Variscan Tectonics of the North Atlantic Region (Geological Society Special Publication 14)

Variscan Tectonics of the North Atlantic Region

Variscan Tectonics of the North Atlantic Region

edited by D. H. W. Hutton Department of Geology, Trinity College, Dublin, Ireland now

at:

Department of Geological Sciences, The University, Durham, U.K.

D. J. Sanderson Department of Geology, Queen's University, Belfast, U.K.

1984 Published for The Geological Society by Blackwell Scientific Publications Oxford London Edinburgh Boston Melbourne Palo Alto

9 1984 The Geological Society Published by Blackwell Scientific Publications Editorial offices: Osney Mead, Oxford OX2 0EL 8 John Street, London WC1N 2ES 9 Forrest Road, Edinburgh EH1 2QH 52 Beacon Street, Boston Massachusetts 02108, USA 706 Cowper Street, Palo Alto California 94301, USA 99 Barry Street, Carlton Victoria 3053, Australia

DISTRIBUTORS USA and Canada Blackwell Scientific Publications Inc PO Box 50009, Palo Alto California 94303 Australia Blackwell Scientific Book Distributors 31 Advantage Road, Highett Victoria 3190 British Library Cataloguing in Publication Data

Variscan tectonics of the North Atlantic region.--(Special publications of the Geological Society, ISSN 0305-8719) 1. Geology, Stratigraphic--Palaeozoic 9 1984 The Geological Society. Authorization to photocopy items for internal or personal use, or the 2. Geology--North Atlantic region internal or personal use of specific clients, is granted by I. Hutton, D. H . W . II. Sanderson, D. J. The Geological Society for libraries and other users III. Geological Society of London IV. Series registered with the Copyright Clearance Center (CCC) 551.7'4 QE654 Transactional Reporting Service, provided that a base fee of $02.00 per copy is paid directly to CCC, 21 ISBN 0-632-01203-X Congress Street, Salem, MA 01970, U.S.A. 0305-8719/84 $02.00. First published 1984 Set by Preface Ltd, Salisbury, Wilts, and printed in Great Britain at the Alden Press, Oxford

Contents Preface: HUTTON, D: H. W. & SANDERSON, D. J . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Page vii

MAINLAND EUROPE

WEBER, K. Variscan events: early Palaeozoic continental rift metamorphism and late Palaeozoic crustal shortening . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . MEISSNER, R., SPRINGER, M. & FL•H, E. Tectonics of the Variscides in North-Western Germany based on seismic reflection measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . FRANKE, W. Late events in the tectonic history of the Saxothuringian zone . . . . . . . . . . . . BURG, J. P., LEYRELOUP, A., MARCHAND, J. & MATTE, Ph. Inverted metamorphic zonation and large-scale thrusting in the Variscan Belt: an example in the French Massif Central QUENARDEL, J.-M. & ROLIN, P. Palaeozoic evolution of the Plateau d'Aigurande (NW Massif Central, France) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . MAqTHEWS, S. C. Northern margins of the Variscides in the North Atlantic region: comments on the tectonic context of the problem . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

3 23 33 47 63 71

BRITAIN

COWARD, M. e. & SMALLWOOD, S. An interpretation of the Variscan tectonics of SW Britain . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . EEVERIDGE, B. E., HOLDER, M. T. & DAY, G. A. Thrust nappe tectonics in the Devonian of south Cornwall and the western English Channel . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPMAN, T. J., FRY, R. L. &; HEAVEY, P. T. A structural cross-section through SW Devon EDWARDS, J. W. F. Interpretations of seismic and gravity surveys over the eastern part of the Cornubian platform . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . SHACKLETON, R. M. Thin-skinned tectonics, basement control and the Variscan front . . . ARTHURTON, R. S. The Ribblesdale fold belt, N W E n g l a n d - - a D i n a n t i a n - e a r l y Namurian dextral shear zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CRITCHLEY, M. F. Variscan tectonics of the Alston block, northern England . . . . . . . . . . .

89 103 113 119 125 131 139

IRELAND

SANDERSON, D. J. Structural variation across the northern margin of the Variscides in N W Europe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . COOPER, M. A., COLLINS, D., FORD, M., MURPHY, F. X. & TRAYNER, P. M. Structural style, shortening estimates and the thrust front of the Irish Variscides . . . . . . . . . . . . . MAX, M. D. & LEFORT, J. P. Does the Variscan front in Ireland follow a dextral shear zone? COLLER, D. W. Variscan structures in the Upper Palaeozoic rocks of west central Ireland

149 167 177 185

NORTH AMERICA RAST, N. The Alleghenian orogeny in eastern North America . . . . . . . . . . . . . . . . . . . . . . . . LEFORI, J.-P. & HAWORTH, R. T. Geophysical evidence for the extension of the Variscan front on to the Canadian continental margin: geodynamic and palaeogeographic consequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . MOSHER, S. & RAST, N. The deformation and metamorphism of Carboniferous rocks in Maritime Canada and New England . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . WINTSCH, R. P. & LEFORT, J.-P. A clockwise rotation of Variscan strain orientation in SE New England and regional implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . BREWER, J. A. Clues to the deep structure of the European Variscides from crustal seismic profiling in North America . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

197

219 233 245 253 265

Preface The Variscan orogen is a broad zone of deformation which lies between central Europe, the southern British Isles, and fringes the east coast of North America and the north-west coast of North Africa. Deformation occurred in this area in a broad time span encompassing the Middle Devonian to the early Permian. We have used the term 'Variscan' since it seems to be the generally preferred usage throughout Europe. 'Hercynian' has come to be used as a synonym for 'Variscan' especially in the mainland European context whilst 'Alleghenian' describes the late Carboniferous to early Permian events of North America and 'Mauritanian' refers to the African part of the orogen. When compared with other orogens the exact nature of the Variscan and the processes which produced it remain particularly unclear. This arises from a number of major problems some of which we can restate. Although described as a fold belt the orogen is very wide (in excess of 2,000 km) and is not particularly belt-like. There are considerable problems in tracing tectonic-stratigraphic units along the strike Of the orogen for any substantial distance. This is exacerbated by the fact that post-Variscan cover in mainland Europe is extensive and much of the Variscan occurs in separated massifs. Nor is along strike comparison helped by the fact that Variscan research has been published in a variety of languages by scientists who have approached the geology with the methodology appropriate to their own differing traditions. There are clear differences between one end of the belt and the other. Thus in central Europe there are large volumes of granitoids and regionally developed low-pressure high-temperature metamorphics: evidence for high heat flow and thinned crust. Yet in North America the Variscides form a longer, more cylindrical belt of Barrovian metamorphics sitting on large-scale west directed thrusts and containing significantly lower volumes of granitic rocks. When these differences and difficulties are combined with a comparative lack of ophiolitic and blueschist remnants in the orogen we find that a plate tectonic interpretation of the Variscides is less easily sustained than in other orogens. A tradition has tended to develop of Mid-Atlantic 'mobilism' versus central European 'fixism' in debates about the ultimate meaning of the Variscides. It was against this background that we decided to organise a meeting in September 1982, under the umbrella of the Tectonic Studies Group of the Geological Society of London and held at Trinity College Dublin, Ireland. However much a discussion of the tectonics and structures of the whole orogen would have been desirable we decided to concentrate on the northern marginal zone around the North Atlantic. This book, although somewhat broader in outlook than the original theme of the Dublin meeting, examines the geology of the marginal zone between central Europe and Alabama. Why should we look at this area? Firstly, it is easier to define orogenic strike on the margins of the orogen than in the interior, thus allowing along strike correlations and comparisons to be attempted. Secondly, the margins of an orogenic belt are the results of a set of boundary conditions in which the observed structures arise from the interference of two things: (a) the pre-convergence geometry (e.g. foreland configuration) and (b) the gross convergence vector. With luck it might be possible to separate these two important parameters. Thirdly, much recent work and new methodology has been devoted to deformation in

viii

Preface

the marginal parts of orogens. The resurgence of interest in thrust tectonics has generated many new models and ideas from Variscan areas and it seemed appropriate to give this an outlet. Fourthly, geophysical studies and particularly crustal scale seismic reflection experiments are being increasingly undertaken at orogenic margins. Hitherto unsuspected major low angle reflectors (thrusts?) are now appearing in reflection studies with great regularity. Finally there has been a long-standing discussion about the existence and nature of the Variscan front: the putative northern edge of the orogen. A discussion of this problem, in view of the developments noted above, is long overdue. The general structure of the book is simple and, we hope, effective. We have attempted to present information on the marginal zone in separate but adjacent (and overlapping) geographic sectors. We begin in the east, in central Europe and finish in the west, in southern U.S.A. Each sector is introduced by a substantial review article followed by one or more regional, geophysically based papers and then a number of more detailed contributions. Some sections are completed by general and 'ideas' papers. We have broken this basic structure in (a) including the Massif Central, which although not part of the northern margins, is of some importance in tracing the German structure westwards, and (b) by including papers in Britain and Ireland which deal with the style of deformation in the orogenic foreland. The facts, ideas and conclusions reached by the authors are difficult to summarize properly here. However, as editors, with a feeling for the 'middle ground', and a little editorial licence, we would like to point to a few limited, but interesting, general conclusions about the northern margins that were not perhaps too apparent until recently. (a) South and south-east dipping thrusts occur throughout the northern marginal zone of the Variscides: although the intensity and, we suspect, the translation on these thrusts varies from place to place. (b) Although brittle dextral transcurrent deformation is well known as a late stage event in the Variscides, we feel there is evidence of earlier dextral movements. These are ductile in nature and occurred broadly synchronously with the overthrusting, especially in the central part of the belt between Newfoundland and western Europe. This all appears to us to be consistent with the view of a roughly N W - S E directed collision. Thus in the central sector the oblique collision produces variable amounts of overthrusting and strike-slip movement, whereas in North America closure was more normal to the N E - S W trend of the margin and large-scale thrusting predominated. A post-collision continuation of these movements may then have imposed the better known brittle dextral shear system. The nature of the Variscan front, as described and discussed by many authors herein, ultimately reflects all of this. Thus it appears relatively sharl~ in thrust dominated regimes yet more diffuse and difficult to pin-point in areas where transcurrent shear is important. Intensity of thrusting can also be controlled by foreland configuration: intense around promontories, weak in embayments and strike-slip segments. Along strike variations in the position of the front can also depend on gross erosion level within the orogen. On a local scale this may be controlled by whether a thrust tip line is exposed or not. On a more regional level erosion may intersect the upper flat or ramps of crustal scale duplexes, the former giving variable sinuous Outcrop of the thrust front, the latter much straighter outcrop closer to the centre of the orogen. Finally the edge of the deformation

Preface

ix

will tend to be drawn towards the foreland in an irregular manner by small basin developments in the marginal zone. Our final task, and pleasure, as editors is to thank the contributors to this volume and in particular to acknowledge their tolerance over the slow processing of some of the papers. We would also like especially to thank Bob Campbell and Nick Parsons of Blackwells with whom it was a pleasure for us to work. Our final note of gratitude is also one of great sadness. Crosbie Matthews, whose paper appears herein, died in the spring of last year in Uppsala, Sweden. In the early stages of planning the conference and this volume, Crosbie gave freely of his time and advice on Variscan matters particularly in respect of mainland Europe with which he was so familiar. At a later stage he willingly undertook the arduous job of rendering in English the papers of a number of our German and French contributors. We are extremely grateful to him for all this. Crosbie devoted much of his research life to the problems of the enigmatic Variscides. His linguistic ability and personal knowledge of European Variscan geology, geological literature, people and attitudes placed him in a formidable, unique and not always properly appreciated position within our affairs. The paper that appears here is a final and somewhat personalized statement of his views of this rather unique fold belt. D. H . W. H U T T O N D. J. S A N D E R S O N

Variscan events: early Palaeozoic continental rift metamorphism and late Palaeozoic crustal shortening K. Weber SUMMARY: Variscan events are interpreted in terms of a geodynamic process of long duration. It began in the early Palaeozoic, possibly in the late Precambrian in some regions, with widespread rifting of the continental lithosphere. Granulite facies metamorphism, widely in evidence in the Ordovician and Silurian, coincides in time with igneous activity and with continuous accumulation of sediment at the surface. That association is taken to indicate continental rift metamorphism above anomalous regions of the mantle. Folding and metamorphism of what is now regarded as the basement began early in the Devonian. By Upper Devonian at the latest wide areas of crystalline basement had been deeply exposed by erosion. The orogenic crustal shortening which began early in Devonian time induced intensive development of nappe tectonics involving the basement rocks. This resulted in deep-reaching crustal imbrications, especially well shown at the Moldanubian-Saxothuringian zone and Saxothuringian-Rhenohercynian zone boundaries, which evolved to carry crystalline basement rocks towards their foreland regions over distances greater than 100 km. During the course of these nappe developments folding of the adjacent sedimentary troughs proceeded. A geodynamic model of the northern flank of the central European Variscan orogen is presented. (Rb/Sr data are given using ~ R b 87 = 1.42 x 1 0 - 1 1 y-1. The error limits are taken from the original papers, original data are given in brackets.)

