MEG (Shand 1927). A/CNK: molar Al2O3/(CaO þ Na2O þ K2O); A/NK: molar Al2O3/(Na2O þ K2O). Fields for I-and S-type granitoids are taken from Chappell & White (1974, 1992). See Figure 6 for the data sources.

The origin of the MEG has been a subject of various studies. Delaloye & Bingo¨l (2000) have suggested that the western Anatolian plutons, including the Salihli granitoid, originated from the Palaeocene and younger magmatism associated with the Hellenic subduction zone. Is¸ik et al. (2004) have reported that the syn-extensional Egrigo¨z and Koyunoba plutons in the footwall of the Simav Detachment were emplaced in the early stages of continental extension in the Aegean province. These granitoids are hybrid in nature with dominantly upper crustal compositions similar to the coeval Oligo-Miocene granitoids in the central Aegean Sea region. Haso¨zbek et al. (2004) and Akay (2008) argued for a

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Fig. 9. Harker diagrams of the MEG illustrating the variations of major oxides and trace elements with SiO2. See Figure 6 for the data sources.

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211

Fig. 11. Chondrite-normalized REE patterns for the MEG (a) and Cyclades granitoids (b). Chondrite normalizing values are from Boynton (1984). Data sources for Cyclades granitoids are from Pe-Piper et al. (2002), Pe-Piper & Piper (2001) and Altherr & Siebel (2002). Upper crust (UC), lower crust (L) and middle crust (MC) values are from Taylor & McLennan (1985).

Fig. 10. PM (Primitive mantle) normalized multi-element patterns for the MEG (a), Cyclades and Rhodope granitoids (b) and metamorphic basement rocks of the Menderes massif (c). PM normalizing values are from Sun & McDonough (1989). Data sources for the Cyclades: Pe-Piper & Piper (2001), Pe-Piper et al. (2002), Alther & Siebel (2002); Rhodope granitoids: Christofides et al. (1998). See Figure 6 for the data sources for the other granitoids.

hybrid magma source produced under compres¨ zgenc¸ & sional regime for the same granitoids. O ˙Ilbeyli (2008) have proposed that the Egrigo¨z pluton formed by partial melting of mafic lowercrustal rocks during post-collisional extensional tectonics in the region. Aydogan et al. (2008) have argued that parental magmas of the Baklan pluton were produced by partial melting of a juvenile lower crust, and that the underplated mantle-derived basaltic magmas, which provided the necessary heat for partial melting, had chemical and isotopic signatures similar to those of an enriched mantle. Catlos et al. (2008) have argued that the trace-element geochemical features of the Salihli and Turgutlu granitoids are consistent with their continental arc origin and that their magmas were produced under a compressional regime above the north-dipping Hellenic subduction zone. The existing interpretations of the melt source and the magmatic evolution of the MEG are, therefore, varied.

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Melt sources and evolution of the MEG The Salihli, Turgutlu, Egrigo¨z and Baklan plutons are all intrusive into the high-grade metamorphic rocks of the Menderes core complex. They show similar major and trace element characteristics and overlapping Zr/Nb, La/Nb, Rb/Nb and Ce/Y, suggesting that their melt source(s) were similar (Figs 9, 10 & 11a). The A/CNK molecular ratios of the MEG between 0.81 and 1.27 indicate that these plutons are made predominantly of metaluminous, I-type ¨ zgenc¸ & I˙lbeyli 2008; Aydogan et al. granitoids (O 2008) and slightly to mildly peraluminous, rare S-type (Is¸ik et al. 2004), two-mica granitoids (White & Chappell 1977; Chappell & White 1992). The MEG samples display similar trace elemental patterns to middle-upper continental crust and metamorphic basement rocks (Fig. 10a, c) that were likely inherited from crustal melts of variable sources and compositions. However, the higher A/CNK values and lower Mg-numbers of the metamorphic basement rocks in comparison to the MEG are not consistent with derivation of the MEG magmas from these basement units alone. Given that the metamorphic basement rocks are strongly peraluminous (Fig. 7), large amounts of crustal component contribution into the magmas would increase their A/CNK ratio and would cause the formation of strongly peraluminous S-type granitoids; yet, we do not see these features in the MEG. Thus, the metaluminous I-type character of the MEG eliminates the metapelitic rocks as suitable source material and points to, instead, an igneous protolith such as metabasalt, juvenile K-rich basaltic underplate and/or mantle rocks (Roberts & Clemens 1993; Tepper et al. 1993; Pearce 1996; Patino Douce & McCarthy 1998; Von Blanckenburg et al. 1998; Ashwal et al. 2002; Altherr & Siebel 2002). Experimental studies suggest that hydrous melting of basalts or amphibolites could yield tonalitic magmas, which evolve toward granodioritic to granitic compositions by crustal interaction and/or fractional crystallization (Wyllie 1984; Rapp & Watson 1995; Petford & Gallagher 2001). Although some chemical features [such as (Na2O þ K2O)/ (FeO þ MgO þ TiO2) and CaO/(FeO þ MgO þ TiO2)] of the MEG, as shown in Figure 12 (a & b) appear to be derived from metabasalts (Patino Douce 1996, 1999), partial melts of metabasalts are characterized by relatively high contents of Na2O and low Mg numbers (Fig. 12c, d), regardless of the degree of partial melting. These features are not displayed by the MEG samples. Furthermore, Roberts & Clemens (1993) and Ashwal et al. (2002) have argued that metabasaltic rocks are not suitable source rocks for the generation of high-K,

calc-alkaline, I-type granitoids because such mafic rocks contain low-K2O and insufficient incompatible elements to form appreciable volumes of granitic melts. Therefore, we infer that high-K, calc-alkaline and incompatible element-enriched nature of the MEG is inconsistent with derivation of their magmas solely from melting of metabasalts. Intermediate to felsic products of the MEG show PM-normalized multi-element patterns (e.g. enrichment in Rb, Th and K and negative anomalies in Ba, Sr, P and Ti), suggesting their possible derivation from basaltic magmas through crystal fractionation (Fig. 10a). Trace-element patterns of this group are consistent with derivation of their magmas from an incompatible element-enriched source, as evidenced by negative Ta and Nb anomalies, enriched LREE, and low Rb/Sr ratios. These features of the MEG are similar to those of igneous rocks forming at convergent margin settings (Thorpe et al. 1982; Davidson et al. 1991; Pearce & Peate 1995). High incompatible element abundances (e.g. K, Rb, Nb and Ba) and inter-element relationships (e.g. Ce/Y, Zr/Ba, Th/Yb and Ba/Nb ratios; Figs 9 & 10a) of the MEG indicate a subduction-enriched, heterogeneous sub-continental lithospheric mantle source (see Pearce et al. 1990; McCulloch & Gamble 1991; McDonough 1990; Thirlwall et al. 1994; Pearce & Peate 1995). Moreover, the (La/Yb)n and (Gd/Yb)n ratios of the MEG are in the ranges 6.5–14.5 (Fig. 11a) and 1.3–2.1, respectively, that are consistent with derivation from lithospheric mantle melts (Thirlwall et al. 1994). Partial melting model suggested by Thirlwall et al. (1994) shows that the effects of partial melting were more important than fractional crystallization in controlling the compositional variations within the MEG (Fig. 13). The subduction-related enrichment of the western Anatolian lithospheric mantle may have been an artifact of either arc-derived magmas or a subduction component inherited from earlier convergent margin events. Source enrichment through previous subduction events in the region has been suggested for the western Anatolian plutons and related volcanism by some authors (Seyitog˘lu et al. 1997; Genc¸ & Yılmaz 1997; Yılmaz & Polat 1998; Altunkaynak & Yılmaz 1998; Yılmaz et al. 2000, 2001; Aldanmaz et al. 2000; Ko¨pru¨bas¸i & Aldanmaz 2004; Altunkaynak & Dilek 2006; Altunkaynak 2007; Dilek & Altunkaynak 2007). However, although subduction-induced mantle metasomatism can account for enriched source characteristics of the MEG rocks, the Nb/La versus Ba/Rb variations observed in these samples (Fig. 14) cannot be explained solely by this mechanism. The vertical continental crust versus N-MORB-OIB mantle trend indicate mixing of a mantle derived magma with a crustal component,

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Fig. 12. Chemical compositions of the MEG. Outlined fields are adapted from Altherr & Siebel (2002 and references therein) and illustrate compositions of partial melts obtained from experimental studies of dehydration melting of Metabasalts (MB), Metagreywackes (MGW), Metapelites (MP) (Rapp & Watson 1995; Patino Douce 1996, 1999; Patino Douce & McCarthy 1998; and references therein). See Figure 6 for the data sources.

rather than the sole influence of subductiongenerated fluids (Wang et al. 1999; Tatsumi et al. 1986; Peccerillo 1999; Marchev et al. 2004). A critical evaluation of possible contamination by crustal materials is, therefore, necessary. If the MEG magmas acquired their chemical compositions by assimilation of any crustal material as discussed above, it is more likely that these assimilants were composed of lower to middle crustal components rather than the upper crustal host basement metamorphic rocks alone.

Sr – Nd isotopes Nd–Sr isotopic compositions of the Salihli (Table 1) and Baklan granitoids (Aydogan et al.

2008) are consistent with a hybrid origin of their magmas. The Baklan granitoid has low 87Sr/86Sr(i) (0.70331 –0.70452) and high 143Nd/144Nd(i) (0.512305–0.512336) ratios and negative 1Nd(t) (25.0 to 25.6) values that are compositionally similar to EMI-type source. These isotope values point to the production of the Baklan granitoid magmas from interaction of mantle-derived melt with lower crustal amphibolites or granulites (Aydogan et al. 2008). However, the Salihli granitoid displays higher 87Sr/86Sr(i) (0.7107–0.7116), lower 143Nd/144Nd(i) (0.512218– 0.512220) and 1Nd(t) (27.5 to 28.3), indicating a dominant crustal component for the origin of its magmas (Fig. 14). These geochemical features and isotopic compositions of the Salihli granitoid are similar to

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Fig. 13. La/Yb versus La (ppm) diagram illustrating the effects of partial melting and fractionation. Vectors for FC and PM are from Thirlwall et al. (1994). See Figure 6 for the data sources.

those of the granitoids in the Cyclades and in the Rhodope massif (Figs 11c, 12b & 15). Pe-Piper & Piper (2001) and Pe-Piper et al. (2002) have suggested that mafic magma fractionation and/or mixing with felsic crustal material, some of which was derived by crustal anatexis (Alther & Siebel 2002), could produce the granitoid plutons of the Cyclades. On the other hand, felsic magmas of the Rhodope granitoids were most likely derived from some crustal melts that originated from dehydration melting at mid- to deep-crustal levels (Christofides et al. 1998). The most reasonable model for the origin of the MEG magmas involves, therefore, partial melting of

Fig. 14. Nb/La versus Ba/Rb diagram illustrating the effects of crustal contamination and subduction metasomatism during evolution of the MEG. The values for CC (Average Continental Crust) are from McLennan (2001) and for N-MORB and OIB are from Sun & McDonough (1989). See Figure 6 for the data sources.

Fig. 15. Epsilon-Nd(i) versus 87Sr/86Sr(i) diagram showing isotopic compositions of the MEG and Cyclades granitoids. Data for asthenospheric and lithospheric mantle melting array are from Davis & von Blanckenburg (1995), for Aegean metamorphic basement from Briqueu et al. (1986), for Aegean Sea sediments from Altherr et al. (1988) and for Global River Average from Goldstein & Jacobsen (1988). See Figure 6 for the data sources.

a mixed source including varying proportions of an enriched lithospheric mantle and assimilatedmelted lower and middle crustal components (Figs 14 & 15). We infer that crustal-scale shear zones facilitated the uprising of melts derived from a previously enriched lithospheric mantle and their transport to lower to mid-crustal levels, where further melting and magma mixing occurred within the crust.

Fractional crystallization processes The observed linear variations in the Harker diagrams (Fig. 9) and REEcn patterns (Fig. 11a) indicate that fractional crystallization was an important process during the evolution of the MEG magmas. The negative covariances between SiO2 and FeO* (FeO þ Fe2O3), MgO and CaO (Fig. 9) indicate fractionation of olivine and clinopyroxene during the evolution of the MEG magmas. Variations in CaO with silica are almost the same in all groups (Fig. 9). Decreasing CaO with increasing silica indicates that Ca-rich phases such as hornblende and Ca-plagioclase were progressively formed and then removed from the granitic melt. The lack of notable negative Eu and Sr anomalies suggests that plagioclase fractionation was insignificant during the evolution of the Salihli granitoid. However, the Eu/Eu* ratios (Eu/Eu*: 022– 088) and their relationship with the increasing silica content of the Turgutlu and Egrigo¨z granitoids indicate that plagioclase fractionation was important during formation of the more siliceous members of the MEG

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(Fig. 11). The increasing Rb and decreasing Sr contents of the MEG vary with increasing SiO2, pointing out significant feldspar fractionation during the evolution of the MEG. The depletion of Ti and P is also consistent with fractional crystallization of Fe– Ti oxides and apatite (Fig. 10a).