An increasingly wide availability of radiometric age-dates in the Variscides (the term should be understood in the sense of European Hercynides, cf. Ziegler 1982) has given new impetus to the discussion of the geodynamic development of the Variscides and of the range of age of the events involved. J/iger's (1977) comprehensive account of radiometric age-dates in central and western Europe led to the conclusion that oceanic sedimentation in late Precambrian and early Palaeozoic time was followed by three phases of orogeny: the Cadomian, the Caledonian and the Variscan. Vidal et al. (1981) have shown that from the evidence of 87Sr/86Sr initial ratios, magmatites in the greater part of the central and west European crust are probably not older than 700 Ma. These ratios rise with time, and this suggests to Vidal et al. (1981) that the Variscides in central and western Europe can be regarded as a closed system with Variscan magmatic rocks derived as melts in relatively young continental crust. There are, of course, older pre-Cambrian rocks within the Variscan crust. This is especially clear in the northern part of the Armorican massif (Cogn6 1974; Cogn6 & Wright 1980; A u t r a n & Cogn6 1980) and in the Bohemian massif (Vejnar 1971; Jakeg et al. 1979). However, in the interpretation of Vidal et al. (1981) based on the trends of the 87Sr/86Sr

initial ratios, these can make up no volumetrically large part of the Variscan crust. Gebauer & Gr/inenfelder (1983), on the other hand, have ffsed U-Pb zircon data to suggest that mafic and ultramafic protoliths older than 1 Ba are detectable in metasedimentary country rocks in the French massif central, in the Moldanubian region and in the central and western Alps. The metamorphism of these mafic and ultramafic protoliths can, nevertheless, be regarded as being everywhere of early Palaeozoic age, and this introduces the question of the meaning of any 'Caledonian' event in the Variscan basement. A Caledonian event in northern Europe is incontrovertibly of orogenic character and is explained there in terms of the closing of the Iapetus Ocean. But in central and western Europe it has not been clearly established that any major crustal shortening was involved in the events which have been called, because of their age, Caledonian.

The 'Caledonian' event within the Variscan orogen Certain peculiarities make it difficult to regard a 'Caledonian' event within the Variscan realm as representing an orogenic event in which significant crustal shortening was achieved: (a) During the time span of 5 0 0 - 4 0 0 Ma

4

K. W e b e r

FIG. 1. Structural map of the European Variscides (after Engel & Franke 1983). Devonian and Carboniferous flysch at outcrop (close stipple) and presumed extent (spaced stipple). Upper Carboniferous parallic molasse, at outcrop (cross-hatched) and presumed extent (hatched). Arrows: tectonic polarity. Black patches: crystalline nappes in the Saxothuringian zone (from west to east: Mfinchberg massif (MM), Wildenfels, Fankenberg, and Gory Sowie (8)) and in Galicia of NW Spain (from north to south: Hesperian massif, Braganca and Morais). SG: Saxonian Granulitgebirge, 1: Flechtingen Hills, 2: Harz Mountains, 3: Rheinisches Schiefergebirge, 4: Ardennes, 5: Odenwald, 6: Spessart, 7: Thfiringer Wald (Thuringian Forest), 8: Gory Sowie (Eulengebirge), 9: Schwarzwald (Black Forest), 10: Vosges, 11: Waldviertel of lower Austria, 12: Elbe line, 13: south Portuguese basin, 14: Alto Alentejo, 15: Sierra Morena. i.e. during the Ordovician and Silurian, enormous quantities of granitic melts of calcalkaline to peralkaline composition were produced. These granites were emplaced in preVariscan crust, and were at a later date deformed to produce orthogneisses. Such orthogneisses produced from pre-orogenic granitoids are widespread in the Variscan basement. Some examples can be cited: Alkaline to peralkaline granitoids in the Alto Alentejo in Portugal have, according to Priem et al. (1970) an intrusion age of 470 Ma. These alkali granites were later converted into Variscan orthogneisses. In the Cordoba-Abrantes shear zone, this deformation produced mylonites and ultramylonites (Kosinowski 1982; Sattler-Kosinowski 1982). In the Hesperian massif, Kuijper (1979) has considered that ages of between 2 and 2.5 Ba on detrital zircons from orthogneisses and paragneisses indicate the existence of distinctly old basement rocks. An age of 1.5 Ba is suggested for sedimentation of the protoliths of the paragneisses (Kuijper 1979). Rb/Sr whole rock ages and likewise U-Pb ages on zircons (lower intercept) from orthogneisses suggest that widespread granitic intrusions took place between 500 and 450 Ma (van Calsteren &

Den Tex 1978; Kuijper 1979; Den Tex 1981, 1982; van der Meer Mohr et al. 1981). The southern part of the Armorican massif and the Massif Central in France have produced evidence of abundant pre-orogenic granites in the range 500-400 Ma. Data from these cases have been summarized by, for example, Dornsiepen (1979) and Autran & Cogng (1980). In the Schwarzwald, where sedimentation began, according to Hofmann & K6hler (1973), later than 900 Ma, an anatexis to which the pre-tectonic granites are attributed and a diatexis which brought gneissification of granites fall in the range of 470-490 Ma. U-Pb age determinations on zircons from the diatexites (Steiger, Bfir & Busch 1972) gave a lower intercept of 489 -+ 26 Ma (473 -+ 26) published by Hofmann & K6hler (1973). According to Hofmann (1979) these data do not provide a sufficient basis for more far-reaching conclusion on the pre-Devonian orogenic history. On the other hand, they do serve to show that there is good evidence that the so-called Rotgneisses of the Saxothuringian zone (see below) were originally emplaced there as granites during the Ordovician and Silurian. Intensive early Palaeozoic granitic magmat-

Variscan events

ism is also known from the Alpine basement. According to von Raumer (1981) all of the Variscan massifs in the western Alps have coarse grained granites, produced during a first anatexis and having intrusion ages between 420 and 450 Ma (Arnold 1970) which were later deformed to become gneisses. U-Pb zircon and monazite ages from granitic gneisses in the Swiss Central Alps (K6ppel, G/inthert & Gr/inenfelder 1981) point to intrusion between 450 and 400 Ma. Further, pre-Variscan granites in the eastern Alps show a maximum in age of intrusion in the range 460-420 Ma (Sch6nlaub & Scharbert 1978; Schmidt 1976a; Heinisch & Schmidt 1976). The timing of magmatic events in the M/inchberg massif of the southern part of the Saxothuringian zone is reasonably well established by the geochronological investigations carried out by S611ner, K6hler & Mfiller-Sohnius (1981) and Gebauer & Gr/inenfelder (1979). According to these studies, sedimentation of the paragneisses of the 'Liegendserie' is no older than 700-1000 Ma. The basaltic protoliths of the M/inchberg eclogites are of Cambrian age. The peak of regional metamorphism that led to the formation of eclogites and kyanite-staurolite gneisses occurred at about 380 Ma. (b) The Ordovician-Silurian magmatism is broadly contemporaneous with a granulite facies metamorphism which was in progress at depth in many regions where continuous stratigraphic successions of that range of age were accumulating at the surface. Two regions must be regarded as providing classic examples of granulite facies metamorphism with concurrent stratigraphic continuity during the Lower Palaeozoic. One is the Granulitgebirge in Saxony and the other is the Hesperian massif. The Granulitgebirge (Figs 1 and 3) lies within the Saxothuringian zone, whose weakly metamorphic sedimentary sequence proceeds in an essentially continuous succession from late Precambrian to Carboniferous. In the Granulitgebirge area itself, the stratigraphic succession ranges from Upper Proterozoic to Devonian. A pre-granulitic migmatisation (anatexis 1), whose relics survived in metatectic structures, was followed at around 450 Ma by granulite facies metamorphism (J~iger & Watznauer 1969; Watznauer 1974; J~iger 1977) at approximately 8 kB and 700-800~ (Behr 1980; Weber & Behr 1983). These data suggest middle pressure granulites. As in the case of the Waldviertel granulites of Lower Austria and the granulites of the Hesperian massif, the Granulitgebirge in Saxony

5

has lensoid eclogites and pyriclasites which were retrograded during the granulite facies metamorphism. The ascent of the granulitic body to higher crustal levels is associated with an amphibolite facies overprint and local anatexis (anatexis II): features which are especially evident in the peripheral parts of the Granulitgebirge area. During the time of the retrograde metamorphism of the granulitic body its stratigraphic envelope therefore experienced a prograde metamorphism. Water release during this prograde metamorphism migrated into the marginal parts of the granulitic body and brought about anatexis II under amphibolite facies conditions (Behr 1980). Behr's (1961) analyses of fabrics have shown that quartz fabrics in the core granulites are dominantly those of small lenticular quartzes having small circle configurations with girdle axes normal to the metamorphic layering. Within the amphibolite facies marginal regime the highly symmetrical small-circle configurations give way to crossed-girdle fabrics whose opening angle decreases with increasing migmatization. Blasto-mylonitic effects at the margins and in the enveloping mica schists show oblique girdle fabrics. Lister & Dornsiepen (1982) interpret these fabrics to mean that the small-circle pattern in the core granulites is typical of strain histories intermediate between axially symmetrical shortening and plane strain, whereas the 90~ crossed girdle patterns at the rims of the granulitic body are typical of plane strain. The oblique girdle fabrics in the surrounding micaschists and gneisses and also the overprinted crossed girdle fabrics with small opening angles reflect a non-coaxial strain path. In the region occupied by the Hesperian massif the sedimentary record is practically continuous from late Precambrian to midDevonian (Kuijper 1979; van der Meer Mohr etal. 1981). Local gaps can be attributed to block faulting. Bimodal volcanism occurs particularly in the Cambrian but is also seen in the Ordovician and Silurian. In the Ordovician, the episode of heightened igneous activity and granulite facies metamorphism corresponds with a widening of the area receiving sediment and a succeeding period of block faulting. The Ordovician granulite facies metamorphism in the Hesperian massif suggest 10-11 kb and 850~ (Kuijper 1979). Just as the Granulitgebirge in Saxony lacks preferred lattice orientation where granoblastic pyriclasites show a granulite facies tempering (Watznauer 1974), so the granoblastic fabric which locally survives as a representative of the granulite facies metamorphism (M1) likewise

6

K. W e b e r

shows no PLO. The granulites exhibit a retrograde course of metamorphism which runs through hornblende granulite facies with local migmatization to the amphibolite facies to greenschist facies. The ascent of the granulites to higher crustal levels can be read in relation to several phases of deformation. Deformation associated with hornblende granulite facies conditions had taken place at a deep crustal level, with temperature between 600 and 750~ and pressure in the range 8-12 kb (Hubregste 1973; Maaskant 1970). According to van Calsteren et al. (1979) the migmatization process came to an end at 347 _+ 17 Ma. The main difference between granulite facies and hornblende granulite facies metamorphism lies in an increase of P fluid during the M 2 metamorphism (Engels 1972; Kuijper 1979). Amphibolite facies metamorphism was accompanied by a penetrative deformation (F4) responsible for the subvertical, N W - S E trending main foliation, preferred lattice orientation of hornblende and blastomylonitic textures (Kuijper 1979). In the course of the further rise of the granulites the deformation produced cold worked fabrics. Deformation history and metamorphic succession are therefore closely comparable with those in the Granulitgebirge of Saxony. The Moldanubian granulites in Lower Austria again reveal retrograde metamorphism during the course of tectonic deformation. Granulite facies metamorphism took place, according to Scharbert (1977a), at approximately 11 kb and 760~ These are predominantly light-coloured, quartz-rich granulites, with insertions of subsidiary amounts of garnet pyroxenites which bear a granulite facies overprint. These latter rocks, according to Scharbert (1977a), may in their original condition have been upper mantle material which moved into the lower crust where they acquired their granulitic character. The granulite facies rocks of Lower Austria and Czechoslovakia occur in tectonic nappes which rest on rocks of lower metamorphic grades (Fuchs 1971, 1983; Thiele 1976a,b; Tollmann 1982). The retrograded marginal parts of the granulite bodies show ribbon quartzes like those in the Granulitgebirge. Their preferred lattice orientations indicate non-coaxial deformation. The primary granulitic fabrics, in contrast, with their discoidal quartzes, suggest coaxial deformation. The age of the granulite facies metamorphism is 446 + 35 Ma (431 -+ 35) according to Arnold & Scharbert (1973). An Sr homogenization at 485 + 11 Ma (469 -+ 11) is regarded as indicating the age of the educts of

the granulites. It appears difficult, however, to interpret this educt age as an age of sedimentation, because the Gf6hler orthogneiss, which is associated with the granulites has given an educt age of 474 +_ 23 Ma by Arnold (in Scharbert 1977b). A likely suggestion would be that the Sr homogenization should be related to Ordovician rift metamorphism (see below) with the 446 Ma (431) date indicating the granulite facies dewatering of the rocks. The leptyno-amphibolitic group of the Massif Central contains acidic and mafic granulites. According to Burg (1977) and Burg & Matte (1978) (and see Burg et al., this volume) the granulite facies metamorphism is older than the main deformation (F 1 and F2) in the Massif Central. Some granulite bodies contain ghosts of isoclinal folds older than the static recrystallization under granulite facies conditions which, according to Dufour, Piboule & Duthou (1983), took place at 7-8 kb and 800-825~ Granites of crustal derivation with Rb/Sr whole rock ages between 450 and 550 Ma, which were intruded pre-tectonically, were later transformed into orthogneisses under amphibolite facies conditions. In the area of the Monts du Lyonnais this amphibolite facies metamorphism took place at 5-6 kb and 700-725~ (Dufour et al. 1983). It has overprinted the older granulite facies rocks.

FIG. 2. Diagrammatic sketches of the early Palaeozoic continental rifting and associated rift metamorphism (a) and the development of an injective granulite fold during later orogenic crustal shortening (b). The present SE-vergence of the Sfichsische Granulitgebirge (which is not shown in this sketch) results from the younger back-folding. (Further explanations in the text.)