Geodynamics of late Cenozoic magmatism and its effects on extensional tectonics in the Aegean region The collision of the Sakarya and Anatolide –Tauride continental blocks in the late Palaeocene –early Eocene caused crustal thickening and orogen-wide burial metamorphism. This collision-driven regional metamorphism was responsible for the development of high-grade rocks in the Menderes metamorphic massif. Partial underplating of the buoyant Anatolide –Tauride block beneath the Sakarya continent jammed the north-dipping Tethyan subduction temporarily, while the continued sinking of lithospheric mantle resulted in slab breakoff (Fig. 16a). The emplacement of the widespread middle to late Eocene granitoid plutons (i.e. Orhaneli, Topuk, Gu¨rgenyayla, Kapidag) along the IASZ and into the Sakarya continent has been interpreted to have resulted from slab breakoff-related asthenospheric upwelling and associated partial melting of the subduction-metasomatized continental lithospheric mantle (Fig. 16a; Altunkaynak 2007; Dilek & Altunkaynak 2007). Resumed Tethyan subduction and associated slab rollback triggered upper plate extension, leading a tectonic collapse of the thermally weakened orogenic crust in western Anatolia during the late Oligocene –Miocene. This tectonic phase coincides with bimodal volcanism and widespread ignimbrite flare-up in the region (Pe-Piper & Piper 2006 and references therein). The Kazdag core complex in NW Anatolia began its initial exhumation in the latest Oligocene – early Miocene (Okay & Satir 2000) and the Menderes core complex in Central Western Anatolia underwent its initial exhumation in the earliest Miocene (Is¸ik et al. 2004; Thomson & Ring 2006; Bozkurt 2007; Dilek & Altunkaynak 2007). Some of the collision-generated thrust faults may have been reactivated during this time as crustal-scale low-angle detachment faults, (i.e. Simav detachment fault, SW Anatolian shear zone) facilitating the region-wide extension (Thomson & Ring 2006; C ¸ emen et al. 2006). Starting in the middle Miocene, the subcontinental lithospheric mantle beneath the Aegean region was delaminated as a result of peeling of its base due to rapid slab rollback at the Hellenic trench (Fig. 16b). Asthenospheric upwelling caused by this lithospheric delamination led to melting of the

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subduction-metasomatized lithospheric mantle that in turn provided heat to the overlying crust. Invasion of the lower and middle crust by lithospheric mantle-derived melts triggered MASH-type processes (melting, assimilation, storage, homogenization; Hildreth & Moorbath 1988), resulting in the production of hybrid magmas of the MEG (Fig. 17a). Thus, lithospheric mantle and crustal melts were involved in the evolution of hybrid magmatism in the middle Miocene. Thermal relaxation associated with this magmatic phase induced lithospheric-scale extension and accelerated lower crustal exhumation and doming across the Aegean region (Figs 16b & 17a; Lips et al. 2001). Sufficient cooling of the exhumed mid–lower crustal rocks (including the MEG plutons) in the Menderes core complex was followed around 7 + 1 Ma by the development of high-angle normal faults forming graben structures (Hetzel et al. 1995, 1998; Gessner et al. 2001; Lips et al. 2001). These faults crosscut the low-angle detachment surfaces and the earlier extensional deformational fabrics in the granitoid plutons and their metamorphic host rocks (Fig. 17b). This late-stage normal faulting caused relative uplifting of the graben shoulders and further exhumation of the detachment footwalls. The MEG plutons continued to be deformed cataclastically and brittlely during this phase as both their metamorphic host rocks and they were further uplifted tectonically. The development of the major graben systems (i.e. Bu¨yu¨k Menderes, Ku¨c¸u¨k Menderes, Alasehir, Simav; Fig. 17b) during the advanced stages of extensional tectonics in the late Miocene – Quaternary further attenuated the continental crust in the region. This extensional phase, accompanied by increased geothermal gradients, resulted in asthenospheric upwelling and corresponding decompression melting of the sub-asthenospheric mantle. This melting episode generated magmas with high concentrations of LILE and HFSE, radiogenic Nd, unradiogenic Sr and MORB and OIB-like Pb signatures (Alici et al. 2002; Dilek & Altunkaynak unpublished data) that produced the Kula volcanics erupted on the Menderes massif (Fig. 16c). Grabenbounding faults and lithospheric-scale shear zones facilitated the upward transport of these alkaline magmas to the surface with little or no crustal contamination. Hydrous melting of the mantle wedge peridotites above the southward-retreating Hellenic subduction zone produced the South Aegean arc volcanism farther south since the late Miocene (Fig. 16c; Pe-Piper & Piper 2006). The close temporal and spatial relationships between the late Cenozoic tectonic extension and magmatism in the broader Aegean province indicate that lithospheric-scale melting played a significant role in weakening the young orogenic crust,

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Fig. 16. Cenozoic geodynamic evolution of the western Anatolian region through collisional and extensional processes in the upper plate of north-dipping subduction zone(s) within the Tethyan realm. (a) Collision and partial subduction of the Anatolide–Tauride continent (ATC) beneath the Sakarya continent leads to slab break-off of the Tethyan lithosphere. Asthenospheric upwelling through the breakoff-induced window facilitates partial melting of the subduction-metasomatized mantle beneath the suture zone and the Sakarya continent that in turn generates the Eocene to Oligo-Miocene volcanoplutonic complexes in NW Anatolia. (b) Rapid slab reatreat of the northward subducting Southern Tethyan oceanic lithosphere triggers lithospheric delamination in the middle Miocene. Associated asthenospheric upwelling results in asthenospheric and lithospheric mantle melting. Produced melt reacts with the lower and middle crust where it undergoes fractional crystallization, assimilation, storage and further melting, generating hybrid magmas of the MEG granitoids. This phase of magmatism was synchronous with and facilitated lithospheric-scale extension in the region. Similar melting processes, magmatism and extension are inferred for the Kazdag and Rhodope massifs to the north. (c) Advanced stages of tectonic extension and crustal attenuation induce asthenospheric upwelling, leading to decompression melting of the asthenospheric or sub-asthenospheric mantle. This melting event in turn generates MORB- and OIB-like magmas of the Kula volcanic domain. Menderes, Kazdag and Rhodope massifs are substantially cooled off (due to their unroofing) by this time period and have undergone high-angle normal faulting. The South Aegean arc volcanism is fed by hydrous melting of the mantle wedge above the Hellenic subduction zone. Key to lettering: ATC, Anatolide–Tauride continent; IASZ, Izmir–Ankara suture zone; IPSZO, Intra-Pontide suture zone ophiolites; LO, Lycian ophiolites; MBL, Mechanical boundary layer; NAFZ, North Anatolian fault zone; RM, Rhodope massif; RPC, Rhodope-Pontide continent; SC, Sakarya continent; TBL, Thermal boundary layer.

facilitating vertical crustal flow and exhumation of core complexes (Fig. 17a). Similar close relations between magmatism and extension have been documented from farther south on the Aegean islands (Avigad & Garfunkel 1991; Vanderhaeghe 2004)

and from other extended terranes such as the Basin and Range in the North American Cordillera (Gans et al. 1989) and the D’Entrecasteaux Islands of Papua-New Guinea in the SW Pacific (Baldwin et al. 1993).

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Fig. 17. Schematic sequential cross-sections, depicting the inferred tectonomagmatic evolution of the metamorphic core complexes in western Anatolia during the middle Miocene (a) and late Miocene (b). See text for discussion. Key to lettering: AG, Alasehir graben; BMG, Bu¨yu¨k Menderes graben; CGD, C ¸ ataldag granitoid; DF, Datc¸a fault; IASZ, Izmir–Ankara suture zone; IGD, Ilica granitoid; KDCC, Kazdag core complex; KMG, Ku¨c¸u¨k Menderes graben; LON, Lycian ophiolite nappes; MEG, Menderes granitoids; SG, Simav graben; TBS, Tavsanli blueschists; TO, Tethyan ophiolites.

Conclusions The Miocene granitoid plutons in the metamorphic core complexes in western Anatolia are synextensional intrusions, indicating close spatial and temporal relations between magmatism ad extensional tectonics during the late Cenozoic geodynamic evolution of this region. This interpretation is supported by: (1) the nearly coeval crystallization and cooling ages of the granitoid plutons and the deformation ages of their host metamorphic rocks and (2) the spatial progression from undeformed granitoid rocks at depth toward highly deformed, mylonitic–ultramylonitic and cataclastic plutonic rocks structurally upward into the shear zones associated with the detachment surfaces in the core complexes. This granitic magmatism was instrumental in crustal weakening that led to the collapse and tectonic thinning of the early Cenozoic orogenic belt in the Aegean region, in tandem with

rapid rollback of the north-dipping Hellenic subduction zone. The compositional variations of the synextensional MEG are likely to have resulted from different degrees of partial melting of a mixed source including varying proportions of an enriched lithospheric mantle component and assimilatedmelted lower and middle crustal components. Crustal-scale extensional shear zones gave rise to uprising of melts derived from previously enriched lithospheric mantle to lower- and mid-crustal levels, where further melting and mixing occurred within the crust. At the present level of exposures, there is no evidence of basic intrusive rocks within the Menderes massif. Nevertheless, mantle-derived magmas likely contributed to the granite petrogenesis by invading the crust. Partial melting of the subduction-metasomatized lithospheric mantle and the overlying lower crust that led to the formation of the MEG magmas was

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induced by asthenospheric upwelling caused by lithospheric delamination. This inferred lithospheric delamination was triggered by peeling of the base of the subcontinental lithosphere as a result of the rapid slab retreat of the newly established post-Eocene Hellenic subduction zone. The combination of the slab rollback-generated upper plate extension and lithospheric-scale partial melting, aided by asthenospheric upwelling, migrated southward in time, resulting in the younging of core complex formation and associated magmatism toward the Hellenic trench. This study was supported by research grants from the Scientific & Technical Research Council of Turkey (TUBITAK-CAYDAG-101Y006), Istanbul Technical University (Bilimsel Arastirma ve Gelistirme Destekleme Programi – BAP) and Miami University Committee on Faculty Research. The fieldwork of Z. O. in western Turkey was supported by the Geological Society of America and the American Association of Petroleum Geology Grant-in-Aid research funds. We thank Y. K. Kadioglu for his help with the geochemical analyses of the Salihli granotoid samples in his laboratory in the University of Ankara (Turkey). Constructive and thorough reviews by Uwe Ring, Stuart N. Thomson and an anonymous referee helped us improve the paper greatly.

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A case study of lateral spreading: the Precambrian Svecofennian Orogen ANNAKAISA KORJA*, PAULA KOSUNEN & PEKKA HEIKKINEN Institute of Seismology, POB 68, FI-00014 University of Helsinki, Finland *Corresponding author (e-mail: [email protected]) Abstract: We have studied the crustal structures of the Palaeoproterozoic Svecofennian (c. 1.9 Ga) Orogeny with the help of large scale seismic reflection surveys (FIRE 1– 3), preliminary structural field work and geological and geophysical databases. The central part of the orogen is occupied by the Central Finland Granitoid Complex, which comprises two suites of granitoid rocks and associated mafic and volcanic rocks. The complex and the surrounding supracrustal belts are cut and deformed by numerous shear zones and faults; here divided into six groups. The most prominent reflections are usually shear zones or faults on outcrops. The granitoid complex is interpreted as a deep, lower-level section of an old core complex, where the younger granitoid intrusions form the basins and older granitoid intrusions and associated volcanic rocks form the horsts. The upper–middle crust detachment zone is exposed at the northeastern edge of the complex and middle crust is exposed in the migmatitic domes at northern and western margins. The seismic reflection sections display a frozen image of orogenic thickening and lateral spreading. The decoupling of the upper, middle and lower crust during spreading resulted in the formation of layered superstructure–infrastructure of the crust.

One third of the Earth’s surface is covered with continental lithospheric lithosphere and the rest with oceanic lithosphere. The growth of these rigid, buoyant continents takes place at convergent plate boundaries, where material is added via subduction and/or via accretion of smaller and bigger fragments of island arcs, sedimentary basins or continents. The process of merging the independently arriving lithospheric entities with the rim of a continental plate is complex and has three phases: precollisional, collisional and postcollisional, each of which will leave different tectonic markers. The first phase brings in material and the second one accretes the material to the continent, and the last phase determines whether the terrane is destroyed or preserved in the geological record. The third, post-collisional phase also determines the final geometrical relationships of the geological and geophysical records. Although seismic datasets from different parts of the world (Cook et al. 1979, 1999; BABEL Working Group 1993; Beaumont & Quinlan 1994; Brown et al. 1996; van der Velden & Cook 1999; Beaumont et al. 2000; TRANSALP Working Group 2002; Culshaw et al. 2006; Korja & Heikkinen 2008) have shown that the crustal structure is dependent on the formation environment irrespective of its age, most of the structures observed in inactive areas are interpreted in terms of the precollisional and collisional phases (Meissner 1996; Cook et al. 1999) and the last, post-collisional phase (Dewey

1988; Platt 1993) that largely controls the present distribution of geological formations and terranes is largely neglected (Culshaw et al. 2006). Culshaw et al. (2006) challenged the reflection seismologists to reinterpret their data sets with respect to superstructure–infrastructure produced in the post-collisional phase. In conventional collision models (Coward 1994), the lithosphere thickens via breaking and stacking of either the footwall (lower plate) or the hanging wall (upper plate) as well as via stacking of the accreting allochthonous material. In the crustal blocks, the deformation is mainly restricted to discrete deformation zones, the stacking surfaces. This deformation mode is called semiplastic by Meissner (1996). In crustal flow models (Beaumont et al. 2001, 2004; Jamieson et al. 2004), the thickening is more homogenous and takes place by flow of warm viscous material and the deformation should leave a pervasive flow structure. In both cases, the orogenic front moves towards the thinner areas and i.e. the orogen grows forward it may also grow sideways by lateral escape. During thickening, part of the kinetic energy of plate movement is transformed into potential and thermal energy of the thickened crust, whose potential energy level is raised relative to its surroundings. Lateral spreading of the orogen to the sides tends to equilibrate the potential energy differences instantly and thus the orogenic crust is thinned and adjacent crust is thickened (Rey et al. 2001).