Variscan events

Continental rift metamorphism during Ordovician-Silurian time It is in fact not a simple matter to explain the coincidence of granulite facies metamorphism deep in the crust, intensive pre-tectonic igneous activity and development of a continuous sedimentary sequence at the surface. It is made more difficult if one assumes that the process of producing granulite facies metamorphism is necessarily bound up with orogenic crustal shortening. The fact that many of the early Palaeozoic granitoids predate deformation and also the widespread evidence of more or less uninterrupted Lower Palaeozoic stratigraphic successions in many parts of Variscan Europe would tend to discredit any such assumption. The following proposalS on continental rift metamorphism during Ordovician-Silurian time are based mainly on the model developed by Den Tex, van Calsteren and their coworkers for the Hesperian massif. The basic idea is that a continental rift develops on top of an anomalous mantle and the heat transferred into the lower crust produces granulite facies metamorphism (Fig. 2a). From work on recent passive continental margins we learn that rifting processes in continental crust promote ductile stretching in the lower crust, which contrasts with the brittle manner of reaction in the upper 10-15 km of the crust where graben formation proceeds (De Charpal et al. 1978; Montadert et al. 1979; Le Pichon & Sibuet 1981; Le Pichon, Angelier & Sibuet 1982). This concept, introduced by McKenzie (1978a,b) has been widely accepted (Christie & Slater 1980; Royden, Sclater & yon Herzen 1980) and might provide a means of interpreting the fact that 'Caledonian' granulite facies metamorphism was contemporaneous with sedimentation and pre-orogenic igneous activity. The continental crust is underlain by lithospheric mantle. Rifting of the continental crust must have some association with a rifting in the lithospheric mantle which promotes ascent of hot asthenospheric mantle material. Such a convective supply of heat may be regarded as a course of heightened temperature at the crust-mantle transition. Partial melts of tholeiitic composition may be transformed to eclogites in the higher parts of the upper mantle or at the crust-mantle boundary (van Calsteren & Den Tex 1978) or else they invade the lower crust and become metagabbros or amphibolites. The P T conditions for granulite facies metamorphism (7-11 kb and 700-850~ suggest that granulites could be produced at the base of

7

an approximately 30-40 km thick continental crust. However, the granulite metamorphism in almost all of the Ordovician granulites follows on a pre-granulite migmatization which corresponds to anatexis I. This is an understandable relationship, for the granulite facies dewatering of deep crustal rocks is a gradual process. The expulsion of water coincides with an introduction of mantle CO 2. Lead isotope ratios in K-feldspars from several metamorphic and granitic rocks in the southern Schwarzwald, which suggest a very early formation of the basement, have been reinterpreted by Kober & Lippolt (1983) to be the result of crust-mantle interaction during anatexis I: mantle lead was injected upwards out of a degassing mantle region during genesis of 'Caledonian' magma. Increasing temperature, and expulsion of water from the granulites led to the formation of calcalkaline granite, granite magmas which invaded the higher crust and which were later deformed, during crustal shortening, to produce orthogneisses such as the Rotgneise in the Saxothuringian zone. The fact that the Ordovician-Silurian granitoids pre-date Variscan deformation, and that the primary quartz fabrics in the granulites are highly symmetrical, leads to an interpretation in terms of metamorphism under stretching conditions, with a dominantly coaxial deformation path.

Crustal shortening and the formation of nappes with rocks in granulite facies In order to understand the origin of tectonic nappes which involve granulitic facies rocks it is necessary to consider the rheological characteristics of such rocks and to be aware of the distinct rheological character of water-rich amphibolite facies rocks. Quartz-rich crustal rocks, in which quartz is the stress-supporting mineral, are drastically affected by hydrolytic weakening and grain boundary migration when the recrystallization temperature of quartz is exceeded. For 500~ and a strain rate of 10 -14 s -I steady state creep stress of the order of 100 b can be assumed. The steady state creep stress falls to about 1 0 b for 600~ and to approximately 1 b for 800~ (Mercier, Anderson & Carter 1977). A different set of considerations applies in rocks free of water. In quartz and feldspar as the stress-supporting minerals, steady state creep stresses at 500-600~ can be expected to be of the order of 1-2 kb (Heard 1976). High

8

K. Weber

steady state creep must be assumed to apply in basic granulites and in eclogites also. Early granulites, later involved in crustal shortening, will therefore, even at high temperatures, behave mechanically in a much 'stiffer' fashion than would 'wet' rocks. Consequently, granulites, in the course of crustal shortening, will promote the formation of large-scale fold structures. Since the granulites are overlain by water-rich, amphibolite facies, migmatitic rocks, they may penetrate the superjacent parts of the crust in the form of 'injective' folds. The ascent of the granulites would here be ascribed to amplification of large-scale folds (Fig. 2b) rather than to a diapiric mechanism of the kind proposed by Lehmann (1984), Watznauer (1974) and Behr (1961, 1980). The granulites take on a retrograde effect where they are in relatively close association with their envelopes, whereas these surrounding rocks, as in the case of the Granulitgebirge in Saxony, show a prograde metamorphism due to the ascending hot granulites, which can lead to the development of a zone of contact metamorphism (Behr 1961). Anatexis II relates to this orogenic crustal shortening. During the amplification process the peripheral parts of the granulitic mass and the country rocks around take on a blastomylonitic deformative effect (again, the Granulitgebirge provides examples). The oblique girdle fabrics and the overprinted cross girdle fabrics reflect the non-coaxial strain path taken during the course of amplification of the large-scale granulitic fold. The Granulitgebirge in Saxony shows a southeastwards tectonic vergence which is due to a later, SE directed tectonic overprint recognizable in other parts of the Saxothuringian zone (Weber & Behr 1983; Franke, this volume). That local redeformation apart, the Granulitgebirge did not develop tectonically beyond the stage of the 'injective' folding. Elsewhere, e.g. in the Moldanubian of the Bohemian massif, in the Hesperian massif, in the catazonal complexes of Braganca-Moreis and the French Massif Central, a more effective crustal shortening has produced extensive nappe complexes involving granulite facies rocks. These, in all of the cases mentioned, now rest on much more weakly metamorphic rocks. As in the basement complexes of the Saxothuringian zone, it is possible to recognize pre-Middle Devonian folding and metamorphism in the basement of the central zone of the Variscides. In the southern part of the Vosges (Maass & Stoppel 1982) and in the Schwarzwald (Maass 1981) non-metamorphic

Upper Devonian rests on crystalline basement. The folding and metamorphism must be older than 375 Ma (360 Ma), because all of the lateto post-tectonic granites are younger than that (Brewer & Lippolt 1974). Deep-reaching erosion of the basement during this time is suggested also by the fact that the post-tectonic granites, according to Emmermann (1976), were emplaced at high crustal level, within rocks which must have been at greater depth during the preceding anatexis. Indications of relatively early folding and metamorphism are available in an Sr-homogenization earlier than 370 Ma (358 Ma) in the phyllites of the northern Vosges (Steige and Vill~ schists) which Clauer & Bonhomme (1970) interpreted as an age of metamorphism. Rb/Sr isochron ages on diatectic and anatectic granites in the Massif Central give 375 Ma in the Limousin (Duthou 1978), in the Vend6e 385 Ma, and 3 7 5 M a at Morbihan (Vidal 1976), suggest a minimum age of the metamorphism (Autran & Cogn6 1980). K/Ar ages on metamorphic hornblendes and Rb/Sr ages on muscovites, all in the range 360-350 Ma (Autran & Cogn~ 1980), date the cooling of the basement to 500-400~ Intrusion of the posttectonic granites begins, as in the Schwarzwald, in the higher part of the Upper Devonian. Stratigraphic evidence, too, suggests a preUpper Devonian age for the basement. North of Lyons the most highly metamorphic rocks in the Massif Central are overlain by epizonal volcano-clastic Upper Devonian sediments, which are in turn overlain by non-metamorphic Vis~an (Burg & Matte 1978). This series of examples could be extended to include other Variscan basement complexes. In what follows, the single further case of the mid-German crystalline rise will be examined. Here, too, it is possible to recognize a postSilurian, pre-Upper Devonian folding and metamorphism. The mid-German crystalline rise (Scholtz 1930; Brinkmann 1948) forms the northern part of the Saxothuringian zone (Fig. 3). It can be followed from the Saar district in the west to at least the Elbe line in the east. Metamorphic and magmatic rocks of the mid-German crystalline rise are exposed in the Odenwald, Spessart, Kyffhfiuser and Ruhla crystalline complexes. Radiometric and stratigraphic evidence point to the existence of an orogenic event during Lower Devonian time (Fig. 3). Stratigraphic evidence of pre-Middle Devonian deformation and metamorphism is available in the Saar region, where nonmetamorphic sediments of Middle Devonian

Variscan events

9

F1G. 3. Structural map of the Rhenohercynian and Saxothuringian zones. age, encountered in the Saar 1 borehole, rest on chloritized albite granite of Lower Devonian age (394 -+ 24 Ma; Lenz & Mfiller 1976). The metasediments of the Spessart possibly represent a time span from late Precambrian in the south to at least Ordovician in the north (Matthes 1954; Bederke 1957; Okrusch, Streit & Weinelt 1967; Matthes & Okrusch 1977). In the Odenwald (B611steiner Odenwald) and in the Spessart pretectonic granites were emplaced 398-419 Ma ago (Kreuzer etal. 1973; Lippolt, Barany & Raczek 1976). The granites were transformed during later metamorphism into muscovite-biotite gneisses. These gneisses are regarded as equivalents of the widespread Saxothuringian Rotgneis (Scheumann 1932, 1939; Bederke 1957; Matthes & Okrusch 1965, 1977). In the central part of the crystalline Spessart this metamorphism took place at 5-6 kB and 600-650~ (Matthes & Okrusch 1977). In the western part of the Odenwald (Bergstr~Ber Odenwald) the conditions of regional metamorphism were 4-6 kb and 650-670~ (Okrusch et al. 1975). The exact age of this regional metamorphism is not known. But it can be roughly bracketed by the following considerations. The Bergstrfiger Odenwald is dominated by a sequence of plutonic rocks ranging from older gab-

bros to younger diorites and granodiorites. The intrusion of this igneous sequence post-dates the regional metamorphism (Maggetti 1975). Table 1 summarizes the available radiometric data from the Odenwald and Spessart. The metamorphism must be younger than the 'redgneiss' intrusions and older than the hornblende cooling ages. Therefore, the discordia intercept age of 380 Ma determined by Todt (1979) on zircons of the grain-size fraction <2 m from the B611steiner Odenwald granodioritic gneiss and its metasedimentary gneiss cover could correspond approximately to the time of regional metamorphism. In that case, the regional metamorphism would be of the same age as in the Mtinchberg Massif. After the time of deformation and regional metamorphism, or more probably starting at some point during this time, fast uplift of the mid-German crystalline rise must have taken place. In the area of the Saar 1 borehole a Lower Devonian crystalline basement is covered by non-metamorphic Middle Devonian sediments (Zimmerle 1976). In the Odenwald 10-12 km of rocks must have been eroded prior to the emplacement of the igneous intrusions since the depth of intrusions was about 3.5-5 km and this magmatic event post-dates the regional metamorphism (Maggetti 1974, 1975) that took place at a depth of about

10

K. W e b e r

TABLE 1. Radiometric data from the Odenwald and Spessart (Rb/Sr ages are recalculated using ;t = 1.42 • 10-Nyr -1) 1 'Red-gneiss' intrusion

398-419

2 Granodioritic gneiss and paragneiss cover of the B611stein Odenwald

380

3 Metamorphism at 4-6 kB and 600-650~ (regional metamorphism) 4 Cooling of the metamorphic rocks to about 400-500~ 5 Intrusion of gabbros, diorites and granodiorites. Depth of intrusion 3.5-5 km (1-1.5 kB) 6 Cooling of the magmatites to about 400-500~ 7 Cooling ofthe Odenwald to about 300~

Approx. 370 Post-regional metamorphism but syntectonic 340 (northern part) 335 (southern part) 330-325

15 km. The K / A r cooling ages of the magmatic rocks cannot be far from the intrusion age, because the closing temperature of hornblende is about 400-500~ (Kreuzer & Harre 1975), and the temperature of the country rocks has been estimated by Maggetti (1974, 1975) at about 200~ The cooling of the igneous rocks as indicated by the K / A r cooling ages of hornblende and biotite (Kreuzer & Harre 1975; Hellmann, Lippolt & Todt 1982) proceeded from 400-500~ at about 342-335 Ma, to 200-300~ at about 330-325 Ma. It should be mentioned that the cooling ages of hornblende and biotite in the northern Odenwald seem to be somewhat older than those of the southern Odenwald. During the time of uplift, deformation continued in the Odenwald and Spessart. This is indicated by the flaser-gneiss-fabric of some of the biotite-diorites and granodiorites. This fabric is interpreted by Nickel & Maggetti (1974) and Maggetti & Nickel (1976) as the result of syntectonic intrusion. It is furthermore indicated by the widespread syntectonic retrograde metamorphism and phyllonitization (Murawski 1958) under greenschist facies conditions (Matthes 1954; Matthes & Okrusch 1977; Gabert 1957). In the quartzite-micaschist series of the northern Spessart this diaphthoresis is related to the south-facing late tectonic overthrust (Michelbach overthrust) of the north-western amphibolite paragneiss (Bederke 1957; Gabert 1957; Plessmann 1957; Murawski 1958).