From: RING , U. & WERNICKE , B. (eds) Extending a Continent: Architecture, Rheology and Heat Budget. Geological Society, London, Special Publications, 321, 225–251. DOI: 10.1144/SP321.11 0305-8719/09/$15.00 # The Geological Society of London 2009.

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In young orogens, the process is called orogenic collapse, extensional collapse or gravitational collapse (Dewey 1988; Rey et al. 2001). The equivalent stabilization process of the older Precambrian orogens is traditionally called cratonization, which is not very well quantified. The thermal energy surplus creates new pressure–temperature fields and induces regional metamorphism. Associated partial melting drives lithological differentiation and density stratification of the crust, which also help to isostatically balance the crust in vertical direction. Beaumont et al. (2001, 2004) have suggested that thermal weakening results in the formation of orogenic superstructure–infrastructure (Culshaw et al. 2006) of large hot orogens and is partially responsible for at least enhancing the layering of the crust. In addition to plate tectonic forces and variations in temperature and pressure, the mode of deformation in collisional environments is largely dependent on initial distribution of lithologies and especially their rheological properties. The crust may respond differently to stress at different depths of the crust and it may decouple to two or more differently behaving layers. Rheology, together with geothermal gradient, determines whether detachment surfaces form at upper, middle or lower crustal levels (Kusznir & Park 1987). The uppermost crystalline crust always deforms in brittle mode along discrete shear planes. The deformation mode changes gradually from discrete to more pervasive plastic deformation as the depth changes from upper to lower level. Concurrently, the fold structures change from upright to flat lying. This gradual change is deformation style may be one of the reasons why reflection sections are dominated by low angle structures at depth, whereas the exposed surface is dominated by upright folds. Lower lever of deformation is reached already in amphibolite facies, and thus areas with regional amphibolite/granulite metamorphism, like the Precambrian cratons, are expected to be characterized by pervasive low-angle fabrics. These areas could be used to test how the seismic patterns change across metamorphic boundaries. Another problem may, however, rise from the fact that the Precambrian terranes usually have flat topography and low-angle deformation patterns are therefore not so easily observed. The strength of the thickened crust/lithosphere is dependent on the composition and geothermal gradient of the arriving blocks and the receiving continent and on the total thickness of the collage (Kusznir & Park 1987). The overall distribution of tectonic forces at the convergent margin determines the deformation rates and the lateral extension rate at which the crust is thinned. The extension rate seems to be the controlling factor on whether the

extensional environment is frozen as a consequence of strain hardening or is destroyed by strain weakening leading to crustal separation (Kusznir & Park 1987). Slow to intermediate strain rates (10216 –10215 s21) in regions with moderate to high geothermal gradients should produce wide regions of strain hardened, stable crust (Kusznir & Park 1987). The extensional collapse process produces crustal and lithospheric-scale structures that can be classified as narrow rifts, wide rifts or core complexes (Buck 1991). Although core-complex-mode extension has been suggested to take place in areas of gravitational collapse (Brun 1999; Corti et al. 2003), narrow rift or wide rift modes may also be applicable in orogens where several tectonic processes: sidewise spreading of the orogen, gravitational collapse, retreating subduction and slabrollback may contribute to the extension. Recently, a high resolution deep seismic survey (FIRE) across the Palaeoproterozoic Svecofennian orogen (Kukkonen et al. 2006) revealed a layered crust with deformational features characteristic of collision followed by extension (Korja et al. 2006; Nironen et al. 2006; Sorjonen-Ward 2006; Korja & Heikkinen 2008). The focus of this paper is to describe crustal structures produced by the postcollisional phase of the Svecofennian Orogeny, and to discuss the reasons why the Svecofennian crust was laterally extending and why it became the stable nucleus of the Fennoscandian Shield. We will analyse the frozen seismic images and combine the data with surface geological observations and with older crustal and lithospheric data sets in order to compare the observed crustal structure to the existing extensional models and to seek answers to such questions as what was the mode of extension, what drove the extension and why the extension ceased.

The Svecofennian orogen The Palaeoproterozoic Svecofennian orogen (Fig. 1) shows evidence of rapid crustal evolution from island arc environments to stable continental crust during 2.1– 1.87 Ga (Park 1991; Gorbatschev & Bogdanova 1993; Korsman et al. 1997, 1999; Lahtinen et al. 2005 and references therein) in an environment reminiscent of the modern Indonesian archipelago (Ward 1987). In Indonesia, new continental crust is being produced in a complex tectonic setting, where oceanic plates and an oceanic part of a continental plate (Australia) are consumed in oblique subduction zones above which arcs live only briefly and new microplates are forming (Hall 2002). Sooner or later, parts of the microplates will accrete to the larger continental plates.

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Fig. 1. (a) FIRE 1–3 reflection lines on a schematic tectonic map of the central part of the Svecofennian orogen. The black lines indicate structural discontinuities and shear zones, the stippled grey lines denote a crustal-scale electrical conductors after Korja et al. (2002). The red and blue lines within the CFGC denote sets of Group 1 and 2 orthogonal faults (see Fig. 3), the blue arrows indicate the spreading direction of the orogen. The boxes denote areas shown in Figures 2 (outer) and 3 (inner). The major tectonic units are: BB, Bothnian Belt; CFGC, Central Finland Granitoid Complex; KA, Archaean Karelian domain; OB, Outokumpu Belt; SAC, Southern Finland Arc Complex; SB, Savo Belt; TB, Tampere Belt. The Ladoga– Bothnian Bay (LBB) wrench fault zone is indicated with arrows. (b) A schematic plate-tectonic reconstruction before a double-sided collision, modified after Lahtinen et al. (2005). The plate movement vectors are shear wave anisotropy directions from the lithospheric mantle (Plomerova et al. 2006), interpreted as paleo-stress vectors. The velocities of the plates are tentative. The geometry of the triple junction resembles that of the electrical conductors in (a).

In its central part, the Svecofennian orogen comprises rocks formed within the Karelian continent, Savo Arc and Western Finland Arc Complex and is bordered in the south by the Southern Finland Arc Complex (Korsman et al. 1999). The major terrane boundaries are marked by crustal-scale electrical conductors (Korja et al. 1993; T. Korja et al. 2002; Fig. 1a) and crustal-scale reflective zones (Korja & Heikkinen 2008). A recent evolutionary model (Lahtinen et al. 2005; Korja et al. 2006) suggests that the central part of the orogen is underlain by a Palaeoproterozoic microplate (Keitele) under which oceanic crust was subducting from the east and south producing the Savo and Tampere arcs, respectively (Fig. 1). After the surrounding ocean was consumed, the microplate collided with the continental parts of the arriving plates (Karelia and Bergslagen) and the orogeny was initiated. The crust overthickened and collapsed and in these processes the crust attained its characteristic large thickness (.56 km), thick high velocity lower crust (vp . 6.8 km s21 between

depths 30 and 60 km; Korja et al. 1993) and thick high velocity lithosphere (.200 km; Sandoval et al. 2004). Figure 1b shows one solution for the plate tectonic setting just prior to the doublesided collision. The Karelian continent (KA; Figs 1a & 2) comprises Archaean granitoid–gneiss complexes and greenstone belts (3.2–2.5 Ga) and is overlain locally by autochthonous Palaeoproterozoic supracrustal cover rocks (Sorjonen-Ward & Luukkonen 2005). During the orogeny, the western parts suffered from metamorphic overprinting with resetting of K –Ar isotopes and localized shearing. Mantled gneiss domes and metaturbiditic sequences hosting ophiolites (1.97 Ga) were thrust on the continental margin from the west (Park et al. 1984; Kontinen 1987). Both the amount of intruding granites and the metamorphic degree increase towards the suture (Korsman & Glebovitsky 1999). The suture with the Savo Arc is overprinted by a large-scale, NW –SE-trending strike slip fault zone, the Ladoga –Bothnian Bay Zone (LBB; Figs 1a & 2).

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Fig. 2. FIRE profiles 1, 2 and 3a on a geological map of the Central Finland Granitoid Complex, modified after Korsman et al. (1997) and Nironen (2003). Legend: BB, Bothnian Belt; CFGC, Central Finland Granitoid Complex; KA, Archaean Karelian domain; LBB, Ladoga– Bothnian Bay wrench fault zone; OB, Outokumpu Belt; SAC, Southern Finland Arc Complex, SB, Savo Belt; TB, Tampere Belt. The box denotes the area shown in Figure 3. Magenta lines denote the location of detailed crustal sections of Figure 7.

The NW–SE-trending Savo Arc (SA; Figs 1a & 2) consists of metavolcanic and -turbiditic rocks that are interlayered with gneissic tonalites (1.92 Ga) with positive 1Nd(T ) indicating juvenile sources (Lahtinen & Huhma 1997). The heavily

deformed and steeply dipping volcano-sedimentary complex were intruded by calc-alkaline plutons and extruded by their associated volcanic rocks (1.89– 1.88 Ga) during the peak of deformation (Korsman et al. 1999). The exposed rocks of the

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Savo Arc have been metamorphosed under HT –LP conditions and have attained upper amphibolitesfacies conditions with a few blocks in granulite facies. The intrusion of only slightly deformed, locally pyroxene-bearing granites (1.885 Ga) ended the magmatic evolution of the arc complex. The suture between the Savo arc and the Western Finland Arc Complex is hidden beneath the Central Finland Granitoid Complex. The Western Finland Arc Complex (WAC; Figs 1a & 2) comprises supracrustal belts (Bothnian Belt, BB; Tampere Belt, TB) surrounding the Central Finland Granitoid Complex (CFGC). The Bothnian Belt (1.90–1.87 Ga) in the west comprises a large migmatite –granitoid complex (Ma¨kitie 2000) as well as metaturbiditic sequences hosting mafic volcanic units (Korsman et al. 1997). The Tampere Belt (1.90–1.89 Ga) comprises well-preserved greywacke and volcanic rocks formed in a mature oceanic arc or close to a continental margin. Migmatitic greywackes south of Tampere are interpreted to be remnants of an accretionary prism (Ka¨hko¨nen 1987; Korsman et al. 1999). Both Bothnian and Tampere supracrustal belts have suffered from regional metamorphism and poly-phase deformation and have been intruded by calc-alkaline and minor alkaline granodiorite and granite intrusions with ages of 1.89– 1.87 Ga (Patchett & Kouvo 1986; Nironen 1989; Kilpela¨inen 1998; Ma¨kitie 2000; Korsman et al. 1999). The Central Finland Granitoid Complex (CFGC; Figs 1a & 2) consists of two suites of granitoid intrusions and associated mafic and volcanic rocks. The available isotopic and geochemical data indicate the existence of evolved crust and associated lithospheric mantle beneath the Central Finland Granitoid Complex already prior to 1.91 Ga (Lahtinen & Huhma 1997). This unexposed continental block has tentatively been called Keitele in recent plate tectonic models (Lahtinen et al. 2005; Korja et al. 2006). The marginally older suite of granitoid rocks is composed of calcalkaline granodiorites and granites (1.90–1.88 Ga) with minor amounts of mafic plutonic rocks and co-magmatic volcanic sequences (Harris et al. 1986; Korsman et al. 1997; Nironen et al. 2000). The suite has traditionally been called syncollisional (Korsman et al. 1997), but it post-dates collision and contains remnants of folded supracrustal fragments. The intrusion of the igneous rocks was simultaneous with the peak of the HT –LP metamorphism (T ¼ 700 –800 8C, P ¼ 400 –500 MPa) and deformation (Korsman et al. 1999). The calc-alkaline granitoid rocks were later deformed pervasively, as recorded by consistent lineations oriented towards E–SE (Kilpela¨inen et al. 2008). Soon after or contemporaneously with the calc-alkaline

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suite, a series of alkaline (A-C-type), occasionally pyroxene-bearing granites with minor granodiorite and quartz monzonite plutons intruded the complex (at 1.88–1.86 Ga). These, as well as a series of coeval gabbros and mafic dykes, intruded along NW–SE-trending shear zones in transtensional environments (Elliott et al. 1998; Nironen et al. 2000). In the alkaline suite, deformation is more restricted and localized into NE– SW and NW –SE oriented shear zones.

Aeromagnetic data Regional-scale low-altitude aeromagnetic anomaly map provides an excellent tool for studying various regional structural features, especially shear and fault zones. In our study area (Fig. 3b), the more magnetic supracrustal and mafic rocks form local anomaly maxima in the background composed of two suites of granitoid intrusions. The post-collisional alkaline granites are displayed as homogeneous, featureless magnetic minima and in outcrop they show discrete and localized shearing and faulting. The older calc-alkaline granitoid rocks are inhomogeneous and they are pervasively deformed with a shallow, 10–308 ESE-plunging lineation (Kilpela¨inen et al. 2008). Shear zones cutting the bedrock are displayed as linear magnetic minima or maxima disrupting the surrounding deformation patterns. Based on lineament interpretation, previous published structural maps (Kuosmanen 1988; Nironen 2003; Korsman et al. 1997; Korsman & Glebovitsky 1999) and reconnaissance field studies (Fig. 3a) we have divided the predominant lineaments into six groups (Fig. 3b), three of which (Groups 1 –3) coincide with the three groups previously described by Nironen (2003). The faults of Groups 1 –4 form two orthogonal pairs at 458 angle; Groups 5 and 6 are Riedel shears. Group 1 (green, Fig. 3b) consists of NE–SWstriking lineaments, which are especially abundant in the western part of the area and are typically observed as smooth changes in magnetic anomaly levels. For example at Soini, the local magnetic maximum is introduced by a moderately SEdipping (strike 508, dip 130/408) felsic metavolcanic rock exposed next to sheared metagabbro (Fig. 4a, Photo 1). Locally, Group 1 faults give rise to metamorphic block structure and thus Nironen (2003) has interpreted those to have a vertical component. Group 2 faults (Fig. 3b, blue), that form an orthogonal pair with faults of Group 1, are observed throughout the area as sharp NW –SE-trending minima. The Group 2 faults seem to deform and displace the faults of Group 1 and field observations

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Fig. 3. Structural observations, shear zones and some alkaline granites and related gabbro and diorite intrusions on an aeromagnetic anomaly map of the study area in Central Finland (the map courtesy of Geological Survey of Finland). Locations of detailed upper crustal seismic sections of FIRE 1 & 3a (Figs 8– 11) are indicated by coloured lines and locations of outcrop photographs in Figure 4 are indicated with numbered dots. (a) Revised structural observations and examples of different shear and fault zone groups are high-lighted in transparent colours. (b) Magnetic anomalies interpreted as shear and fault zones are divided into Groups 1 –6 and marked with transparent coloured lines.

indicate them to contain a dextral horizontal displacement component (Nironen 2003). Group 2 faults are often at the fringes of post-collisional alkaline granite intrusions and seem to have partially controlled their emplacement, under extensional or transtensional conditions (Nironen 2003). A transextensional environment is also implied by local development of Riedel faults (Group 5; yellow) between Group 2 faults and dragging of older structures to the fault plane (Fig. 3b).

An example of a subvertical shear zone of Group 2 is found in a quarry in alkaline granite close to Karstula (Fig. 4b, Photo 2). The massive, coarsegrained porphyritic granite is locally deformed by a nearly vertical, NW–SE-directed shear zone. The lithology changes from undeformed porphyritic granite (Fig. 4b, Photo 2A) through sheared augen gneiss (2B) to ultaramylonite (2C, D) within a distance of 200 m. The shear indicators (quartz veins, rotated crystals, splitting of mylonite, lineations)

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Fig. 3. (Continued).

suggest oblique, right-lateral movement and relative downward motion of the NE side. The preserved parts of Group 2 shear zones are typically found on the NE flanks of topographic depressions. Another, magnetically much less distinctive, orthogonal pair is found at 458 angle to the first pair. The second pair is formed by the lineaments of Groups 3 and 4. Group 3 lineaments (red, Fig. 3b) run north–south and are most pronounced in the eastern part of the study area. According to Nironen (2003), field evidence points to both

vertical and dextral horizontal displacement components in these faults; development of Riedel faults (Group 6; orange) between Group 3 faults in the east (Fig. 3b) supports the idea of rightlateral movement. A prominent example of Group 3 shear zone is located at Pyha¨ja¨rvi. It is over 30 km long and 0.5 km wide NNE-trending shear zone (Fig. 4c, Photo 3) across which the lihtologies change from granitoid rocks to flow structured diatexites and metatexites (Fig. 4d, Photo 3A) and the

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Fig. 4. Outcrop photographs on exposed shear zones. Photograph locations are shown in Figure 3. (a) Photo 1. Sheared metagabbro in contact with felsic volcanic rocks at Soini is an example of the Group 1 listric shear zones. The contact and the schistosity is ESE dipping (130/408). (A) A vertical surface of the metagabbro, and (B) a horizontal surface of the felsic metavolcanic rock. The magnetic anomaly is rising from the metavolcanic rocks carrying high

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metamorphic grade changes from andalusite to granulite facies (Kuosmanen 1988; Ho¨ltta¨ 1988; Marttila 1992). The deformation fabric of the host rock granitoid and metavolcanic rocks to varies between strongly stretched to mylonitic. Group 3 deformation is illustrated in Figure 4c on a sheared mafic metavolcanic rock (uralite porphyrite) that in mylonitic bands has changed into epidote- and calcite-bearing talc –chlorite schist. The magnetic signature of the east –west lineaments of Group 4 (Fig. 3b; magenta) is relatively weak. Field observations from the Group 4 structures indicate low angle slip planes (Fig. 4e– g; Kilpela¨inen 2007, Kilpela¨inen et al. 2008). The most prominent known Group 4 structure, the Ela¨ma¨ja¨rvi shear zone, is located in the northern part of the study area. It is a 10 km wide and 70 km long macro-structure, where highly strained rocks are found in low-angle orientations. The shear zone is best observed where it deforms the 1870 Ma old, alkaline Pyytsalo granite (Photos 4 & 4A in Figs 3b & 4e; Kousa et al. 1994), but it can be followed into the surrounding lithologies as well. Within the Pyytsalo granite, the lithologies range from sheared coarse porphyritic granite to augen gneiss, mylonite and to ultramylonites cutting mylonites suggesting that the shear zone was active for a long time and that the style of deformation changed from lower level plastic (sheared granite) to upper level brittle (ultramylonite breccia) deformation. The Ela¨ma¨ja¨rvi shear zone displays oblique shearing with shallow plunging lineations (10–208) towards ESE and kinematic indicators suggesting westward movement of the hanging wall along low-angle oblique slip planes (Photo 4A in Figs 3b & 4f; Kilpela¨inen 2007; Kilpela¨inen et al. 2008). Similar low-angle oblique

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shearing is observed in a young shallow-dipping aplite dyke/sill at Kivija¨rvi (Photo 4B, Fig. 4g).

Seismic data The seismic data used in this study is from deep seismic reflection profiles FIRE 1–3, which transect the major structural units of the Svecofennian orogen in NNE–SSW and ESE –WNW directions (Figs 1 & 2). The profiles cross within the study area and provide an excellent opportunity for three-dimensional study of the crustal structure. Preliminary interpretations of the data sets have been presented by the Fire Working Group (Kukkonen et al. 2006; Nironen et al. 2006; Korja et al. 2006; Sorjonen-Ward 2006; Korja & Heikkinen 2008). The seismic sections shown here (Figs 5– 11) are either crustal-scale migrated NMO-stacks or migrated near-surface sections of DMO-stacks of the FIRE profiles 1– 3. Seismic reflections rise from lithological contacts with large velocity and density contrasts. In highly metamorphosed and deformed areas, lithological contacts can be either primary contacts like dykes or igneous batholiths intruding deformed supracrustal rocks or tectonic contacts formed during folding and faulting. Weak reflectivity indicates a rock mass with little or no density and velocity contrasts, e.g. monotonous intrusions and older crustal pieces in which the internal structure has prior to deformation been homogenized in the scale of reflectivity. On seismic sections compression produces concave (thrusts) and extension convex (listric) reflections in general (Meissner 1996). Detachment zones and decollements are continuous subhorizontal reflections toward which other reflection

Fig. 4. (Continued) susceptibility values. (b) Photo 2. An example of a several metres wide, subvertical shear zone of Group 2 found in an alkaline granite quarry close to Karstula. (A) The typical texture of the massive, porphyritic Karstula post-collisional granite. (B) Closing in on the shear zone (2208/858), the granite becomes epidotized and deformed by shear-bands, S-C structure indicates dextral movement at horizontal surface. (C) Vertical surface of intensely sheared Karstula granite at the nearly vertical shear zone (2208/858); the texture varies from augen gneiss to ultra-mylonite. (D) A moderately plunging lineation (1358/508) developed on the mylonite surfaces indicating oblique shearing on subvertical surface. (c) Photo 3 & 3A. An example of Group 5 shear zones at Pyha¨ja¨rvi. Within the shear zone mafic metavolcanic rock (uralite porphyrite) host rock has changed into chlorite– biotite–epidote- and calcite-bearing schist. (d) Photo 3A. A horizontal surface of flow structured metatexite, where amphibolite fragments are rotated towards NNW indicating the direction of magma flow. (e) Photo 4. The low angle of the Ela¨ma¨ja¨rvi shear zone is reflected by outcrops of the post-collisional Pyytsalo granite: the outcrops form low east–west ridges with steep southern edges and more gently dipping north sides. (f) Photo 4A. The vertical surface of a southern edge of an outcrop of the Pyytsalo granite (view towards north) is an example of Group 4 structures. The originally coarse-porphyritic granite has been deformed into augen gneiss with schistosity (C) dipping relatively shallowly towards north (0028/358) and a shallow lineation plunging ENE (0628/108). The shear indicators (S-C) suggest left-handed oblique shear (roof towards W). (g) Photo 4B. An example of Group 4 low-angle shear zones deforming younger subhorizontal aplite dyke. Subhorizontal foliation (C-structure) in (070/18) direction and lineation in (1048/208) direction has develop in response to low angle oblique shearing. (h) Photo 5. A pair of SW-dipping (2258/308) normal faults (Group 2/5) transecting a fine-grained plagioclase-porphyry at a quarry in the Kalmari volcanic sequence NW of Saarija¨rvi; Riedel-fractures (2108/758) are found between the faults. (i) Photo 6. The vertical surface of a sheared granite/granodiorite at Kyyja¨rvi displays the strong, shallow plunging lineation (0628/288) typical of the Kyyja¨rvi deformation zone.

234 A. KORJA ET AL. Fig. 5. An interpretation of the crustal structure along deep seismic reflection profile FIRE 1 & 2 shown as a straight line. Distance is given in CMPs and kilometres. No vertical exaggeration. The seismic data are shown as an averaged instantaneous amplitude section. The grey-scale intensities vary between 0 (white) and 20dB (black). Lithology at surface is after Korsman et al. (1997). Crossing-point of FIRE 3a is marked with a vertical line. Note the difference in reflective properties of upper, middle and lower crust. South-pointing herringbone structures and symmetrical thinning in its central parts characterize the middle crust. Subvertical transparent zones dip symmetrically towards the centre of the profile.

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Fig. 6. An interpretation of the crustal structure along deep seismic reflection profile FIRE 3a shown as a straight line. Distance is given in CMPs and kilometres. No vertical exaggeration. The seismic data are shown as an averaged instantaneous amplitude section. The grey-scale intensities vary between 0 (white) and 20 dB (black). Lithology at surface is after Korsman et al. (1997). Crossing-point of FIRE 1 & 2 is marked with a vertical line. The section images the rearrangement of collisional stacking structure by orogenic collapse. Note the difference in reflective properties of upper, middle and lower crust. The middle crust is characterized by listric reflections soling out the middle to lower crustal boundary and by anticlinal ramps.

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Fig. 7. Detailed images of crustal structures on FIRE 1 & 3a seismic reflection profiles traversing the Central Finland Granitoid Complex. Averaged instantaneous amplitude sections combined with automatic line drawing. Locations of the lines are in Figure 2. (a) A detailed image on FIRE 3a (CMPs 8000 and 10000) shows listric reflections soling out at the middle-lower crustal boundary (U-M) at the depth of 38 km. (b) A detailed image on FIRE 1 (CMPs 15000 and 17000) highlights the difference in deformation patterns above and blow the U-M detachment zone at the depth of 8 km. The middle crust (8 –25 km) shows isolated small scale reflections that form crustal scale herringbone structure.

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Fig. 8. A three-dimensional block diagram of the crustal structure beneath the Central Finland Granitoid Complex at the junction of FIRE 1 & 2 and 3a. View to south. Distance is given in CMPs and kilometres. No vertical exaggeration.The upper-middle crustal detachment is subhorizontal, the middle –lower crust detachment dips shallowly south to SE. The core of a poorly reflective, anticlinal structure in FIRE 3a is the poorly reflective core of a fishbone structure in FIRE 1 & 2.

238 A. KORJA ET AL. Fig. 9. Crustal structure of the uppermost 8 km of the northeastern edge of the Central Finland Granitoid Complex along FIRE 1 & 2. Major crustal reflections, modified after Korja et al. (2006), have been drawn on a colour-coded seismic reflection section. Lithology at surface is after Korja et al. (2006). Location of the section and the photographs are given in Figure 3. Shallowly south-dipping high-amplitude reflections are displaced by low-angle, lower amplitude reflections dipping north and by transparent sub-vertical zones. The section is interpreted to image how the low-angle normal faults (Ela¨ma¨ja¨rvi shear zone) and subvertical strike slip faults expose the upper– middle crust detachment surface and the underlying high temperature migmatites in the northern part (see Fig. 4).

LATERAL SPREADING OF THE SVECOFENNIAN Fig. 10. An interpretation of the uppermost 10 km of FIRE 3a between CMP points 10930 and 13430. Location of the section is shown in Figure 3. (a) Upper panel: grey-scale variable intensity DMO section. (b) Lower panel: an interpreted grey-scale variable intensity DMO section. Lithology at surface is after Korsman et al. (1997). Numbers denote locations of photographs shown in Figure 11. The depth extent of the granitoid plutons is estimated based on continuity of reflectivity pattern. The section displays an extensional environment in brittle regime. Note the sequential development of listric reflections and extensional duplexes.