Rb/Sr whole rocks (Kreuzer et al. 1973; Lippolt et al. 1976) Spessart and B611stein Odenwald U/Pb zircon, grain-size fraction < 2 #m (Todt 1979) age of intrusion or metamorphism? Spessart 5-6 kB/600-650~ (Matthes & Okrusch 1977), Odenwald 4-6 kB/650670~ (Okrusch et al. 1975) K/Ar hornblende (Kreuzer & Harre 1975) Odenwald Maggetti (1974, 1975) Odenwald K/Ar Hornblende (Kreuzer & Harre 1975; Hellmann et al. 1982) Odenwald K/Ar biotite (Kreuzer & Harre 1975; Hellmann eta'. 1982) Odenwald

The Carboniferous event The major phase of folding and metamorphism occurred early in Devonian time in the Variscan basement, but the weakly metamorphic external zones of the Variscides were first folded during the course of the Carboniferous. This latter episode, however, did bring significant further shortening in the central zones of the Variscides, documented in folding and imbrication of weakly metamorphic Devonian and Carboniferous sequences which are discordant on crystalline basement. The style of deformation was now different. The uplift due to orogenic shortening and the consequent cooling had set the stage for inhomogeneous deformation of the crust with, at macroscopic scale, dominantly brittle behaviour. At mesoscopic and microscopic scales this deformation shows itself as blastomylonites and at still lower temperatures and/or higher strain rates by development of ultramylonites. Such mylonite zones are normally interpreted as shear zones. They arise as the reactions of a cooling and therefore more rigid crust to orogenic foreshortening. Such deep-reaching shear zones, associated with uprise of crystalline basement, are found at, for example, the northern margin of the midGerman crystalline rise, and at the boundary between the Saxothuringian and Moldanubian zones along the Erbendorf line, and in the northern Schwarzwald and the northern part of the Vosges. A further example is the Badenweiler-Lenzkirch zone in the southern Schwarzwald. Deep-reaching crustal imbrication is especially well displayed at the southern

Variscan events

FIG. 4. Diagrammatic sketch of the development of the Giessen greywacke nappe. limit of the Ossa Morena zone, where tectonic transport is toward the south Portuguese basin--the evidence is well seen in the Arancena area in southern Spain. In eastern Bavaria a N W - S E belt of blastomylonitic gneisses (Perlgneise) is exposed. The mylonitic overprint of pre-existing para- and orthogneisses which began under P T conditions of 550~ and 2.8 kB (Blfimel, 1983, personal communication), ended in the Perlgneise with the paragenesis quartz + white mica + chlorite + albite in low temperature ultramylonites (Bltimel 1977). Rb/Sr thin slab measurements on these blastomylonites carried out by K6hler, Christinas & M/iller-Sohnius (1983) suggest an intrusion age of 474 _+ 18 Ma for the protoliths of the orthogneisses and a 346 _+ 29 Ma age for the blastomylonitic overprint with the production of the Perlgneise. Further uprise along this shear zone in the course of the Carboniferous can be estimated to have continued until at least 310 Ma on the basis of biotite cooling ages. The progressive development of these large shear zones provides a basis for suggesting more or less continuous crustal shortening during the Devonian and Carboniferous. This is,

11

however, a very broad statement. The widespread volcanism, already evident in the Upper Devonian and continuing into the Permian, would suggest the intervention of phases of crustal stretching. Ascent of the early Variscan basement can also be documented in the greywacke influxes into the adjacent troughs. The example of the Rheinisches Schiefergebirge shows that erosion of the mid-German crystalline rise and receipt of turbidites in a basin on its north began at least as early as low Upper Devonian. These Upper Devonian greywackes are encountered today as allochthonous units in the southern Harz, the southern part of the Hesse depression ('Hessische Senke') and the southern Rheinische Schiefergebirge (Figs 3 & 4). They were at one time parts of a nappe complex continuously developed from the southern Schiefergebirge to the southern Harz (Engel et al. 1983; Weber 1978, 1981; Weber & Behr 1983). The original site of deposition of these greywackes is not now seen--it was overridden by the mid-German crystalline rise. From this one can estimate that the midGerman crystalline rise must have travelled at least 100 km toward its foreland and one notes that the greywacke nappes, for their part, have suffered deformation which has gone as far as producing recumbent folds. The boundary between the mid-German crystalline rise and the Rhenohercynian zone is one of the most significant boundaries in the Variscides. Its geodynamic character as a zone of deep-reaching crustal imbrication (a subfluence zone) was recognized by Weber (1978, 1981). The transition from the crystalline rocks of the mid-German crystalline rise into the very low grade metamorphic state of the rocks in the Rhenohercynian zone is represented in a zone of phyUitic rocks which can be traced from Dtippenweiler in the northern part of the Saar north-eastward to the Wippra zone in the southern Harz Mountains (Figs 3 & 5). Deep boreholes show that this northern phyllite zone continues farther to the NE. In the Wippra zone Ordovician to Upper Devonian sedimentary rocks are known. In the southern part of the Hessische Senke and in the southern Taunus bimodal metavolcanites are widespread. The more acid variants are in some cases ignimbrites. The presence of Mg-riebeckite indicates alkali volcanism. The age of the educts of these metavolcanites is not known. It may, however, be suggested that these bimodal volcanites possibly belong among the widely developed Ordovician rift-related volcanics. It is the case that in the remainder of the

12

K. Weber

FIG. 5. Diagrammatic cross-section through the Rhenohercynian crust in the area of the Rheinisches Schiefergebirge. The lower cross-section shows the position of horizons of high magentotelluric conductivity (conductivity data after Giese et al. 1983). Rhenohercynian zone (with the exception of parts of the Stavelot-Venn m a s s i f - - K r a m m 1982; Schreyer & Abraham 1978) metamorphic temperatures did not exceed 350~ and in most places were in the range 200-300~ But 4 0 0 - 4 5 0 ~ was reached in the northern phyllite zone (Meisl 1970). Taking account of the structural evolution of the northern phyllite zone and of the likely original configuration of sedimentary thicknesses in the southern part of the Rheinische Schiefergebirge, Weber (1978, 1981) has proposed that the rocks of the northern phyllite zone were at one time at considerably greater depth. This suggestion has been confirmed by work done by Massonne & Schreyer (1983). Using a new phengite barometer Massonne & Schreyer

have come to the conclusion that the metamorphic conditions were 10-12 kb at 400-450~ These findings (they correspond to a geothermal gradient of approximately 10~ km -~) would resemble what is proposed for B-subduction zones. But since the ophiolites and andesites typical of B-subduction are lacking, as is also any indication of a paired metamorphic belt, it is reasonable to await further information before proceeding to any relatively ambitious interpretation of these estimates of pressure. Typical of the whole orogenic character of the Rhenohercynian zone is a NW directed polarity (Figs 5 & 6). This is evident not only in the strong NW vergence of the folding and thrust structures (Weber 1978, 1981) it is also

13

Variscan events

oo

~,..~

0 .=

0

14

K. W e b e r

developed in the flysch sediments which were accommodated in a succession of basins which migrated from south to north in late Devonian and Lower Carboniferous time, but also in the migration of folding and metamorphism from south to north. K/Ar datings on white micas from rocks which were originally vitric tufts in the Rheinische Schiefergebirge, and which were syntectonically recrystallized at temperatures below 350~ indicate that the prograde metamorphism is of an age around 330 Ma in the southern part of the Rheinische Schiefergebirge and 300 Ma in the northern part (Ahrendt, Hunziker & Weber 1978; Ahrendt et al. 1983). (Fig. 3). The beginning of prograde metamo/'phism in the southern part of the Rheinische Schiefergebirge is therefore contemporaneous with the cooling of the igneous rocks in the Odenwald to temperatures around 300~ (Table 1).

Geodynamic interpretation The preceding discussion makes it clear that the geodynamic development of the European Variscides cannot be treated as an exclusively Carboniferous set of events. The whole scheme of evolution began with continental rifting in the early Palaeozoic which had an association with metamorphism and anatexis in the lower crust. One cannot exclude the possibility that the process of continental rifting may locally have led to limited developments of oceanic crust. Orogenic crustal shortening began with the Acadian event in the early Devonian. B-subduction during this episode seems likely only in the neighbourhood of the Ligerian suture. How and indeed whether such a suture continued towards what later became the Alpine region is not known. This being the case, considerable uncertainty surrounds any suggestion that the Acadian event is the result of a widely developed B-subduction operating on the southern side of Variscides. Interpretation of the Rhenohercynian depositional basin as a Devonian and Lower Carboniferous back-arc basin (Ziegler 1982; Floyd 1982) is tempting. But one would encounter serious difficulty in proceeding from there to geodynamic explanation of the Saxothuringian depositional site, with its quasi-continuous late Precambrian to Lower Carboniferous sequence. Further, there is the problem that the northward directed nappe and thrust tectonics (Weber 1978, 1981; Weber & Behr 1983; Enge[ et al. 1983; Behr 1978; Behr, Engel & Franke 1980, 1982; Meissner, Bartelsen & Murawski 1981; Giese

et al. 1983; Franke, this volume) are no less impressive in the Variscides than those nappes directed southward (Burg & Matte 1978; Burg et al. 1984; Tollmann 1982; Engel, Feist & Franke 1978, 1981). Compared to the Alps or the Himalayas, for example, the European Variscides show a clearly expressed bilateral symmetry. This bilateral symmetry invites the suggestion that similar geodynamic processes were producing crustal shortening on both flanks of the orogen. It is, however, evident that a proposal of classical B-subduction applies altogether less readily on the north side than on the south. The geodynamic model presented in what follows is based on the evidence of the north flank of the European Variscides, especially the Rhenohercynian zone and is based on the view that this was an essentially ensialic tectonic process. It takes account of the well determined polarity of structural and metamorphic developments. Geometric and kinematic analyses of the exposed suprastructures show (Weber 1978, 1981) that the frequent NW verging overthrusts in the Rheinische Schiefergebirge are listric overthrusts. These overthrusts arise at different depths out of horizontal thrust planes. Major structures can thus be related to the horizontal layering in the infrastructure, which has been interpreted by Weber (1978, 1981) as resulting from a horizontal flow regime. The horizontal thrust planes which pass upward into listric overthrusts tend to produce an uncoupling of the infrastructure from the folded and imbricate suprastructure (Fig. 5). Based on homologous temperatures for wet granite solidus, Vetter & Meissner (1979) have divided the continental crust into two layers: an upper layer with T / T m < 0.6-0.65 showing brittle response to stresses and where steady state creep does not seem to be possible in times < 10~Ma, and a lower layer with T I T m > 0.6-0.65 showing transient creep in its upper part and steady state creep in its lower part. The transition from brittle to creepdominated behaviour lies, in areas of higher heat flow, at a depth of 10-20 km, and in older continental shields at a depth of 40-60 kin. In view of the high heat flow in the Variscan orogen this proposed division is in good agreement with the structural interpretations and the seismic and rheological data mentioned above. Behr (1978), Weber (1978, 1981) and Behr & Weber (1980) have applied the term 'subfluence' (in the sense in which it was used by Ampferer 1906; see also Schmidt 1976b) to the progress of underflowing mass transfer of mat-

15

Variscan events

erial suggested to have been active in the lower crustduring the evolution of an orogen. Weber (1981) has proposed that the cause of the subfluence is a relative movement of continental crust with respect to lithospheric mantle. This assumption is based on the fact that in the neighbourhood of the Moho there is a critical lithological change. Predominantly peridotitic composition is suggested for the lithospheric mantle (Ringwood 1975). Olivine and pyroxene can be regarded as the stress supporting minerals in such rocks. Available theological data on olivine as well as on rocks of peridotitic composition indicate very high steady state creep stresses, extending up to several kilobars, in the Moho range of temperature of 500-600~ (Meissner & Strehlau 1982; Heard 1976; Stocker & Ashby 1973; Ashby & Verrall 1977; Mercier 1980; Goetze 1978; Post 1977; Kirby 1980; Carter 1976; Nicolas & Poirier 1976). It is then to be anticipated that there is a well-developed, geodynamically effective rheological boundary zone between the 'dry' peridotitic upper mantle and the 'wet' quartz-feldspar-rich continental crust. In some areas of former continental rift zones where an intervening layer of higher stiffness can be developed at the base of the continental crust some perturbations can occur at the crust-mantle boundary which give rise to the formation of granulite nappe complexes as described above. A relative movement of the lithospheric mantle in a sense opposite to the one indicated at high levels by the tectonic structures would provide a plausible explanation of the tectonic movement picture. The fact that the tectonic structures on the southern side of the European Variscides have a southward vergence and

those on the north a northward vergence could then be seen to be consistent with the assumption of convergent movement in the lithospheric mantle. For mechanical and geometrical reasons, such a convergence is possible only if subduction of lithospheric mantle is involved. The causes of such a subduction are unknown. Since the lithospheric mantle is, by reason of its higher density as compared to the asthenosphere below, in an unstable equilibrium, it could be suggested that a potential instability established during an earlier rifting stage might during later convergent plate movement lead to downward detachment of lithospheric mantle ('A-subduction', Weber 1981). Descent of large volumes of lithospheric mantle into the asthenosphere would require, in exchange, upward transfer of large volumes of asthenospheric material (Fig. 7). Since the Iapetus Ocean, to the north of the Variscides, had closed, since the opening of the Atlantic had not yet begun, and since the presence of the Gondwana continent in the south (a possible ocean was closed, at least by Upper Devonian time) excluded availability there of any mid-ocean ridge at which massequilibration could be effected, any subducting segment of lithospheric mantle has to retreat in order to allow mass-equilibration. A possible means of return is available in Andrews & Sleep's (1974) model of forced convection, below a crustal segment, induced by the subducting slab of lithospheric mantle (Fig. 7). This Andrews-Sleep cell would shift, during the retreat of the descending lithospheric mantle, outwards toward the margin of the developing orogen. In the case of the bilateral Variscides we have to assume two such Andrews-Sleep

,j

NW

o~

F 1 - folding

0

~

F 2 - backfotding

co~~

SE

subsequent rifting

1O0km

FIG. 7. Diagrammatic sketch of the geodynamic development of the northern branch of the mid-European Variscides. (Further explanation in the text.)

16

K. W e b e r

cells operating on either side of the orogen and in opposite senses. (In the sketch in Fig. 7, only the northern A-subduction is shown.) An assumption basic to this model is that the orogenic shortening took place in a wide region of crustal convergence. Considerations arising from the history of development of the orogen require that any explanation should also take into account the fact that earlier processes of continental rifting have introduced mechanical inhomogeneities which led to local and regional perturbations in the subsequent orogenic evolution. A further assumption is that the regional stresses were first resolved in what later became the central region of the orogen. The siting of that particular region prescribed the zone in which lithospheric mantle became detached from the crust and so initiated the subduction processes. In the central European Variscides the zone thus prescribed is the Moldanubian zone. The central zone is also the region in which uprise of hot asthenospheric material began and initiated thermal weakening of the crust. Away from the central zone, in places where continental crust and lithospheric mantle remained in contact with one another, there must nevertheless have been relative movement of these two. The frictional stresses induced could have brought about a deformation of the lower crust. The frictional stresses need not, however, have been so great that the entire crust was overtaken by this deformation. One could in this way explain the Lower Devonian age (392 + 10Ma, Schoell, Leuz & Harre 1973) of deformation and metamorphism of the Ecker Gneiss (Fig. 2), which later, by means that are not yet understood, became involved in the late Carboniferous folding and metamorphism that are imposed on the sedimentary rocks of the Harz Mountains. As the two subduction fronts withdraw from one another a new episode begins in the development of the two flanks of the orogen. Interaction between the descending slabs of lithospheric mantle and the opposed sense of rotation in the Andrews-Sleep cells produces a compressive stress field between the slab and its associated Andrews-Sleep cell, the effect of which spread into the crust (Fig. 7). This second deformation must, however, because of the sense of rotation of the Andrews-Sleep cell, be antivergent to the first deformation. This is a possible explanation of the south-eastward vergence of the isoclinal, synmetamorphic F 2 folds encountered in the Saxothuringian zone (Weber & Behr 1983;

Franke 1984) and the southernmost part of the Rhenohercynian zone. Also the SE-facing structures in the non-metamorphic sediments of the Saarbriicken anticline and the postmetamorphic SE-directed, suprastructural overthrusts of Dfippenweiler and the northern Spessart (Michelbach overthrust) can be attributed to this secondary stress field. The presence of an antivergent homoaxial F 2 folding is a general phenomenon in the more internal parts of the Variscan orogen which is not only seen in Europe. The mechanism proposed above could provide a possibility of understanding this phenomenon. The Ordovician granulites seem to have been formed under a geothermal gradient of about 20-30~ -1 (Zwart & Dornsiepen 1978; Behr et al. 1980). During Acadian time (Lower to Middle Devonian) the gradient increased to 30-40~ km -1, and in the Carboniferous and Permian perhaps locally reached even 80-100~ km -1 (Zwart 1976; Zwart & Dornsiepen 1978; Behr etal. 1980; Weber 1978; Buntebarth 1982; Buntebarth, Koppe & Teichmfiller 1982). The steepest increase is to be observed during the Upper Carboniferous and Lower Permian. Thus, in the Variscan crust, there was a pronounced increase of the geothermal gradient from the early Palaeozoic (largely Ordovician) rifting stage to the final stage of the Variscan orogeny. A cause of the steepening of geothermal gradients in the course of the Variscan progression of events may lie in the buffering of endothermic prograde metamorphic reactions. The rate of metamorphism is determined by the net input of heat into the metamorphic pile, the enthalpy of metamorphic reactions and the heat capacity of the rock-forming minerals. The net input of heat can be understood as the sum of heat which enters the system and which is generated inside the metamorphic pile minus the heat which leaves by advection and conduction. The suggestions would be that this buffer effect was especially effective in, for example, the Saxothuringian zone with its thick pile of sediments reaching back in time to late Precambrian. Steeper geothermal gradients associated with the rift metamorphism would in such circumstances be expected to exist in the deeper crust only. The whole thickness of crust, on an overall average, would suggest only a low geothermal gradient. Acadian metamorphism shows widespread occurrences of kyanitebearing middle pressure metamorphic assemblages. They reflect the temperature increase at mid-levels of the crust. Later, when nappe

17

Variscan events

development became more intense they proceeded into higher levels of the crust. There, such rocks took on a high temperature/low pressure overprint. The widespread surface near very weakly metamorphosed rocks, which were never deeply buried and which first encountered deformation and metamorphism at a later stage of orogeny, also became exposed to Abukumatype metamorphic conditions. Here it is necessary to take the view that in addition to possibly enhanced radioactive heat production and synorogenic granite intrusion the underlying metamorphic crust exerted a 'socle effect', which accentuated the upward transfer of mantle heat which is brought at the base of the crust by Andrews-Sleep convection. This 'late orogenic' heat is also regarded as being responsible for the Abukuma-type overprint of the older higher grade metamorphic rocks which were brought into a near-surface position by nappe and thrust tectonics. The regions subjected to pre-Middle Devonian folding and prograde regional metamorphism had already been deeply eroded by the Middle Devonian. Therefore, younger sediments suffered a weak prograde metamorphism, whereas the older metamorphic rocks were uplifted and subjected to retrograde overprinting under Abukuma-type metamorphic conditions. The development is, however, not uniform, but mirrors heterogeneities of the crustal structure and the heat flow. There are areas with strong late to post-tectonic igneous activity, e.g. the Odenwald, which contrasts with the neighbouring Saar-Province where the post-Lower Devonian sediments overlying the crystalline basement have remained nonmetamorphic even at a depth of 5000 m. In the case of the subsequent Permian magmatism the generation of the rhyolitic magmas requires a high temperature at the crust-mantle boundary. The model in Fig. 7 attributes the high temperature to the ascending limb of an Andrews-Sleep cell that migrates towards the orogenic foreland. The fact that rhyolitic volcanism had already occurred in the Black Forest in the Lower Carboniferous allows the interpretation that the ascending limb of the convection cell has spread out, from the Lower Carboniferous to the Permian, from the Black Forest to the Saar-Nahe trough. It could be assumed that the convective uprise of asthenospheric material below the crust produces spreading movements in the overlying crust similar to back-arc situations. However, the bipolar structure of the Variscan orogen and the assumed bilateral subcrustal

subduction implies strong convergent movements. Such a movement pattern does not allow the kind of crustal spreading encountered in marginal seas. Only the late orogenic magmatic activity can be understood as an expression of 'subsequent' rifting processes induced into the crust by the ascending limbs of Andrews-Sleep cells at a time when during the late stages of A-subduction and retreat of the subducting slabs the convergent movements became less and less effective. Finally during the Lower Permian the subcrustal subduction and crustal convergence came to an end.

Conclusions The so-called 'Caledonian' thermal event represents one of the main problems of the Variscan crustal development of central, western and southern Europe. The geological data discussed in this paper allow an interpretation of this event as resulting from continental rift processes. These took place on top of an anomalous mantle which induced igneous and bimodal volcanic activity during almost continuous sedimentation, and high grade metamorphism at the base of the crust. These rifting processes could have taken place inside a stationary crustal field between Laurasia and Gondwana as supposed by Zwart & Dornsiepen (1978) and Weber and Behr (1983) or, in the sense of Ziegler (1982), at the northern border of Gondwana. From here, Cadomian (Panafrican) consolidated crustal fragments in the form of allochthonous terranes were transported to the north where they were incorporated by collision into the Variscan belt. More palaeomagnetic data are necessary to prove these models. Nevertheless, the existence of Lower Palaeozoic rift processes seems to be well established, no matter which of the two models one prefers. The main phase of crustal shortening and accompanying regional metamorphism in the central parts of the Variscan orogen took place during the Lower Devonian. However, crustal shortening and thrust and nappe tectonics were active in the central zones up to the end of the Carboniferous, and the external zones of the Variscides were first deformed at this time. The former rift zones were the sites from which deep reaching crustal imbrications (subfluence zones in the sense of Weber 1978,1981) and nappe tectonics developed and which trace

18

K. Weber

out the main structural boundaries, e.g. the boundaries between the R h e n o h e r c y n i a n and Saxothuringian zones and the Saxothuringian and Moldanubian zones. A-subduction, i.e. subduction of lithospheric mantle underneath continental crust (Ampferer-subduction in the sense of Weber 1981 or delamination in the sense of Bird 1978), is regarded as the driving mechanism of crustal shortening. A-subduction follows B-subduction when the collisional stage is reached. That applies especially to the Ligerian suture. Whether small oceanic basins have been developed in other parts of the Variscides, particularly in the northern part along the boundaries between the Rhenohercynian and Saxothuringian zones and Saxothuringian and Moldanubian zones cannot yet definitely be answered. However, the main effects which can be observed there are the result of A-subduction. A peculiarity in the development of the European Variscides in comparison to Cordilleran and island arc type orogens is the missing availability of any mid-ocean ridges at which

mass-equilibration could be effected. Therefore, any subducting segment of lithospheric mantle has to retreat in order to allow massequilibration. This mass-equilibration which takes place in front of the subducting slabs leads to an uprise of hot asthenospheric material at the base of the overlying continental crust. High heat flow (low pressure/high temperature metamorphism) and the formation of a secondary stress field (back-folding) might be attributed to forced convection possibly in the form of an A n d r e w s - S l e e p cell. Finally, the formation of vast masses of late to post-orogenic granites, of bimodal volcanics and ignimbrites might be interpreted as the result of a restricted subsequent rifting event, which evolved on top of the convecting asthenosphere when A-subduction and crustal convergence gradually ceased. ACKNOWLEDGMENTS: I owe my thanks to H. Ahrendt, H. J. Behr, W. Engel and W. Franke for many helpful discussions and to S. C. Matthews for providing the translation.

References

ARENDT,H., CLAUER,N., HUNZIKER, J. C. & WEBER, K. 1983. Migration of folding and metamorphism in the Rheinische Schiefergebirge deduced from K-Ar and Rb-Sr age determinations. In: MARTIN, H. • EDER, W. (eds) Intracontinental Fold Belts--Case Studies in the Variscan Belt o f Europe and the Damara Orogen o f Namibia, 323-38. Springer-Verlag, Berlin. --, HUNZIKER, J. C. • WEBER, K. 1978. K/ArAltersbestimmungen an schwach metamorphen Gesteinen des Rheinischen Schiefergebirges. Z. dt. geol. Ges. 129, 229-47. AMPFERER, O. 1906. Uber das Bewegungsbild von Faltengebirgen. Jb. geol. Bundesanst., Wien, 56, 539-622. ANDREWS, D. J. & SLEEP, N. H. 1974. Numerical modelling of tectonic flow behind island arcs. Geophys. J. R. astr. Soc. 38, 237-51. ARNOLD, A. 1970. On the history of the Gotthard Massif (Central Alps, Switzerland). Eclog. geol. Helv. 63, 29-30. -&; SCHARBERT, H. G. 1973. Rb-Sr Altersbestimmungen an Granuliten der stidlichen B6hmischen Masse in Osterreich. Schweiz. miner, petrogr. Mitt. 53, 61-78. ASHBY, M. F. & VERRALL, R. A. 1977. Micromechanisms of flow and fracture, and their relevance to the rheology of the upper mantle. Phil. Trans. R. Soc. A, 288, 59-93. AUTRAN, A. & COGNI~, J. 1980. La zone interne l'orog~ne varisque dans l'ouest de la France et sa place dans le dtveloppement de la chMne hercynienne. In: COGNE, J. & SLANSKV, M. (eds)

G~ologie de l'Europe du Prdcambrien aux bassins s~dimentaires post-Hercyniens. Mem. Bur. Rech. Ggol. Min. 108, 90-111. BEDERKE, E. 1957. Alter und Metamorphose des kristallinen Grundgebirges im Spessart. Ahb. hess. Landesanst Bodenforsch. 18, 7-19. BEHR, H. J. 1961. Beitr~ige zur petrographischen und tektonischen Analyse des S/ichsischen Granulitgebirges. Freiberger ForschHft. 119, 1-146. 1978. Subfluenz-Prozesse im GrundgebirgsStockwerk Mitteleuropas. Z. dt. geol. Ges. 129, 283-318. - 1980. Polyphase shear zones in the granulite belts along the margins of the Bohemian Massif. J. struct. Geol. 2, 1-2, 249-54. ~, ENGEL, W. & FRANKE, W. 1980. Mtinchberger Gneismasse und Bayerischer Wald. Guide to Excursion, int. Conf.: the Effect of Deformation on Rocks. Gtttingen. 100 pp. --, & ~ 1982. Variscan Wildflysch and nappe tectonics in the Saxothuringian Zone (Northeast Bavaria, West Germany). Am. J. Sci. 282, 1438-70. & WEBER, K. 1980. Subduktion oder Subfluenz im mitteleuropaischen Varistikum. Berl. geowiss. Abh. (A) 19, 22-23 (abstr.), Int. A. Wegener-Symposium. BIRD, P. 1978. Initiation of intracontinental subduction in the Himalaya. J. geophys. Res. 83, 4975-87. BLOMEL, P. 1977. Stoffbestand und Metamorphose im Bayerischen Moldanubikum, insbesondere der