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240 A. KORJA ET AL. Fig. 11. An interpretation of the uppermost 10 km of FIRE 1 & 2 between CMP points 15250 and 17750 after Korja et al. (2006). Location of the section is shown in Figure 3. (a) Upper panel: grey-scale variable intensity DMO section. (b) Lower panel: an interpreted grey-scale variable intensity DMO section. Lithology at surface is after Korja et al. (2006). Numbers denote locations of photographs shown in Figure 4. The depth extent of the granitoid plutons is estimated based on continuity of reflectivity pattern. The section displays the geometrical relationships between the apparently flat-lying high-amplitude reflections, the subvertical transparent zones and the less prominent low-angle reflections. The low-angle reflections could also be oblique cuts of more upright reflections.

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truncate. Near vertical strike-slip zones are observed indirectly as they are associated with transparent zones with decreased reflectivity and displacement of continuous reflections (Harding 1985). The decreased reflectivity is probably related to sets of closely spaced fracture zones destroying the continuity of the seismic reflectors and elastic properties. A well-known example is the San Andreas Fault, from which reflections have disappeared (Zhu 2000). Seismic reflection method is biased as it images better gently dipping surfaces than steeply dipping ones. A consequence of this that the tops and crests of synforms and antiforms are imaged as piece-wise continuous subhorizontal reflectors whereas sides are not so easily distinguished (Ji & Long 2006). This problem can be partly avoided by displaying the data as grey shaded instantaneous amplitude sections (Figs 5–7), which enhance differences in reflectivity between adjacent blocks, helping to define steeper block boundaries. All the sections are plotted without normalization i.e. the amplitudes of the different areas in each section are comparable. The instantaneous amplitude sections are averaged both horizontally and vertically and plotted as grey-scale intensities; the difference between black and white is 20 dB. Automatic line drawings are used to highlight the dip of the reflections in the more detailed sections (Fig. 7; for technical details, see Korja & Heikkinen 2005; Kukkonen et al. 2006). The steeply dipping structures close to the surface are better imaged in DMO corrected sections (DMO, dip move out). The coherency filtered migrated DMO sections imaging only the upper crust are displayed in variable area (Figs 8–10) mode and as colour-coded instantaneous amplitude section overlain by traces plotted in variable area mode. The seismic sections are two dimensional projections of three dimensional structures. In the sections, the observed, apparent dip angle (b) is always dependent on the interception angle (a) between the geological structure and the reflection line as well as the true dip angle (a) of the structure; sin˙ b ¼ sin a  sin a. Because the apparent dip of the reflection is always smaller than the true angle, the structures are steeper than they appear on the sections, unless the line runs perpendicular to the structure (a ¼ 908). If a line runs parallel to the strike (a ¼ 08), the structures in the section appear horizontal. True dip-directions of the reflections can been determined from the turning points or cross-points of two lines. Examples illustrating the three dimensional nature of the observations are illustrated with block diagrams (Figs 6 & 10) from the junction of profiles 1 and 3.

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In the following, we will analyse the reflection images and try to separate traces left by different tectonic processes. We will give examples of different types of fault zones, shear zones and detachment zones as well as flow structures, low-angle faults and transfer faults.

Crustal scale features The FIRE profiles display three crustal layers with different reflection properties: upper, middle and lower (Figs 5–8). The upper crust is associated with velocities of 5.8–6.2 km s21 and Vp/Vs ratios of 1.68–1.70, the middle crust with 6.3 –6.6 km s21 and 1.71 –1.74 and the lower crust with 6.8 –7.6 km s21 and 1.74–1.76, respectively (Korja et al. 1993; Korsman et al. 1999; Hyvo¨nen et al. 2006). The layers are separated by subcontinuos subhorizontal reflective boundaries interpreted as detachment surfaces (U-M, light violet, L-M, violet in Figs 5 & 6). A low velocity layer (LVL; Grad & Luosto 1987) is found in the uppermost part of the middle crust. The lower crust has a subdued and patchy subhorizontal reflectivity that dies out at the lower crust –upper mantle boundary: the Moho boundary (green in Figs 5 & 6). The subhorizontal reflectivity pattern suggests a ductile environment during the formation/deformation of the lower crust (Meissner 1996; Beaumont et al. 2004). FIRE 1 & 2 (Fig. 5) displays minor upwarping of the Moho boundary. The Moho boundary is marked by a small amplitude decrease in background reflectivity that coincides with the wide-angle reflection Moho observed on SVEKA81 (Grad & Luosto 1987) running parallel to the FIRE 1 & 2 line. We interpret that the Moho boundary has not been acting as a major deformation zone during the thinning of the crust and that it is more likely a lithological boundary. The upper crust is generally detached from the middle crust via a subhorizontal, highly reflective surface (U-M; Figs 5 & 6) that locally is also the upper boundary of the low-velocity layer. The middle crust has two components: large blocks of poorly reflective background material boarded by high-amplitude, shallow to steeply dipping crustal-scale reflections and smaller scale, highamplitude reflections. The crustal-scale reflections, mostly listric in shape, sole out at the boundary between the middle and lower crust (Figs 6 & 7a). These reflections have previously been interpreted to image crustal-scale stacking (Sorjonen-Ward 2006; Korja & Heikkinen 2008). The smaller scale reflections are arranged as herringbone structures on FIRE 1 & 2 (Figs 5 & 7b) and as westward ramping anticlinal structures on FIRE 3a (Fig. 6), both indicative of mid-crustal flow (infrastructure) according to Beaumont et al. (2001, 2004) and

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Culshaw et al. (2006). On a three-dimensional block diagram (Fig. 8), the mid-crustal structures seem to be dipping south indicating southward movement. On FIRE 1 & 2 (Fig. 5), the middle crust images symmetrical thinning, upwarping of the lower crust and subvertical crustal faults dipping symmetrically towards the centre, which leaves the crustal structure reminiscent of a pure shear rift (McKenzie 1978). On FIRE 3a (Fig. 6), the listric structures could also be interpreted to image sets of extensional faults in which case the structural pattern is reminiscent of an asymmetric simple shear rift with listric shear zones propagating towards the NW (Wernicke 1985; Lister et al. 1986).

The upper crust The upper crust (0– 12 km) displays fine layering with fine-scale deformation patterns (Figs 5 & 6). On profile FIRE 3a (Fig. 6), the uppermost crust displays listric reflections flattening at a subhorizontal surface, interpreted as a detachment surface, between the depths of 8 and 12 km. In the perpendicular direction, along FIRE 1 & 2 (Figs 5 & 7b), the same listric reflections are imaged as subhorizontal reflection surfaces, above which reflections mimicking half grabens and graben and horst structures are found and from where some near-vertical transparent zones spin off. The detachment surface is gently inclined southwards: it is at the depth of 12 km in the southern part of FIRE 1 & 2 (CMP 17200) and may reach the surface at Pihtipudas (CMPs 12350– 12800; Fig. 5). It may also be exposed in the western part of FIRE 3a (CMPs 13000–13400, Fig. 6). In the following, we will use the detailed upper crustal seismic sections and block diagrams (Figs 9– 12) to analyse how the reflection patterns are associated with the geological and geophysical observations at surface. The first example is from the NE boundary of the Central Finland Granitoid Complex (Fig. 9), where the upper– middle crust detachment zone (U –M) is exposed; the second example is from the crossing of the two reflection lines (Figs 10–12). In the NE corner of the Central Finland Granitoid Complex, FIRE 1 & 2 runs NNE– SSW direction (Fig. 3b) and it is perpendicular to Group 2 structures (incident angle c. 908), sub-parallel to Group 1 structures (incident angle c. 08), and obliquely to both Groups 3 and 4 (incident angles 508 –708). In the smaller scale seismic section (Fig. 9), the upper– middle crust detachment zone (U-M) is displayed as a series of highly reflective subhorizontal structures (blue). Those are cross-cut and displaced upwards by weaker, low-angle reflections (dark blue) and subvertical transparent zones (stipled black).

At depth, the thickness of the upper –middle crustal detachment zone (U-M) is c. 2– 3 km (Figs 5 & 9). If it were to exhume along low-angle shear zones dipping 10 –208 as implied by the seismic data (CMPs 12350 –12800; Fig. 9) its horizontal width would be c. 8–10 km, close to that observed for the Group 4 low angle shear zone at Ela¨ma¨ja¨rvi (Figs 3b & 4f). Thus, the Ela¨ma¨ja¨rvi macrostructure is wide (and long) enough to represent a major crustal deformation zone separating the upper and middle crust. In the Pihtipudas block, which is south of the Ela¨ma¨ja¨rvi zone and thus structurally above the U-M detachment, the supracrustal rocks are well-preserved and record upper amphibolite facies metamorphic conditions (Aho 1979; Ho¨ltta¨ 1988; Rastas & Kilpela¨inen 1991). All rocks regardless of age record shallow, 10–308 ESE plunging lineations (Kilpela¨inen et al. 2008). North of the Ela¨ma¨ja¨rvi shear zone (CMPs 12100 –12200; Fig. 9), a near vertical transparent zone is exposed as a Group 3 Pyha¨ja¨rvi shear zone (Fig. 4c). The reflection image suggests that reflective units below the detachment are exposed on the northern side of the shear zone. The northern block comprises flow structured diatexites and metatexites (Fig. 4d) grading northwards into migmatitic mica gneiss. Preliminary structural studies suggest relative upward movement of the northern block (Kilpela¨inen et al. 2008). Earlier metamorphic studies on this block indicate that the HT –LP (650 8C, 5 kbar) metamorphism post-dates the 1.89 Ga calc-alkaline magmatism in the area (Ho¨ltta¨ 1988; Rastas & Kilpela¨inen 1991). The second example (Fig. 10) is from the central part of the Central Finland Granitoid Complex, where the FIRE profiles 3a and 1 & 2 cross nearly at right angles (Figs 2 & 3). The area is mainly composed of the two post-collisional suites of granitoid intrusions and associated gabbroic and volcanic rocks that have been deformed and displaced to various amounts. In this area, FIRE 3a trends NW– SE (Fig. 3) and it is perpendicular to Group 1 structures (incident angle c. 908), parallel to Group 2 structures (incident angle c. 08) and oblique to the Group 3 and 4 structures (incident angle c. 458). The upper crust is dominated by several generations of SE-dipping listric reflections and associated synthetic and antithetic reflection fans (purple, light purple; Fig. 10). The reflections also suggests development of extensional duplexes below structural highs (blue; CMPs 11300 – 12100). A core complex might be exposed in the western part (brown; CMPs 13000–13400), where the listric shear zones flatten close to the surface and cross-cut an anticlinal structure that has

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Fig. 12. Block diagrams displaying the upper crustal structure at the crossing of the FIRE profiles. Listric shear zones are in blue/violet and transfer faults in stippled purple. Their surface counterparts give rise to Group 1 and 2 structures identified on a low-altitude aeromagnetic map (Fig. 3). North and south dipping conjugate faults in red and orange are associated with Group 4 structures. (a) View is to the north; (b) view to the south.

developed in the footwall of the listric reflections. The de´collement surfaces, onto which the listric reflections sole out, become shallower towards the west and also cross-cut reflections soling out at deeper levels, indicating that the structures become younger towards the west. The same younging direction is also implied by the development of extensional duplexes. The upper crustal reflection image along FIRE 3a (Fig. 10) is similar to those developed in extensional environments in corecomplex areas (Allmendinger et al. 1983; Cook et al. 1992; Varsek & Cook 1994). The older calc-alkaline magmatic suite is pervasively deformed, dipping SE to ESE with a shallow to moderate ESE plunging lineation (Pipping & Vaarma 1993; Sorjonen-Ward 2006). At CMP 12450 (Fig. 10), where one of the listric reflections is coming to the surface, the survey cut across Group 1 magnetic lineament rising from strongly sheared sheared SE-dipping felsic metavolcanic rocks (strike 0508, dip 120/408) and metagabbros

(Fig. 4a). This suggests that the listric reflections are, at least partly, related to deformation patterns found in the older calc-alkaline suite. The larger porphyritic alkaline granite intrusions (Figs 2 & 3) are emplaced in graben structures bordered by listric synthetic and antithethic fans (CMPs 10900 –11600 and 11800–12250; Fig. 10); minor granite intrusions are found in half-grabens floored by synthetic faults. The host rock is observable in the central highs or horsts, which are underlain by extensional duplexes (CMPs 11400 –12000). The horsts are composed of older, deformed granitoids of the calc-alkaline suite (containing rafts of folded schist), and associated with well-preserved felsic volcanic rocks that display only brittle deformation. Pairs of brittle normal faults are found in the central highs; a clear example can be seen in a quarry in the Kalmari volcanic sequence NW of Saarija¨rvi, where two normal faults (dip 225/308, Group 2 or 5) transect a fine-grained plagioclaseporphyry (Fig. 4h).