Variscan events progressiv metamorphen Serie des n6drlichen Bayerischen Waldes ("Arber-Osser Serie"). ln: La chafne Varisque d'Europe Moyenne et Occidentale, Colleques int. Cent. natn. Rech. scient. 243,349-57. BREWER, M. S. & LIPPOLT, H. J. 1972. Isotopische Altersbestimmungen an Schwarzwals-Gesteinen , eine Ubersicht. Fortschr. Miner. 5 0 , 2, 42-50. .... & 1974. Petrogenesis of basement rocks of the Upper Rhine region elucidated by Rubidium-Strontium systematics. Contr. Miner. Petrol. 45, 123-41. BRINKMAN, R. 1948. Die Mitteldeutsche Schwelle. Geol. Rdsch. 36, 56-66, Stuttgart. BUNTEBARTH, G. 1982. Zur Pal~iogeothermik im Permokarbon der Saar-Nahe-Senke. Nachr. dt. geol. Ges. 27, 44. --, KOPPE, J. & TEICHMLrLLER, M. 1982. Palaeogeothermics in the Ruhr Basin. In: CERMAK, V. & HAENEL, E. (eds) Geothermics and Geothermal Energy. 45-55, Stuttgart. BURG, J. P. 1977. Tectonique et microtectonique des sdries cristallophylliennes du Haut-Allier, et de la val6e de la Truy6re. These 3Ome cycle, 79, Montpellier. -&: MATTE, P. J. 1978. A cross-section through the French Massif Central and the scope of its Variscan geodynarnic evolution. Z. dt. geol. Ges. 129, 429-60. CALSTEREN, VAN, P. W. & DEN TEX, E. 1978. An early Palaeozoic continental rift in Galicia (W Spain). In: RAMBERG, J. B. & NEtJMANN, E. R. (eds) Tectonics and Geophysics of Continental Rifts, 125-32. Reidel, Dordrecht. - - , BOELRIJK, N. A. I. M., HEBEDA, E. H., PRIEM, H. N. A., TEX, E. DEN, VERDURMEN, E. A. TH. & VERSCHURE, R. H. 1979. Isotopic dating of older elements (including the Cabo Ortegal mafic-ultramafic complex) in the Hercynian orogen of NW Spain: manifestations of a presumed early Paleozoic mantle plume. Chem. Geol. 24, 35-56. CARTER, N. L. 1976. Steady state flow of rocks. Rev. Geophys. Space Phys. 14, 301-60. COGNr J. 1974. Le Massif Armoricain. In: DEBELMAS, J. (ed.) Gdologie de la France, 105-61. Paris. & WRIGHT, A. E. 1980. L'orogene cadomien. In: COGNE, J. & SLANSKY,M. (eds) Gdologie de l'Europe du Pr(cambrien aux Bassins S(dimentaires post-Hercyniens. Mere. Bur. Rech. GEol. Min. 108, 29-55. DE CHARPAL, O., GUENNOC, P., MONTADERT, L. & ROBERTS, O. G. 1978. Rifting, crustal attenuation and subsidence in the Bay of Biscay. Nature, 275, 706-11. CHRISTIE, P. A. F. & SCLATER, J. G. 1980. An extensional Origin for the Buchan and Witchground Graben of the North Sea. Nature, 283, 729-32. CLAUER, N. & BONHOMME, M. 1970. Datations rubidium-strontium dans les schistes de Steige et la s4rie de Ville (Vosges). Bull. Serv. Carte g4ol. Als. Loft. 23, 191-208. DORNSIEPEN, U. F. 1979. Rb/Sr whole rock ages

19

within the European Hercynian. A review. Krystallinikum, 14, 33-49. DUFOUR, E., PIBOULE, M. & DUTHOU, J. L. 1983. Les granulites des Moats du Lyonnais (Massif Central Frangais): evolution m6tamorphique et premiers r6sultats radiom6triques Rb/Sr. Terra Cognita, 3, 197 (abstr.) DUTHOU, M. 1978. Les granitoides du HautLimousin. Chronologie Rb/Sr. Le thermom6tamorphisme carbonif~re. Bull. Soc. gOol. Fr. (7), XX, 3,229-36. EMMERMANN, R. 1976. A petrogenetic model for the origin and evolution of the Schwarzwald. Neues Miner. Pali~ont. Mh. Abh. 128, 219-53. ENGEL, W., FEIST, R. & FRANKE, W. 1978. Synorogenic gravitational transport in the Carboniferous of the Montagne Noire (S. France). Z. dt. geol. Ges. 129, 461-72. & 1981. Le Carbonif6re ant6Stdphanien de la Montagne Noire: Rapports entre mise en place des nappes et s6dimentation. Bull. Bur. Rech. GOol. Min. (deuxi6me s6rie), Sec. I, No. 4, 341-89. --, FRANKE, W., GROTE, G., WEBER, K., AHRENDT, H. & EDER, W. 1983. Nappe tectonics in the southeastern part of the Rheinische Schiefergebirge. In: MARTIN, H. & LoeB, W. (eds) Intracontinental Fold Belts--Case Studies in the Variscan Belt of Europe and the Damara Orogen o f Namibia, 267-88. Springer-Verlag, Berlin. & - 1983. Flysch-sedimentation: its relations to tectonism in the European Variscides. In: MARTIN, H & EDER, W. (eds) Intracontinental Fold Belts--Case Studies in the Variscan Belt of Europe and the Damara Belt o f Namibia, 289-321. Springer-Verlag, Berlin. ENGELS, J. P. 1972. The catazonal polymetamorphic rocks of Cabo Ortegal (NW Spain); a structural and petrofabric study. Leidse geol. Med. 48, 83-133. FLOYD, P. A. 1982. Chemical variation in Hercynian basalts relative to plate tectonics. J. geol. Soc. London, 139, 4, 507-20. Fucns, G. 1971. Zur Tektonik des 6stlichen Waldviertels (N. t3.). Verh. Geol. Bundesanst., Wien, 424-40. 1983. The evolution of the Bohemian Massif in Austria. Terra Cognita, 3, 198 (abstr.) GABERT, G. 1957. Zur Geologie und Tektonik des nord6stlichen kristaUinen Vorspessart. Abh. hess. Landesanst. Bodenforsch. 18, 101-34. GEBAUER, D. & GR~NENFELDER,M. 1979. U - P b zircon and Rb-Sr mineral dating of eclogites and their country rocks. Example: Mfinchberg Gneiss Massif, northeast Bavaria. Earth planet. Sci. Lett. 42, 35-44. & 1983. On the oldest rocks of the central European Hercynides--rock chemistry, REE-, Sm-Nd and U-Pb zircon data. Terra Cognita, 3, 2-3, Second L U G meeting, p. 198 (abstr.). GIESE, P., JODICKE, H., PRODEHL, C. • WEBER, K. 1983. The crustal structure of the Hercynian Mountain System--a model for crustal thickening by staking. In: MARTIN, H. & EDER, W. (eds)

20

K. Weber

Intracontinental Fold Belts--Case Studies in the Variscan Belt o f Europe and the Damara Orogen o f Namibia, 405-26. Springer-Verlag, Berlin. GOETZE, C. 1978. The mechanisms of creep in olivine. Phil. Trans. R. Soc. A , 288, 99-119. GRAUERT, B., HANNY, R. & SOPTRAJANOVA, G. 1971. Isotopic ages of paragneisses and anatectic rocks of the Moldanubikum of Eastern Bavaria (abstr.) A n n & Soc. g~ol. Belg. 94, 115. HEARD, H. C. 1976. Comparison of flow properties of rocks at crustal conditions. Phil. Trans. R. Soc. A, 283, 173-86. HEINISCH, H. & SCHMIDT, K. 1976. Zur kaledonischen Orogenese in den Ostalpen. Geol. Rdsch. 62, 2, 459-82, Stuttgart. HELLMANN, K. N., L1PPOLT, H. J. & TODT, W. 1982. Interpretation der Kalium Argon-Alter eines Odenw/ilder Granodioritporphyritgafiges und seiner Nebengesteine. Aufschlufl, 33, 155-64, 6 Abb., Heidelberg. HOFMANN, A. W. 1979. Geochronology of the crystalline rocks of the Schwarzwald. In: J.a,GER, E. & HUNZIKER, J. C. (eds) Lectures in Isotope Geology, 215-21, Springer-Verlag, Berlin. & KOHLER, H. 1973. Whole rock Rb-Sr ages of anatectic gneisses from the Schwarzwald, SW Germany. Neues Jb. Miner. Abh. 119, 163-87. HUBREGTSE, J. J. M. V. 1973. Petrology of the Mellid area, a Precambrian polymetamorphic rock complex, Galicia, NW Spain. Leidse geol. Med. 49, 9-31. JAGER, E. 1977. The evolution of the central and west European continent. In: La Chafn Varisque d'Europe Moyenne et Occidentale. Colloques int. Cent. natn. Rech. scient. 2 4 3 , 2 2 7 - 3 9 . & WATZNAUER, A. 1969. Einige Rb/SrDatierungen an Granuliten des S/ichsichen Granulitgebirges. Mber. dt. Akad. Wiss. Berl. 11, 420-26. JAKI~S, P., ZOUBEK, J., ZOUBKOVA, J. & FRANKE, W. 1979. Graywackes and metagraywackes of the Tepla-Barrandian Proterozoic area. J. geol. Sci. 33, 83-122. KIRBY, S. H. 1980. Tectonic stresses in the lithosphere: constraints provided by the experimental deformation of rocks. J. geophys. Res. 85, 6353-63. KOBER, B. & LIPPOLT, H. J. 1983. Lead isotopes in K-feldspars and rocks from the southern Schwarzwald, SW Germany. Terra Cognita, 3, no. 2-3, 199 (abstr.) KOHLER, H., CHRISTINAS,P. & MULLER-SOHNIUS,D. 1983. Rb-Sr thin slab measurements on blastomylonitic rocks from the Bavarian Moldanubicum. Terra Cognita, 3, 200 (abstr.) KOPPEL, V., GUNTHERT, A. & GRUNENFELDER, M. 1981. Patterns of U-Pb zircon and monazite ages in polymetamorphic units of the Swiss Central Alps. Schweiz. miner, petrogr. Mitt. 61, 97-119. KOSINOWSKI, M. 1982. Die strukturelle Entwicklung der Scherzone Cordoba/Abrantes zwischen Elvas und Portalegre (Portugal) und ihre Stellung in den iberischen Varisziden (Metamorphose, Mikrogefiige, Magmatismus). Diss. University of G6ttingen. 120 pp.

KRAMM, U. 1982. Die Metamorphose des VennStavelot-Massivs, nordwestliches Rheinisches Schiefergebirge: Grad, Alter und Ursache. Decheniana (Bonn), 135, 121-78. KREUZER, H. & HARRE, W. 1975. K/Ar Altersdatierungen an Hornblenden und Biotiten des kristallinen Odenwaldes. Aufschlufl Sonderheft, 27, 71-78. - - , LENZ, H., HARRE, W., MATFHES, S., OKRUSCH, M. & RICHTER, P. 1973. Zur Altersstellung der R6tgneise im Spessart, Rb/Sr Gesamtgesteins datierungen. Geol. Jb. 9, 69-88. KUIJPER, R. P. 1979. U-Pb systematics and the petrogenetic evolution of infra-crustal rocks in the Paleozoic basement of western Galicia (NW Spain). Verh. Nr. 5, Z W O Laboratorium voor lsotopen-Geology, Amsterdam, 1-110. LEHMANN, J. 1984. Untersuchungen fiber die Entstehung der altkristallinen Schiefergebirge mit besonderer Bezugnahme auf das S/ichsische Granulitgebirge, Erzgebirge, Fichtelgebirge und bayerisch-b6hmische Grenzgebirge. Bonn. LENZ, H. & MOLLER, P. 1976. Radiometrische Altersbestimmungen am Kristallin der Bohrung Saar 1. Geol. Jb. 27, 429-32. LWPOLT, H. J., BARANY, J. & RACZEK, J. 1976. Rb/Sr chronology of orthogneisses in the eastern Odenwald and southern Spessart (Germany). Abstract ECOG W , Amsterdam. L~STER, G. S. & DORNSlEPEN, U. F. 1982. Fabric transitions in the Saxony granulite terrain. J. struct. Geol. 4, 81-92. MAASKANT, P. 1970. Chemical petrology of polymetamorphic ultramafic rocks from Galicia, NW Spain. Leidse geol. Med. 45, 237-325. MAASS, R. 1981. The Variscan Black Forest. Geologic Mi]nb. 60, 1, 137-44. -& STOPPEL, D. 1982. Nachweis von Oberdevon bei Markstein (B1. Munster, Siidvogesen). Z. geol. Ges. 133, 403-8. MAGGETTI, M. 1974. Zur Dioritbildung im kristallinen Odenwald. Schweiz. miner, petrog. Mitt. 54, 1, 39-57. 1975. Die Tiefengesteine des Bergstr/il3er Odenwaldes. Aufschlufl, Sonderband 27, 87-107, Heidelberg. -& NICKEL, E. 1976. Konvergenzen zwischen Metamorphiten und Magmatiten, Geol. Jb. 104, 147-60. MASSONNE, H.-J. & SCHREVER, W. 1983. A new experimental phengite barometer and its application to a variscan subduction zone at the southern margin of the Rhenohercynicum. Terra Cognita, 3, 187 (abstr.). MATTHES, S. 1954. Die Paragneise im mittleren kristallinen Vorspessart und ihre Metamorphose. Abh. hess. Landesanst. Bodenforsch. 8, 1-86, Wiesbaden. -& -1965. Spessart. Samml. geol. Fiihr. 44, Borntr~iger. -& -1977. The Spessart, crystalline complex, north-west Bavaria: rock series, metamorphism, and position within the Central German Crystalline Rise. In: La Chatne Varisque d'Europe Moyenne et Oecidentale. Colloques int. Cent. natn. Rech. seient. 243,375-90.

Variscan events MCKENZ1E, D . 1978a. Some remarks on the development of sedimentary basins. Earth planet. Sci. Lett. 40, 25-32. 1978b. Active tectonics of the AlpineHimalayan belt: the Aegean Sea and surrounding regions. Geophys. J. R. astr. Soc. 5 5 , 217-54. MEER MOHR VAN DER, C. G., KUIJPER, R. P., VAN CALSTEREN, P. W. C. & DEN TEX, E. 1981. The Hesperian Massif: from Iapetus aulacogen to ensialic orogen. A model for its development. Geol. Rdsch. 70, 2,459-72. MEISSNER, R., BARTELSEN, H. & MURAWSKI, H. 1981. Thin-skinned tectonics in the northern Rhenish Massif, Germany. Nature, 290, 5805, 399-401. MERCIER, J.-C. 1980. Magnitude of the continental lithospheric stresses inferred from rheomorphic petrology. J. geophys. Res. 85, B1, 6293-303. , ANDERSON, D. A. (~ CARTER, N. L. 1977. Stress in the lithosphere: inferences from steady-state flow of rocks. Pure appl. Geophys. 115, 199-226. MONTADERT, L., DE CHARPAL, O., ROBERTS, D. C. & GUENNOC, P. 1979. Rifting and subsidence of the Northern Continental Margin of the Bay of Biscay. Init. Rep. Deep Sea drill. Pro]. 48, 1025-60. MURAWSKI, H. 1958. Der geologische BaH des zentralen Vorspessart. Z. dt. geol. Ges. 110, 360-88. NICKEL, E. & MAGGETTI, M. 1974. Magmententwicklung und Dioritbildung im synorogenen konsolidierten Grundgebirge des Bergstrasser Odenwaldes. Geol. Rdsch. 63, 618-54. NICOLAS, A. & POIRIER, J. P. 1976. Crystalline Plasticity and Solid State Flow in Metamorphic Rocks. Wiley, London, New York. 444 pp. OKRUSCH, M., RAUMER, J., MATZHES, S. & SCHUBERT, W. 1975. Mineralfazies und Stelling des Odenwaldkristallins. Aufschlufl, Sonderband 27, 109-34, Heidelberg. , STREIT, R. & WEINELT, W. 1967. Erliiuterungen zur Geologischen Karte von Bayern 1: 25.000, BI. 5920. Alzenau i. Ufr., Miinchen. LE PICHON, X., ANGELIER, J. & SIBUET, J.-C. 1982. Plate boundaries and extensional tectonics. Tectonophys. 81, 239-56. - & SIBUET, J.-C. 1981. Passive margins: a model of fermation. J. geophys. Res. 86, 3708-20. PLESSMAN, W. 1957. Zur Baugeschichte des nordwestlichen kristallinen Spessart. Abh. hess. Landesanst. Bodenforsch. 18, 149-66. POST, R. 1977. High temperature creep of Mt. Burnet dunite. Tectonophys. 42, 75-110. PRIES, H. N. A., BOELRIJK, N. m. I. M., VERSCHURE, R. H., HEBEDA, E. H. & VERDURMEN, E. A. T. 1970. Dating events of acidic plutonism through the Palaeozoic of the Western Iberian Peninsula. Eclog. geol. Helv. 63, 255-74. RAUMER, VON J. F. 1981. Variscan events in the Alpine region. Geologie Mijnb. 1, 67-80. RINGWOOD, A. E. 1975. Composition and Petrology o f the Earth's Mantle. International Series in the Earth and Planetary Sciences. McGraw-Hill, New York. 618 pp. ROYDEN, L., SCLATER, J. C. & VON HERZEN, R. P.

21

1980. Continental margin subsidence and heat flow: important parameters in formation of petroleum hydrocarbons. Bull. Am. Ass. Petrol. Geol. 64, 173-87. SATXLER-KOSlNOWSK1, S. 1982. Die strukturelle Entwicklung der Scherzone Cordoba/Abrantes zwischen Elvas und Portalegre (Portugal) und ihre Stellung in den iberischen Varisziden. (Strukturplan, Mikrothermometrie, Stratigraphie). Diss. University of G6ttingen. 121 pp. SCHARBERT, H. G. 1977a. Tiefe Kruste und oberer Mantel in der Moldanubischen Zone Nieder6sterreichs In: La Chafne Varisque d'Europe Moyenne et Occidentale. Colloques int. Cent. natn. Rech. Scient. 243, 193-8. SCHARBERT, S. 1977b. Neue Ergebnisse radiometrischer Altersbestimmungen an Gesteinen des Waldviertels. In: MURATA, A. (ed.) Fiihrer Arbeitstagung Geol. Bundesanst. Waldviertel, 11-15 Wien (Geol. B.-A.). SCHEUMANN, K. H. 1932. Uber die petrogenetische Ableitung der roten Erzgebirgsgneise. Miner.petrogr..Mitt. 42, 413-54. - 1939. Uber die petrographische und chemische Substanzbestimmung der Gesteinsgruppe der Roten Gneise des Sfichsischen Erzgebirges und der angrenzenden R/iume. SCHMIDT, K. 1976a. Das "kaledonische Ereignis" in Mittel- und Siidwesteuropa. Nova Acta Leopoldina, N. F. 224, 45, 381-401, Halle. - 1976b. "Subfluenze" und "Subduktion" in den Alpen. Z. dt. geol. Ges. 127, 53-72. SCHOELL, M., LENZ, H. & HARRE, W. 1973. Das Alter der Hauptmetamorphose des Eckergneises im Harz auf Grund von Rb/Sr-Datierungen. Geol. Jb. A9, 89-95. SCHONLAUB, H. P. & SCHARBERT, S. H. 1978. The early history of the eastern Alps. Z. dt. geol. Ges. 129, 473-84. SCHOLTZ, H. 1930. Das varistische Bewegungsbild, entwickelt aus der inneren Tektonik eines Profils von der Brhmischen Masse bis zum Massiv von Brabant. Fortschr. Geol. Pali~ont. 8, (25) 235-316. SCHREYER, W. (~ ABRAHAM, K. 1978. Prehnite/ chlorite and actinolite/epidote bearing mineral assemblages in the metamorphic igneous rocks of la Hella and Chales, Ven-Stavelot Massif, Belgium. Ann. Soc. g~ol. Belg. 101, 227-41. SOLLNER, F. K()HLER, H, • M(JLLER-SOHNIUS, D. 1981. Rb/SrAltersbestimmungen an Gesteinen der Miinchberger Gneismasse (MM), NEBayern--Teil 1, Gesamtgesteinsdatierungen. Neues Jb. Miner. Geol. Pal~ont. Abh. 141, 1, 90-112. STEIGER,R., BAR, M. & BUSCH, W. 1972. The zircon age of an anatectic rock of the Central Schwarzwald. Fortschr. Miner. 5 0 , 3, 131-2. STOKER, R. L. & ASHBY, M. F. 1973. On the rheology of the upper mantle. Rev. Geophys. Space Phys. 11, 391-426. TEX, E. DEN 1981. A geological section across the Hesperian Massif in western central Galicia. Geologie Mij~b. 60, 33-40.

22

K. W e b e r

1982. Dynamothermal metamorphism across the continental crust/mantle interface. Fortschr. Miner. 60, 1, 57-80. TmELE, O. 1976a. Ein westvergenter kaledonischer Deckenbau im nieder6sterreichischen Waldviertel? Jb. geol. Bundesanst, Wien, 119, 75-81. 1976b. Zur Tektonik des Waldviertels in Nieder6sterreich (siidliche B6hmische Masse). Nova Acta Leopoldina, N.F., N. 224, 45, 67-82 (Franz Kossmat Symposium). TODT, W. 1979. U-Pb-Datierungen an Zirkonen des kristalinen Odenwaldes. Fortschr. Miner. 57, 1, 153-4. TOLLMANN, A. 1982. Grossrfiumiger variszischer Deckenbau im Moldanubikum und neue Gedanken zum Variszikum Europas. Geotekt. Forsch. 64, I-II, 1-91. VEJNAR, Z. 1971. Grundfragen des Moldanubikums und seine Stellung in der B6hmischen Masse. Geol. Rdsch. 60, 4, 1455-65. VETrER, U. R. & MEISSNER, R. 1979. Rheologic properties of the lithosphere and applications to Passive Continental Margins. Tectonophys. 59, 367-80. VIDAL, P. 1976. L'dvolution polyorog6niques du Massif armoricain: apport de la g6ochronologie et de la gdochimie du strontium. 7hOse. University of Rennes. 140 pp. - - , AUVRAY, B., CHARLOT, R. & COGNE, J. 1981. Precambrian relicts in the Armorican Massif: their age and role in the evolution of the western

and central European Cadomian-Hercynian Belt. Precambrian Res. 14, 1, 1-20. WATZNAUER, A. 1974. Beitrag zur Frage des zeitlichen Ablaufs der Granulitgenese (Sfichsisches Granulitgebirge). Krystalinikum, 10, 181-92, Praha. WEBER, K. 1978. Das Bewegungsbild im Rhenoherzynikum--Abbild einer varistischen Subfluenz. Z. dt. geol. Ges. 129, 249-81. 1981. The structural development of the Rheinische Schiefergebirge. Geologie Mijnb. 60, 1, 149-59. & BEHR, H. J. 1983. Geodynamic interpretation of the mid-European Variscides. In: MARTIN, H. & EDER, W. (eds) Intracontinental Fold Belts-Case Studies in the Variscan Belt of Europe and the Damara Orogen Of Namibia, 427-72. Springer-Verlag, New York. Z1EGLER,P. A. 1982. Geological Atlas of Western and Central Europe. Shell International Petroleum Maatschappij B. V. 130 pp., 40 encls. ZIMMERLE, W. 1976. petrographische Beschreibung und Deutung der erbohrten Schichten. In: Die Tiefbohrung Saar 1, Geol. Jb. A27, 91-305. ZWART, H. J. 1976. Regional metamorphism in the Variscan orogeny of Europe. Nova Acta Leopoldina, N.F. Nr. 224, B. 45, 361-7 (Franz Kossmat-Symposium). & DORNSIEPEN, U. F. 1978. The tectonic framework of Central and Western Europe. Geologie Mi]nb. 57, 627-54. -

-

KLAUS WEBER, Geologisch-Pal/iontologisches Institut und Museum Goldschmidtstr, 3, D-3400, G6ttingen, Germany.

Tectonics of the Variscides in North-Western Germany based on seismic reflection measurements R. Meissner, M. Springer & E. Fifth SUMMARY: In the area of the Variscides in Germany five seismic reflection surveys were carried out between 1968 and 1978. Near Aachen, at the very northern part of the Variscan deformation front, a thin-skinned overthrust fault was found, while farther south, at the Hunsriick border fault, a steep listric fault zone was mapped which seems to have been initiated as an overthrust, but developed into a deep reaching extensional fault during the post-Variscan formation of the Saar-Nahe trough. The reflection signature of the different experiments, the low seismic velocity in the middle and lower crust, and the generally small crustal thickness together with many geological and petrological observations are compatible with the assumption that during the Variscan orogenies an interstacking of predominantly sialic platelets took place in a generally high-temperature environment. The shifting of the collisional belts from SE to NW is opposite to that of the Appalachian orogenies, although time periods and tectonic framework were similar. A simplified concept of approaching thin sialic platelets toward the rugged remnants of the Caledonian orogeny is presented.

It was the concept of a group of geoscientists associated with the programme 'Geotraverse Rhenohercynicum' to examine some of the most interesting parts of the Variscan mountain systems in N W Germany by seismic reflection profiles. As shown by some early experiments in 1964 in Bavaria and presently by the success of the US C O C O R P programme (Oliver 1980, 1982) reflection seismology may be considered as the most powerful tool for investigating the structure of the earth's crust. In addition to the structural data, the detection of fault zones, sometimes cutting the whole crust, and the determination of interval velocities, also for the lower crust, are important features of seismic reflection methods. Within the Variscan mountain system a combination of steep-angle and wide-angle reflection observations was always used. In these experiments mobile refraction stations were set up along the profiles at distances of up to 180 km from the shot point, thereby observing the same explosions as those fired for the reflection work. This, additional, use of refraction stations in the wide-angle area provides us with important information on the general velocity structure of the crust: information which is not obtained by the routine short-spread m e t h o d of steep-angle reflection observations. As the detailed technique of the various seismic reflection experiments in the Variscan mountain system and the basic results are described elsewhere (Clfment 1963; Bartelsen 1970; Meissner & Vetter 1976; Glocke & Meissner 1976; Bless et al. 1980; Meissner et al. 1980, 1981, 1983; Bartelsen et al. 1982), this

paper will concentrate on some new data from the 1978 profile near Aachen and on the consequences of a comparison of the various profiles with regard to the development of the Variscan collision belts.