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Where the seismic sections suggest an anticlinal core of the middle crust (CMPs 13000–13400), the bedrock is composed of highly strained calcalkaline granitoid intrusions with pervasive SE-dipping foliation and shallow to moderately east-plunging lineation (Photo 6 in Figs 3b & 4i; Pipping & Vaarma 1993; Sorjonen-Ward 2006). Sorjonen-Ward (2006) called this area the Kyyja¨rvi deformation zone; but based on its structural position and the regional continuity of the strain patterns commonly observed in the domes of core-complexes (Coney 1980), we reinterpret the Kyyja¨rvi deformation zone as a core complex dome that may have originated as an anticlinal expulsion of the middle crust. Where FIRE 1 & 2 crosses FIRE 3a it trends NNE-SSW (Fig. 3b) and parallels Group 1 structures (incident angle c. 08), is perpendicular to Group 2 structures (incident angle c. 908) and is oblique to Groups 3, 4 and 5 structures (incident angle c. 458). The seismic section (Fig. 11) is characterized by three sets of structures: (1) subhorizontal to shallowly south-dipping, high-amplitude reflections that are displaced upwards (purple) by (2) south/north-dipping low-angle reflections (magenta/orange) or by (3) subvertical transparent zones (stippled) that displace all the above structures (Fig. 11). Both the low-angle and subvertical structures seem to initiate at locations associated with Group 2 shear zones (Fig. 4b, photo 2). Because the subvertical shear zones displace the listric shear zones (seen as flat surfaces because of geometrical projection; see Fig. 12) they are interpreted as transfer faults. The alkaline granite plutons seem to have been emplaced in grabens outlined by low-angle faults (Fig. 11). The exposed margins of the granite plutons are tectonic contacts with moderately to steep plunging lineation. At least locally, they are somewhat younger than the calc-alkaline plutons (Fig. 4b, photo 2). The sections suggest that the subvertical transfer faults close to the margins of the plutons developed as break-away faults to the low-angle fault system. The three dimensional relationships of the reflection structures are illustrated on the schematic block diagram (Fig. 12). The north facing block diagram (Fig. 12a) is dominated by the NE–SW-striking, SE-dipping listric shear zones (Group 1) displaced by the NW–SE-trending transfer faults (Group 2). Both the listric shear zones and the transfer faults terminate downward at the upper –middle crust detachment zone (U –M). A conjugate set of lowangle faults dipping both south and north (Group 4), seem to displace the listric faults (Group 1). The block diagrams image the granites as emplaced in double-sided grabens controlled firstly by the listric shear zones and their synthetic and antithetic

faults and secondly by the graben-forming lowangle faults.

Discussion As in many other orogenic belts (Cook et al. 1979, 1992, 1999; Brown et al. 1996; van der Velden & Cook 1999; Meissner 1996; TRANSALP Working Group 2002), also in the Svecofennian the deformational patterns recognized in the supracrustal segments (Park et al. 1984; Sorjonen-Ward 2006) at surface are dominated by collision related thrust, fold and stacking patterns whereas the crustal seismic sections outline a structure dominated by lateral spreading structures (Figs 5–8). Beaumont et al. (2001, 2004) and Culshaw et al. (2006) suggested that this apparent paradox could be explained by the development of superstructure –infrastructure layering at deeper levels of the crust during the later stages of the orogeny. In the following, we will adapt the superstructure – infrastructure model to the Svecofennian orogen and discuss how it fits the magmatic, metamorphic and deformational observations at surface.

Formation of layered Svecofennian crust with superstructure– infrastructure The coeval convergence from both east– west and north– south directions (Fig. 1b; Lahtinen et al. 2005) probably resulted in extreme over-thickening of the crust and lithosphere in the common hinterland. The thickening resulted in excess of gravitational potential energy (GPE) that drives gravitational collapse (Rey et al. 2001). Stacking of crustal material may also have resulted in isostatic imbalance as the density stratification of the crust was disturbed. The Svecofennian crust is, however, in quasi-isostatic equilibrium (Korsman et al. 1999) and the area is mostly aseismic and the current insignificant intracontinental deformation is commonly attributed to post-glacial uplift or spreading of the Atlantic Ocean sea floor. Thus the late stage orogenic collapse event assumed here succeeded in balancing both the vertical and horizontal gravitational inequilibria. The Svecofennian crust has three layers and their internal boundaries have acted as detachment zones. The middle to lower crustal boundary (M-L) is associated with a velocity jump from 6.6 to 6.8 km s21, which coincides with a change from intermediate to mafic composition (Christensen & Mooney 1995; Mooney et al. 1998; Kuusisto et al. 2006). The deformation has concentrated at this boundary, because it is associated with a change from quartz-dominated to plagioclase-dominated lithologies (Kusznir & Park 1987). The upper to

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middle crustal boundary (U-M) is associated with a velocity jump from 6.2 to 6.3 km s21. This velocity change indicates a change from felsic to intermediate lithologies (Christensen & Mooney 1995). On the other hand, the detachment zone is likely to form where the rheology changes from wet-quartzdominated to dry-quartz-dominated (Kusznir & Park 1987), i.e. from water-bearing mineralogy to water-absent mineralogy or from amphibolite facies to granulite facies (Vanderhaeghe 2001). According to Kusznir & Park (1987), the concurrent formation of mid- and upper-crustal detachment zones (M-L, U-M) suggests high geothermal gradients (70–80 mW m22). Before freezing, extension has thinned the upper and middle crust from 45 km at the edges (Fig. 4; CMP 12000) to 25 km in the centre (Fig. 4; CMPs 16000–18000). Slow to intermediate strain rates (10216 –10215 s21) in regions with moderate to high geothermal gradients should produce wide regions of strain hardened, stable crust (Kusznir & Park 1987). A net strengthening may lead to cessation of extension at one locality and spreading of deformation to adjoining areas with weaker, thicker crust (Kusznir & Park 1987), which may explain the minor age differences of the late-stage post-collisional granite intrusions between adjacent regions of the Central Finland Granitoid Complex. Stacking of hot, juvenile crustal pieces results in a warm/hot, thickened crust (Vanderhaeghe & Teyssier 2001a, b; Thompson et al. 2001). At depths and temperatures where granulite facies is reached (20 km, 750 8C; England & Thompson 1984), anatexis begins and migmatites are formed. A significant change in rheological properties takes place at this migmatite front, where lithologies change from wet to dry or from amphibole- and mica-bearing to pyroxene-bearing (Vanderhaege & Teyssier 2001a; Jamieson et al. 2004) and a detachment surface may form (Beaumont et al. 2004). Below the detachment, the crust melts increasingly and produces calc-alkaline granitoid suites from different melt portions and at different crustal levels (Harris et al. 1986; Clemens & Vielzeuf 1987; Bonin et al. 1998). The melts pond below the detachment zone (Vanderhaege & Teyssir 2001a) thus facilitating the decoupling of the middle and upper crust. The temperature increase with depth will, however, be buffered by energy consuming melting reactions and may thus not increase beyond 900 8C (Vielzeuf et al. 1990) in the underlying crustal layers, which attain granulitefacies mineralogy. After the melting episode, the Svecofennian middle and lower crust is composed of migmatites, melt-depleted residual granulite and igneous rocks crystallized from trapped magmas, whereas the upper crust is composed of greenshist and amphiobolite facies rocks intruded by

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calc-alkaline granitoid plutons that managed to escape from the middle crust. Voluminous calc-alkaline granitoid magmas have intruded the folded upper crustal rocks in the Central Finland Granitoid Complex and the surrounding schist belts at 1.89– 1.88 Ga. They are usually referred to as synkinematic or synorogenic granitoid rocks in the Finnish geological literature. Below the Central Finland Granitoid Complex, a low velocity layer (LVL; Grad & Luosto 1987), is found below the highly reflective U-M detachment zone. The LVL has similar seismic properties with the low velocity layer beneath the Tibetan plateau (Brown et al. 1996) interpreted to originate from partial melts and trapped magma (Wei et al. 2001). We assume that the Central Finland low velocity layer has a similar origin and interpret it to be composed of trapped granitoid intrusions. The rheological behaviour of the middle crust is further complicated by the increased amount of melt. The migmatite is able to flow as viscous fluid as soon as the amount of melt is sufficiently high that the rock loses its solid framework (Brown 1994; Sawyer 1998; Vanderhaeghe 2001). During this anatectic phase, the middle crust is easily deformed and flow structures –midcrustal thrust ramps on FIRE 3a (Fig. 6) and herringbones on FIRE 1 & 2 (Figs 5 & 7) – develop and the crust thins by spreading laterally both forward and sideways. The melting event and the contemporaneous development of the superstructure-infrastructure (Culshaw et al. 2006) enhances the initial layering of the crust. The upper crustal material also slides away from the thickening centre (two-sided orogen and orogenic triple point; Fig. 1b) via the listric upper crustal shear zones and associated transfer zones and via low angle normal faults (Figs 6, 8 & 10). The regional average of the plunging directions of lineations in the northern part of the Central Finland Granitoid Complex (Kilpela¨inen et al. 2008) image persistent movement either towards or away from the Ladoga–Bothnian Bay wrench fault zone (LBB; Figs 1 & 2). The lineation arrows are parallel to the zone within it but rotate to ESE away from the zone. The few observations that we have from the area indicate that the hanging wall was moving towards the west, suggesting that the upper crust was moving upwards and westwards. This direction of spreading is also suggested by the sequential overlapping (younging) of listric shear zones and the formation of extensional duplexes (Fig. 10). The upper crustal structure of the Central Finland Granitoid Complex resembles that of metamorphic core complexes (Coney 1980; Allmendinger et al. 1983; Gans et al. 1985). At surface, the geology could be described as a deeply exhumed

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(lower structural level) metamorphic core complex, where the basins are occupied by the alkaline granites, and where the older calc-alkaline granitoid intrusions and associated volcanic rocks form the central highs within the basins. If the strained granitoid intrusions were parts of the lowest level of the spreading upper crust or parts of the U-M detachment zone, they would be the upper parts of metamorphic domes. Migmatitic cores representing the middle crust would be exposed at the margins of the extending orogenic core. Theoretically, when the hot collapsing orogen started to cool, it would have changed from a thermal dome to a large basin. This phase could explain the late-stage brittle normal faults as well as the deepening of the detachment surface in the middle of the Central Finland Granitoid Complex (Fig. 5). According to Culshaw et al. (2006), the age of metamorphism grows younger with depth and with distance from the suture zone (triple point). The seismic crustal structures on the FIRE profiles suggest middle crustal migmatites to be exposed at the northern and western parts of the Central Finland Granitoid Complex. In the north, the Pyha¨ja¨rvi shear zone exposes the Pyha¨ja¨rvi diatexitic migmatites (Fig. 9) with anticlockwise metamorphic paths and metamorphic ages somewhat younger than the peak of the post-collisional calc-alkaline granitoid intrusions (‘synorogenic’ in Korsman et al. 1999; Ho¨ltta¨ 1988). In the west (Bothnian Belt; Figs 1 –3), the seismic profiles suggest a large core complex or at least anticlinal uprise of the middle crust, where a large diatexitic migmatite complex, the Vaasa Migmatite Complex is exposed (Ma¨kitie 2000; Korsman et al. 1999). Also here, the metamorphic ages (1.88–1.86 Ga) are slightly younger than those of the calc-alkaline granitoid rocks. Kimberlite pipes intruding the collisional boundary between Savo Belt and Karelian continent have sampled the middle and lower crust between the depths of 22–38 km. Single crystal age determinations on zircons with U –Pb and Pb –Pb methods outline age populations that are similar to those found at the surface (Ho¨ltta¨ et al. 2000). The observed Svecofennian ages (1.89, 1.88, 1.87 Ga) are somewhat younger than the syncollisional calc-alkaline magmatism (1.89–1.88 Ga) and they agree better in age with observed ages of alkaline granites and metamorphism (1.88–1.87 Ga), thus suggesting that a slightly younger infrastructure developed in the middle and lower crust. The Central Finland area was also subject to an intrusive magmatic event originating in the mantle and exemplified by the 1.89– 1.88 Ga gabbroic intrusions and dykes (Fig. 2; Korsman et al. 1997; Puranen et al. 1992) that left highly reflective tails behind (Figs 5 & 7b; CMPs 15800 and 18000).

The area of under- and intraplate is characterized by poorly reflective lower crust (FIRE 1 & 2 CMPs 14000 –22000; FIRE 3a CMPs 2000–14500) with a high P-wave velocity (Vp . 7.2 km s21) and high Vp/Vs ratio (Hyvo¨nen et al. 2007). This external heat source raised the lower crustal temperatures further and initiated partial melting of lower crustal granulites giving rise to the post-collisional late-stage, high-temperature alkaline, pyroxenebearing granites with estimated magma temperatures of over 900 8C (Elliott et al. 1998). The source of the mantle magmatism could be either decompressional melting of mantle rising passively as the crust above is thinning; an alternative hypothesis for the melting event is subduction slab break-off and subsequent rise of the asthenosphere (Platt & England 1994). This is beyond the scope of this paper. The alkaline granites were emplaced in pressure minima between the flanking faults of grabens controlled by Group 1–4 structures. On one hand, the granites were emplaced in half grabens controlled by synthetic and antithetic normal faults (Group 1) and their transfer faults (Group 2) in the hanging walls of listric laterally extending shear zones; on the other hand, they intruded in grabens controlled by a set of conjugate low-angle faults dipping N or S (Group 4) (Figs 3b & 10 –12). This implies that all the fault systems were operating at the same time and that the upper crust was spreading in NW– SE and north –south directions contemporaneously. The same picture emerges from the distribution of Group 1 and Group 2 shear zones over the entire Central Finland Granitoid Complex (Fig. 1a). We suggest that the fanning out of the shear zones images the spreading of the orogen in north–south direction and towards NE, away from the palaeotriple point of the double-sided collision (Fig. 1b).

The mode of extension The Fennoscandian shield is underlain by thick lithosphere with anomalously high P- and S-wave velocities and large lateral variations in velocities and anisotropy (Alinaghi et al. 2003; Bruneton et al. 2004; Plomerova et al. 2006). The anisotropy vectors (Plomerova et al. 2006) seem to coincide with the suggested palaeoplate movement directions (Lahtinen et al. 2005). The high velocity anomaly (70– 150 km in depth) has the shape of a saddle with the long axis (250 km) in NE–SW direction and the short axes (120 km) in NW–SE direction. Within the anomalous lithosphere, P-wave velocities are 2 –3% higher, S-wave velocities are 2–3% lower than the mean. Although the high velocity anomaly is spatially associated with the Archaean –Proterozoic suture zone, it transects the present surface suture zone at a 458 angle suggesting

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that its shape might be related to post-collisional rather than collisional processes. A post-collisional melting event in the mantle would lead to relative increase of MgO component in olivine because the Fe-rich component melts at lower temperatures. The increase in MgO component is detected as linear increase in seismic P-wave velocity, decrease in density and Vp/Vs ratio (Lee 2003). Xenoliths coming from the mantle wedge record high degree of melt depletion (Peltonen & Bru¨gman 2006). The tectonic setting could be related to break-off of a subducting slab, thermal relaxation of the over-thickened lithosphere, or back-arc spreading associated with an arc developing to the southern side of the orogen or any combination of these. The anomalous lithospheric mantle is complemented by crustal anomalies. It is overlain by thinned crust with high velocity lower crust (Vp . 7.0 km s21; Korja et al. 1993), positive Bouguer anomalies (Kozlovskaya et al. 2004) and a small density contrast with the mantle; all suggestive of excess of mass in the lower crust. The high densities and velocities of the lower crust have been explained by mantle-derived under- and intraplating (Korja 1995; Korsman et al. 1999) and by older, mafic crustal root (Lahtinen et al. 2005; Korja & Heikkinen 2008). Mafic underplating is indicated by a bimodal suite of alkaline granites and associated gabbroic rocks (Fig. 2), whereas older crustal component has been attributed to negative 1 values (Lahtinen et al. 2005). The geometry of the Svecofennian lithosphere is a mixture of all the extensional end member types: narrow rift, wide rift and core complex (Buck 1991). The narrow rift mode is suggested by the preservation of a 200 km wide Moho rise (Fig. 5), coupled with a 250 km wide high velocity lithosphere anomaly. At surface, the area is characterized by post-collisional bimodal magmatism. The alkaline intrusions are derived from high temperature melting of the lower crust (Elliot et al. 1998) indicative of high heat flow. High heat flow is also implied by the formation of mid- to upper crustal detachments. In the orthogonal direction (SE –NW), the Moho topography is smooth and the extension is distributed over a 400 km wide area whereas the mantle anomaly is only 120 km wide. In this direction, extension was probably in wide rift mode, which implies either slow strain rates or non-uniformly distributed strain (Kusznir & Park 1987; Corti et al. 2003). Because the high heat flow was accompanied by slow strain rates, the extension had time to be distributed over larger volume and instead of a whole lithosphere failure, the crust attained the work-hardened structure of a wide rift.

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Core complex mode (Buck 1991) may be suspected in the northern and western parts of the Svecofennian orogeny, where middle crustal lithologies locally extrude to the surface. Core complex mode may require that lateral spreading of the upper crust is detached from the middle to lower crustal extension and that magmatic underplating smooths out the variations in Moho topography. The results from the Central Finland Granitoid Complex indicate, however, that the core complex mode can also develop above partially molten middle crust and that the contribution of the decoupled lower crust may be smaller.

Conclusions The coeval continental convergence from both east and south (Fig. 1b) probably resulted in extreme over-thickening of the Svecofennian crust and lithosphere. Over-thickening was partially compensated by frontal and sideways spreading of the orogen away from the triple point. The spreading and thermal relaxation resulted in orogenic collapse, during which the crustal material was rearranged and superstructure-infrastructure of the crust was formed. The crust attained its present layered structure, where HT –LP amphibolites-facies rocks characterize the uppermost crust, granulite-facies rocks the middle and lower crust. The spreading of the hot orogen resulted in a lateral extensional environment with high heat flow, and slow strain rates leading to formation of upper and middle crustal detachments and workhardened structure of a wide rift. The upper crust displays the features of a core complex. The upper, middle and lower layers of the crust were decoupled and (they) extended in different fashions. The upper crust spread in a brittle to ductile regime along listric, low-angle and transfer shear zones. The middle crust thinned via ductile flow and extrusion. The lower crust, Moho and upper mantle rose to fill the space created by the thinning middle crust. Later, this structure was modified by intrusion of mantle-derived magmas. The middle crust displays typical large scale lateral flow structures: herringbone and anticlinal ramps. Both structures are rooted to large scale listric surfaces that probably originated as collisional stacking surfaces and evolved later into extensional spreading surfaces. The Central Finland Granitoid Complex is a superstructure that developed above the upper – middle crust detachment zone. The complex is a deep, lower level section of an old core complex, where the younger granitoid intrusions form the basins and the older ones with their associated volcanic rocks form the horsts. The middle crustal

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cores of infrastructure are exposed in the migmatite blocks surrounding the complex in the west and north. The upper– middle crust detachment zone is exposed at the Ela¨ma¨ja¨rvi shear zone and probably at the Kyyja¨rvi deformation zone. This work is part of project SA 20472 financed by the Academy of Finland. Additional field work was funded by K. H. Renlund Foundation. The Geological Survey of Finland provided the low altitude aeromagnetic anomaly map and a large set of existing observation data. Kaisa Wanne and Timo Kilpela¨inen from the University of Turku are acknowledged for their valuable contribution to field observations. The manuscript benefited from constructive criticism by Christopher Talbot and an anonymous reviewer.

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Index Page numbers in italics refer to Figures. Page numbers in bold refer to Tables. Acidere Formation 205 Acigo¨l Fault 199 acoustic fluidization 17 Aegean extensional province tectonic map 198 tectonic setting 197 –198, 200 Aegean Sea geodynamics 181 –183 geological map 140 tectonic map 170, 181 Agridi Fault 126, 127, 129, 130, 131 Akrata Fault 129 Alasehir detachment 199, 205 Alasehir graben 215, 217 Alasehir granitoid 199, 202 Alpine orogeny 34 – 36 Amonton’s Law 11 –12, 17 Amorgos detachment geological setting 171 –173 map 170 timing of detachment methods of analysis 173 fission track analysis 173, 175 results 174 results discussed 175 significance of results 175 – 176 amphibolite facies, Svecofennian 226, 229, 245 Anatolia cross-section 216, 217 geological setting 200 –201 map 199 Anatolide Block 197, 198, 199, 200 andesites, origins of low/high K 74 –76 Andre´ fault 141, 144, 145, 154 apatite FTA see fission track analysis apatite (U –Th)/He dating studies on Ios method 150– 151 results 154 significance of results 158 – 162 Appalachians orogen 2 apparent friction 12 40 Ar/39Ar dating 114, 202 –203 asperities and flash heating 16 Attic –Cycladic Massif 180, 181 Austroalpine nappe system 34 magmatism 40, 41, 42 metamorphism 36, 37, 38 Baklan granitoid 199, 211, 212, 213 Basal Conglomerate Unit 171, 172, 173 Basin and Range Province 2, 3, 4 low-angle normal faults 9 Bergslagen plate 227 Biga Peninsula 200 Bothnian Belt 229 boudins see Varvara boudin brittle/ductile behaviour 11 Burdur Fault 198, 199 Bu¨yu¨k Menderes graben 215, 217 Byerlee friction 9 carbon dioxide, role in faulting 11 cataclasite 11 Cataldag granitoid 199, 217

Catalina-Rincon Mountains (Arizona, USA) 4 Central Anatolide Crystalline Complex 198 Central Finland Granitoid Complex 229 crust detail 242, 245, 246 structure interpretation 237, 238 Central Volcanic Region (New Zealand) 74, 79, 81 –83 chlorite, and fault strength 15 clays, reactions and fault strength 15 Coastal Fault System 141, 144, 145 collisional modelling 225 –226 continental back-arc systems 73 conceptual model 74 studies in New Zealand, extension in relation to local variables crustal structure 78 –81 extension and arc migration 78 extension – rotation relations 78 geodetic strain estimates 76 –77 GPS strain estimates 77 – 78 heat output 81 –83 palaeomagnetic rotations 78 summary of observations 83 –85 tectonic setting 73 –74 volcanic arc migration 74 –76 cooling, and crustal thinning 4 core complexes 4 and tectonites 5 Corinth Rift 119, 120 cratonization 226 Cretan detachment 169, 170 crust structure in New Zealand 78 –81 thinning and cooling 4 Cyclades, regional geology 141 Cycladic Basal and Upper Units 140, 141 Cycladic Blueschist Unit 140, 141, 170 Dara Basin 127, 129, 130 Datc¸a– Kale Fault 198, 199, 217 decollements, seismic character 233 detachment era, defined 1 detachment horizons 10 detachment zones seismic character 233 Svecofennian orogen 241, 242, 244 Dixie Valley (Nevada, USA) 10 Drossopigi Fault 125 dynamic failure 14 dynamic fragmentation 19 –22 earthquakes, and low-angle normal faults 9 East Anatolian Fault Zone 198 East Greenland, Caledonides 2 Ecemis Fault 198 Egrigo¨z granitoid 199, 201 – 202, 211, 212, 213, 214 Eliki Fault 129, 132, 133, 134, 135 Eromonsia Formation 186, 188, 189, 193 Erzurum-Kars Plateau 198 Eu anomaly 214 European Alps 33 geology Alpine evolution 34 –36

Permo-Triassic evolution 36 magmatism 39 – 44 metamorphism 36 – 39 sedimentary basins and volcanics 44 –47 map 35 numerical analysis of geological evolution model set-up 47 –49 thermomechanics 50 –52 numerical and natural data compared 52 –53 magmatism 58 –64 metamorphism 53 –58 sedimentary basins and volcanics 64 Evciler granitoid 199 Evia 140, 141 exhumation, factors affecting 179 extension and isostasy 4 relation to rotation 78 relation to volcanic arc migration 78 extension fault, defined 169 extension rate, effect of 226 extensional collapse 225 –226 External Hellenides 170 extrusion wedge, defined 169 Eybek granitoid 199 fault breccia 11 fault core, defined 11 fault gouge 11 feldspar, reactions and fault strength 15 Feneos Basin 127, 128 fission track analysis (FTA), apatite and zircon Amorgos detachment study methods 173, 175 results 174, 175 –176 Naxos study methods 184 results 184, 189 granodiorite 184, 185, 187 metamorphic rocks 184, 185, 187 sediments 184, 186, 188 results discussed granodiorite 192 – 193 metamorphic rocks 189 –192 sediments 193 fixism and fixist concept, rise and fall of 1–3 flash heating 16 fluids, effect of at faults 12 high pore pressure 13 –14 Flysch Unit 172 folds, upright v. flat-lying 226 friction, coefficient of, for low-angle normal faults 9 friction in LANF, mechanics of 11 –12 evaluation of mechanisms, granular friction 17 –19 dynamic rock fragmentation 19 – 22 mechanical process 16 –17

254 friction in LANF, mechanics of (Continued) requirement for reduction 12 requirements for low strength 12 –13 heating 15 –16 high pore fluid pressure 13 –14 weak faults materials 14 – 15 summary and discussion 22 –25 Gardiki Fault 128, 132, 133 geochemistry Menderes massif study 203, 206 methods of analysis 206 results 206 –209 Zealandia studies Eastern Province Eweburn Tuff 101 –102, 105, 106, 107 Houhora Complex 102 –103, 105, 106, 107 Motu Tuff 103, 104, 105, 106, 107 Shag Valley Ignimbrite 101 –102, 104, 105, 106, 107 Western Province Canavans Quartz Monzonites 102, 105, 107 DSDP rhyolite 103, 104, 107 Stitts Tuff 103, 104, 105, 106, 107 Whataroa Granite 103, 106, 107 geochronology studies on Ios Rb/Sr geochronology introduction 148– 149 methods 149 results 149 –150 thermochronology apatite (U –Th)/He dating method 150 –151 results 154 fission track analysis methods 151, 153 results 152, 153 results discussed age-slip relations 155 –158 cooling history 154 –155 temperature –time relations 158 –159 studies in Menderes granitoids 202 –203 studies in Zealandia Eweburn Tuff 96, 97, 98, 99, 100, 101 Houhora Complex 92, 93, 95 –96, 99 Motu Tuff 100 Shag Valley Ignimbrite 96, 99, 100, 101 Western Province Canavans Quartz Monzonite 93, 95, 96, 100 DSDP rhyolite 93, 95, 96, 99, 100 Stitts Tuff 93, 95, 96, 100 Whataroa Granite 93, 95, 96, 99, 100, 101

INDEX geodesy, strain measurement in New Zealand 76 –77 Go¨ynu¨kbelen granitoid 199 Gondwanaland, break-up 89 GPS, strain measurement in New Zealand 77 – 78 granular friction 17 –19 granulite facies, Svecofennian 226, 229, 245 graphite, reactions and fault strength 15 gravitational collapse 225 –226 Svecofennian Orogen 244 gravitational potential energy, of thickened crust 225 – 226, 244 Greece see Amorgos detachment; Ios; Naxos; Peloponnesus, Northern Gu¨ney detachment 199 Gu¨rgenyayla granitoid 199, 215 Headland Shear Zone 144 –146, 158 heat, effect on fault strength 15 –16 heat output, convective discharge in New Zealand 81 –83 Hellenic trench 197 Helvetic Domain 36, 40, 46 hydrothermal fluid, role in faulting 11 Ilica granitoid 199, 217 illite, and fault strength 15 Indonesian archipeligo 226 infrastructure-superstructure concept 1–2, 226 applied to Svecofennian 244 –246 Inner Tauride suture 198 Intra-Pontide suture zone 198, 199 Ios Detachment Fault 141, 144, 146 Ios metamorphic core complex experimental studies Rb/Sr geochronology introduction 148 –149 methods 149 results 149 –150 thermochronology apatite (U –Th)/He dating method 150 –151 results 154 fission track analysis methods 151, 153 results 152, 153 results discussed age-slip relations 155 –158 cooling history 154 –155 temperature –time relations 158 –159 extensional setting 139 –140 geological setting 141, 142 deformation history D1 143 D2 142 –144 D3-5 144 metamorphic history M1 141 –142 M2 142 –143 petrographic analysis 147 P –T calculation 147 – 148 structural analysis 144 –146 timing of events geometry variations 159 –161 tectonic implications 161 –162 temperature variations 158

isostasy, and extension 4 isotherms, relation to detachment 179 –180 Izmir–Ankara suture zone 198, 199, 200, 217 Japan Sea, Miocene rotations 74 Kamenitsa Fault 126, 129, 133 Kandila Basin 127, 129, 130, 132 Kapidag granitoid 215 Karabiga granitoid 199 Karelia plate 227 Karliova triple junction 198 Katrandag granitoid 199 Kazdag core complex 200, 201, 215, 217 Kazdag mertamorphic massif 199, 200 Kefalonia Fault 198 Keitele microplate 227 Kestanbol granitoid 199 Khelmos Fault 124, 126, 132, 133, 134 Koyunoba granitoid 199, 201– 202 Kozak granitoid 199 Krathis Fault 124, 126, 128 Ku¨c¸u¨k Menderes graben 215, 217 Kusc¸ayiri granitoid 199 Laramide Ranges 2 Levidi Basin 127, 130, 132 Levidi Fault 126, 127, 129, 133 listric faults 10 lithosphere, growth and classification 225 Lord Howe Rise see Zealandia low-angle normal faults (LANF) angle of plane 9 defined 10 initiation 11 relation to friction coefficient 9 relation to relief 9 tectonic settings 11 Lycian ophiolite nappes 140, 199, 200, 201, 216, 217 Lykouria Fault 127, 129, 132, 133 magmatism European Alps geology 39 –44 numerical modelling 47 –49, 50 –52 numerical modelling and geology compared 58 –64 mantle degassing, role in faulting 11 mantle fluid, role in faulting 11 Maras triple junction 198 Marble Unit 171, 172, 173 Marmara granitoids 199, 200 –201 Median Batholith (Zealandia) 89, 90 melting, and fault strength 16 Menderes metamorphic massif 198, 199, 201 granitoids 201 field relations 201 –202 geochemistry 203, 206 methods of analysis 206 results 206 –209 geochronology 202 –203 geodynamic interpretation 215 –216 petrogenetic history fractional crystallization 214– 215 melt source and evolution 212– 213

INDEX previous research 209 –211 Sr –Nd isotopes 213– 214 map 203 Menderes Nappe 170 Metabasite Unit 171, 172 metamorphic core concept, first defined 1 metamorphic facies, Svecofennian 226, 229, 245 metamorphic tectonites 5 metamorphism HP–LT Naxos 180, 182 western Anatolia 200 HT –LP 33 European Alps geology 36 – 39 numerical modelling 47 –49, 50 –52 numerical modelling and geology compared 53 –58 MP –MT, Naxos 182 –183 Svecofennian 226, 229, 245 meteoric fluid, role in faulting 11 mica, reactions and fault strength 15 migmatite, Svecofennian orogen 245 minerals, effect on fault strength 14 – 15 Mohr –Coulomb criterion 11 Monodendri Fault 125, 126 montmorillonite, and fault strength 15 Naxos exhumation evidence 183 –184 fission track analysis study methods 184 results 184, 189 granodiorite 184, 185, 187 metamorphic rocks 184, 185, 187 sediments 184, 186, 188 results discussed granodiorite 192 –193 metamorphic rocks 189 –192 sediments 193 geological setting 180 –181 intrusives 183 landslides 183 sediments 183 tectono-metamorphic units 181– 182 M1 Eocene HP –LT 182 M2 Miocene MP –MT 182 –183 1Nd plots 101, 104, 106, 109, 213, 214 Neotethyan ophiolites 199 New Zealand formation 89 North Island crustal structure 78 –81 extension and arc migration 78 extension–rotation relations 78 geodetic strain estimates 76 –77 GPS strain estimates 77 –78 heat output 81 –83 palaeomagnetic rotations 78 summary of observations 83 –85 tectonic setting 73 –74 volcanic arc migration 74 –76 see also Zealandia, Eastern Province normal faults 1, 9, 169 see also low-angle normal faults North Anatolian Fault Zone 198, 199, 216

North Mainalon Fault Zone 127, 131, 132, 133, 134, 135 Orchomenos Fault 127, 130 Orhaneli granitoid 199, 215 orogenic collapse 225 –226 orthogonal pairs 229 Otago Schists 89 Ovacik Fault 198 palaeomagnetism, measures of rotation in New Zealand 78 Pampak –Sevan fault 198 Pelagonia 197 Pelagonian/Lycian Zone 170 Peloponnesus, Northern 120 history of research 119, 121 lithostratigraphy 121 –124 significance of fault geometry 128 –135 structure and fault geometry 124 –127 Penninic Domain 34 magmatism 41 metamorphism 36, 37 Penteleion Fault 132, 133 Pesulia Formation 184, 186, 188, 189, 193 Phyllite–Quartzite Unit 169, 171 Pindos thrust 125, 127, 129, 130 pore fluid pressure, effect on faulting 13 –14 potential energy see gravitational potential energy Precambrian orogens, characteristics 226 pseudotachylyte 11, 16 Pyrgaki –Mamoussia Fault 132, 133, 134, 135 quartz gouge, and fault strength 15 Rb/Sr geochronology studies on Ios introduction 148 – 149 methods 149 results 149 –150 Rhodope massif 216 rhyolites see Zealandia, felsic volcanics Riedel shears 229 Saitias Fault 132, 133 Sakarya continent 197, 199, 200, 215, 216, 217 Salihli granitoid 199, 202 –203, 204, 205, 207, 208, 209, 211, 212, 213, 214 Samos 140, 141 San Andreas Fault 15, 241 San Andreas Fault Borehole Observatory at Depth (SAFOD) 5 Savo arc 227, 229 seawater, role in faulting 11 serpentinite, and fault strength 15 Sevketiye granitoid 199 shear failure 12 silica gel, and fault strength 15 Simav detachment 199, 209, 215 Simav graben 215, 217 smectite, and fault strength 15 Snake Range (Nevada, USA) 3 South Cyclades shear zone 139, 141, 145, 146, 170 Southalpine Domain 34

255 magmatism 40, 41, 42, 43 metamorphism 36, 38, 39 sedimentary basins and volcanic 45, 46 Southern Finland Arc Complex 227 Southwest Anatolian shear zone 199 Sr anomaly 213, 214 87 Sr/86Sr 106, 109 static failure 13 –14 stockwerk folding hypothesis 1– 2 strain hardening 226 conditions for 245 strain measurement geodetic 76 –77 GPS 77 –78 strength, defined 12 stress, normal 12 strike-slip faults, large displacement 1 Stymfalis Basin 127 superstructure-infrastructure concept 1–2, 226 applied to Svecofennian 244 –246 Svecofennian Orogen aeromagnetic anomalies and lineament recognition 229 –233 crustal features interpreted 241 extension mode 246 –247 lower 241 middle 241 –242 superstructure-infrastructure evolution 244 –246 upper 242 –244 introduction 226 –229 seismic data 233, 241 crustal structure interpretation 237, 238, 239, 240 reflection profiles 234, 235, 236 talc, reactions and fault strength 15 Tampere arc 227 Tampere Belt 229 Taupo Volcanic Zone 74 Tauride Block 199 Tavsanli blueschists 216, 217 Tethyan ophiolites 217 thermal energy, of thickened crust 225 –226 thermal pressurization, effect on fault strength 15 – 16 thermochronology, principles of 179 –180 apatite (U–Th)/He dating on Ios method 150 –151 results 154 significance of results 158 –162 fission track analysis (FTA), apatite and zircon Amorgos detachment study methods 173, 175 results 174, 175 – 176 Naxos study methods 184 results 184, 189 granodiorite 184, 185, 187 metamorphic rocks 184, 185, 187 sediments 184, 186, 188 results discussed granodiorite 192 –193 metamorphic rocks 189– 192 sediments 193

256 thrust faults, large displacement 1 thrust sheet geometry, Peloponnesus 128 –135 thrust-and-nappe structure 2 Tinos 140, 141 Topuk granitoid 199, 215 Torlesse Terrane 89 trace elements, discriminant plots 106, 107 Tripolitza Unit 169, 170, 171 Tsivlos Fault 124, 130, 132, 133, 134 Turgutlu granitoid 199, 202 –203, 211, 212, 213, 214 Tutak Fault 198 Tuzgo¨lu¨ Fault 198 U–Pb ICPMS methods 113 –114 U–Pb TIMS methods 112 –113 U–Th/He dating of apatite studies on Ios method 150 –151 results 154 significance of results 158 –162 Valimi Fault 124, 125, 126, 130, 132, 133, 134 Vardar-Imzir-Ankara Zone 170 Variscan Orogeny 33 –34 European Alps 36 magmatism 39 –44 metamorphism 36 –39 sedimentary basins and volcanics 44 – 47 Varvara boudin 145, 146, 154, 158 volcanic arcs, migration characteristics in New Zealand 74 –76 volcanism, Cretaceous of Zealandia Eastern Province geochronology and geochemistry 40 Ar/39Ar 114 description of samples Eweburn Tuff 91 Houhora Complex 91 Motu Tuff 91 Shag Valley Ignimbrite 91 –92 methods 91, 101 U –Pb ICPMS 113 –114 U –Pb TIMS 112 –113 results of geochemistry Eweburn Tuff 101 –102, 105, 106, 107 Houhora Complex 102 –103, 105, 106, 107 Motu Tuff 103, 104, 105, 106, 107 Shag Valley Ignimbrite 101 –102, 104, 105, 106, 107 results of geochronology Eweburn Tuff 96, 97, 98, 99, 100 Houhora Complex 92, 93, 95 –96, 99 Motu Tuff 100 Shag Valley Ignimbrite 96, 99, 100

INDEX results of geochronology and geochemistry discussed age 108 chemical controls 112 extension initiation 109 –111 tectonic controls 111 –112 wedge magmatism 108 –109 Western Province geochronology and geochemistry description of samples Canavans Quartz Monzonite 92 DSDP rhyolite 92 Stitts Tuff 92 Whatarou Granite 92 methods 40 Ar/39Ar 114 U –Pb ICPMS 113 –114 U –Pb TIMS 112 –113 results of geochemistry Canavans Quartz Monzonites 102, 105, 107 DSDP rhyolite 103, 104, 107 Stitts Tuff 103, 104, 105, 106, 107 Whataroa Granite 103, 106, 107 results of geochronology Canavans Quartz Monzonite 93, 95, 96, 100 DSDP rhyolite 93, 95, 96, 99, 100 Stitts Tuff 93, 95, 96, 100 Whataroa Granite 93, 95, 96, 99, 100, 101 results of geochronology and geochemistry discussed age 108 chemical controls 112 extension initiation 109 –111 tectonic controls 111 –112 wedge magmatism 108 –109 Western Finland Arc Complex 227, 229 Whipple Mountains (California, USA) 3 Yenice granitoid 199 Zarouchla Complex 124, 126, 131, 133 Zealandia 89, 90 Zealandia, Eastern Province geochronology and geochemistry of Cretaceous felsic volcanics 40 Ar/39Ar 114 description of samples Eweburn Tuff 91 Houhora Complex 91

Motu Tuff 91 Shag Valley Ignimbrite 91 –92 methods 91, 101 U –Pb ICPMS 113 –114 U –Pb TIMS 112 –113 results of geochemistry Eweburn Tuff 101– 102, 105, 106, 107 Houhora Complex 102 –103, 105, 106, 107 Motu Tuff 103, 104, 105, 106, 107 Shag Valley Ignimbrite 101– 102, 104, 105, 106, 107 results of geochronology Eweburn Tuff 96, 97, 98, 99, 100, 101 Houhora Complex 92, 93, 95 –96, 99 Motu Tuff 100 Shag Valley Ignimbrite 96, 99, 100, 101 results of geochronology and geochemistry discussed age 108 chemical controls 112 extension initiation 109 –111 tectonic controls 111 –112 wedge magmatism 108– 109 Zealandia, Western Province geochronology and geochemistry of Cretaceous felsic volcanics description of samples Canavans Quartz Monzonite 92 DSDP rhyolite 92 Stitts Tuff 92 Whatarou Granite 92 methods 40 Ar/39Ar 114 U –Pb ICPMS 113 –114 U –Pb TIMS 112 –113 results of geochemistry Canavans Quartz Monzonites 102, 105, 107 DSDP rhyolite 103, 104, 107 Stitts Tuff 103, 104, 105, 106, 107 Whataroa Granite 103, 106, 107 results of geochronology Canavans Quartz Monzonite 93, 95, 96, 100 DSDP rhyolite 93, 95, 96, 99, 100 Stitts Tuff 93, 95, 96, 100 Whataroa Granite 93, 95, 96, 99, 100, 101 results of geochronology and geochemistry discussed age 108 chemical controls 112 extension initiation 109 –111 tectonic controls 111 –112 wedge magmatism 108– 109 zircon FTA see fission track analysis