Important features of the reflection data Figure 1 shows a situation map of the various reflection lines in the Variscan mountain system. One of the important features on the northern profile near Aachen is a strong, shallow reflector at a depth of 3 - 4 km, dipping slightly to the south. Its extrapolation to the north coincides with the well-known Aachen thrust fault which is known as the Faille du Midi or Condroz overthrust in Belgium and France. It marks the northernmost Variscan deformation front. Figure 2(A) shows this reflector in a record section, (B) shows its reflection polarity and (C) is a geological crosssection of the area after Teichmiiller & Teichmfiller (1979). The interpretation of the reflector as the listric continuation of the Aachen overthrust fault below the subsurface of the C a m b r o - O r d o v i cian Hohes Venn anticline is based on the following arguments: (1) the geometric extrapolation as mentioned above; (2) the interpretation of the reflector as a thin layer as deduced from the absence of any break or j u m p of the first (refraction) arrivals; 23

24

R. M e i s s n e r e t al.

---~~riscan

- ~;I~

4_

A/

Foredeep -.%

--..,

/

49 ~

9

/ I" f

+ 6~ 1

tlIla

+~o~d~~b~c + 8* I

I0 ' I

FIG. 1. Situation map of deep-seismic reflection surveys including wide angle observation in the Variscan area in central Europe. A = Aachen-Hohes Venn profile (6-fold coverage, 3 x 7.2 km spread length). HF = Hunsriick border fault profile (3 to 4-fold coverage, 2.4 km spread length). U = Urach profiles (8-fold coverage, 23 km spread length). The other lines show locations of standard exploration surveys (Belgium, France) or short-segment deep-seismic lines (Hunsr/ick, Rhinegraben). (3) the investigation of the polarity of the reflector which showed a predominance of negative reflection coefficients in at least 70% of all traces as seen in Fig. 2(B) (Springer 1982). The negative reflection coefficient is a result of a material change from high to low impedance /9 9V (density times velocity). In this case with a thin layer p 9V inside the layer is smaller than above (and below) it. As the density p generally does not vary as much as V it is justified to speak of a low-velocity layer in the case of negative reflection coefficients: especially in our case where such a strong reflection is observed. The presence of a thin lower-velocity layer is consistent with petrological observations at fault zones, showing that increased pore pressure, grain-size diminution and a hydrolytic weakening in general results in greatly reduced seismic velocities in such a fault zone (Etheridge & Wilkie 1979). From amplitude assessments a thickness of 2 0 - 8 0 m is obtained for this zone (Springer 1982).

Similar reflectors were observed on the Belgian and French exploration profiles in the area south of the Faille du Midi and Condroz overthrust respectively. Here several boreholes penetrated Devonian sediments on top of Carboniferous layers. It therefore seems that sedimentary rocks may be generally present below the overthrust fault. A thin-skinned nappe, very similar to that found in the C O C O R P seismic lines in the southern Appalachians (Cook et al. 1980), marks the northern Variscan deformation front in that area. Looking next to the lateral extension of the thin-skinned nappe, the eastern continuation of the northern Variscan deformation front ( = N V D F ) across the lower Rhine e m b a y m e n t does not look very promising. Although not yet investigated by special seismic profiles, the detailed knowledge of shallow structures in the Ruhr area does not provide any indications of a large nappe. The continuation of the Faille du Midi or Condroz overthrust to the west and

N W German Variscides

25

26

R. Meissner et al.

north-west, however, was identified as a prominent overthrust in France (Laumondais, pets. comm.) and in Wiltshire, England by seismic reflection measurements (Kenolty et al. 1981). In Ireland, however, where the northern Variscan deformation front ( = N V D F ) turns west again, strike-slip features may be important (Sanderson, this volume; Max & Lefort, this volume). Looking to those zones along the N V D F where compression and overthrusting with the formation of thin nappes took place it is evident that the old London Brabant massif (LBM) must have played a key role. It most probably acted as a ramp towards the approaching orogenies from the south. Figure 3 shows a situation map of the N V D F with the L B M and the Welsh massif. It is not surprising that south of the southernmost margin of the LBM is the location of the strongest tectonic events (about the area of our Aachen profile in the Hohes Venn). Additionally the whole SW-part of the N V D F along the L B M and the Welsh massif forms a compressional zone more than 600 km in length, as indicated by the thin-skinned nappe overthrusts found there. Compression east of the Rhine embayment might have been generally smaller as indicated by the more northward extension of the N V D F and by the presence of rather mild compressional features in the Ruhr area and the area

further to the east. Here we feel that no strong ramp was present to interfere with the northward push of the Variscan orogeny. Northward movement of thin-skinned nappes might not have been the only expression of the compressional regime of the Variscan orogeny. Differential movements may and should have taken place also at deeper levels. Such movements may also be observed in the form of reflectors. Unfortunately, the strong dips in the stacked section of our Aachen profile in the middle and lower crust result in a complicated reflection pattern, as seen in Fig. 4(A). Only the Moho boundary appears as a weak but horizontally orientated band of reflections at about 10 s two-way travel time. The structure of the middle and lower crust between about 4.5 and 9.5 s is much better resolved by the migrated section, as seen in Fig. 4(B), showing the upper 10 s two-way travel time, i.e. approximately 30 km in depth. Prominent reflectors in the middle crust show a small syncline below the Venn area and a ramp-like southerly dip towards their northern end. Reflectors in the lower crust show a general northerly dip possibly marking the traces of subduction to the north. The M o h o - - a s seen in Fig. 4 ( A ) l c o n s i s t s of rather weak but horizontally aligned layers as if it were created as a new boundary (after the process of deformation was over) by means of crystallization after fractional differentiation

FIG. 3. The north Variscan deformation front (NVDF) between Ireland and Poland. Large triangles mark areas with seismically determined overthrust fault (thin-skinned nappe tectonics). LBM = London Brabant Massif. WM = Welsh Massif. a = supposed direction of maximum horizontal stress.

N W G e r m a n Variscides

27

FIG. 4. Cross-sections of the Aachen reflection profile with visible reflectors indicated: A = unmigrated section, B = migrated section, R = shallow reflector (thrust fault), CD = Conrad discontinuity, MD -- Moho discontinuity. of intruded basaltic magmas. This concept of cyclic layering was used by Meissner (1967) to explain the lamella-like appearance .of many Moho reflections in central Europe. Recently, Kushiro (1981) found experimental evidence that basaltic magmas at pressures higher than 6 kbars (equivalent to depths greater than 18 km) differentiate into ultramafic cumulates with light anorthositic material above. Several such penetration cycles could easily result in the observed cyclic layering. The other crustal reflectors in this profile at the northern end of the Variscan orogeny indicate differential, nearly horizontal movements of crustal layers to the north. Such movements might have been strongest along the upper

(Aachen) overthrust fault reaching farther to the north than the deeper reflectors which nevertheless might have also acted as thrust planes. Boundaries in velocity and composition are often also boundaries in viscosity (Meissner & Vetter 1979) and may act as easy-gliding planes. A fault of completely different character was found at the southern boundary of the Rhenish massif (Meissner et al. 1980). It had also started as an overthrust fault about 50 Myr earlier than the orogeny at the N V D F and is described by A h o r n e r & Murawski (1975), Weber (1981) and others. As seen in Fig. 5, which contains the seismic and geological information, it cuts the whole crust and even reaches the uppermost

28

R . M e i s s n e r et al.

FIG. 5. Line drawing of migrated cross-section of the reflection profile across the Hunsriick border fault. Seismic data from Meissner et al. (1980); geological data from Murawski (1976) and Weber (1981). a, b = Permian sediments, c = supposed alignment of fault zone. mantle. The listric shape in its lower part seems to be an effect of the extensional stresses which took over in Permian times and opened the Saar-Nahe trough in the south. As indicated by the different dips on both sides of the fault zone, most of the trough-forming movement must have taken place along this steep listric fault zone. The seismic sections of Figs 4 and 5 and some other measurements in the Variscan belts show that here the crust is generally thin, i.e. around 30 km, while shield areas show crustal thicknesses around 45 km and are much older and colder. Moreover, as shown by Meissner & Vetter (1976), the sialic part of the Variscan crust, with low densities and low seismic velocities, is comparatively thick reaching down to 20 and 25 km in places while shields show higher densities and higher velocities from 10 or 15 km downwards. A n o t h e r feature of the seismic sections in the Variscan belts is the high reflectivity of the lower crust which is generally not found in shield areas nor in oceanic surroundings. Figure 6 shows a comparison of the density of reflections, i.e. the frequency of their occurrence

per 0.5 s two-way travel along the five profiles. The increased reflectivity in the lower crust can be explained by an increased interaction between mantle and crust with basaltic magmas entering the hot and low-viscosity lower crust, spreading laterally and crystallizing in seams. Alternatively the reflectors could be interpreted as traces of large horizontal movements created during orogenies and possibly marking the base of listric faulting. Figure 6(A), the reflections histogram of our Aachen profile, shows that here the whole crust, and the lower crust in particular, is thicker than in the profiles (B) and (C) from the southern part of the Rhenish massif. This may be interpreted as evidence of a stronger compressional interstacking near the ramp of the LBM. By contrast the thicker crust in the Saar Nahe trough in Fig. 6(D) may have originated from its asymmetric subsidence along the Hunsrfick border fault in Permian times, the strong dips of reflectors being responsible for a certain smoothing of the frequency of occurrence of reflections. The Urach profiles (E and F) in the Moldanubian zone, show a still thinner crust and a still stronger

NW German Variscides 0

2-~

A

C

29 E

4-

J 6 20

8

1

30

12 t

40

~s

z

(kin)

ITTSlb FIG. 6. Histogram of the density of reflections across the seismic reflection profile in the Variscides as shown in Fig. 1. Reflections are normalized to the maximum reflection density per area. Locations: A = Aachen, B = southern part of Rhenish massif, C = southern Hunsriick, D = Permian Saar Nahe trough, E = Urach WSW-ENE, F = Urach NNW-SSE. Signatures: a = postCarboniferous sediments, b = thrust fault, c = pre-Permian sediments, crystalline crust and upper mantle, M = Moho. concentration of reflectors in the lower crust: possibly an effect of a strong interaction between mantle and crust during the development of the geothermal anomaly in this area (Bartelsen et al. 1982).

The development of the Variscan orogenies Age determination and geological records have shown that the Variscan orogenies moved from south to north (Weber 1981; Ziegler 1978; and many others). Shortly after the Caledonian orogeny petered out in the north the Variscan series of orogenies started in the Moldanubian, while different kinds of submarine volcanic activity around the mid-German high (= MG) indicate a tensional environment at this time. Spilites here, however, are mainly tholeiitic basalts and do not indicate an oceanic mantle (Herrmann & Wedepohl 1970). Also missing from this area are ophiolitic belts, so common in virtually all major orogenic mountain belts (Gass 1982). This may be taken as an indication that any oceanic environment was very limited. Some ophiolites of Variscan age, however, are incorporated in the Alpine belt, i.e. in the very southern part of the Variscan orogeny (Frisch 1977; and pers. comm.). The Appalachians on the other hand, which were formed during the same time interval as the Variscan mountain

system and in a similar plate tectonic environment, show at least two belts of ophiolites, i.e. strong indications of subducted (and partly obducted) oceans in between the approaching continental segments (Cook et al. 1980). Some of the Variscan sediments also show a deep water character (Ziegler 1978), and the assumption of the existence of trenches with at least some subduction seems unavoidable. An indication of subduction associated with the formation of the N V D F is the strong dip of deep crustal reflectors in our Aachen profile. Furthermore, the crustal shortening of several hundred km here cannot be explained without some kind of subduction process of subcrustal material. The foregoing arguments, hence, require subduction but deny the existence of oceanic basins, a situation which is somewhat controversial and certainly not common in today's plate tectonic framework. Weber (1981, this volume) assumes an 'A'-subduction to solve the problem of crustal shortening (A-subduction according to Ampferer 1906). Giese (1978) compares the tectonic style of the Variscides with some of the present Mediterranean features, using velocity and structural data from seismic refraction investigations. It seems indeed that the widespread low-pressure high-temperature character of the metamorphism in the Variscan system within the complicated pattern of continental platelets in the area between two major conti-

30

R. M e & s n e r et al.

FIG. 7. Tectonic development of the Variscan mountain belts. LBM, London Brabant massif. F, Fennoscandian shield. MGH, mid-German high. Mo, Moldanubian. nents (i.e. Eurasia and Africa), each with their rugged boundaries, may provide the key to the solution of the Variscan orogeny. Low-pressure high-temperature metamorphism with the widespread emplacement of granites implies a strong temperature gradient and high temperatures in the lower crust. Rifting in shallow seas possibly did not reach the stage of a true spreading but managed to reduce the crustal segments to rather thin sialic slices. This might have happened in a back-arc situation in front of a retreating subduction zone as postulated by Lorenz (1976) and Giese (1978). It might also indicate strong plume activity according to the stages 1 and 2 of Meissner (1981). The fact that the Variscan platelets were warm and thin is also underlined by their young age. Most Rb-Sr and K-Ar ages of orogenic granitoids are below 460 Ma (Ahrendt et al. 1978) and nowhere in the Variscides have ages larger than 800 Ma been found. It is known from earthquake-depth relationships and from viscosity calculations that even under moderate temperature conditions the lower crust is able to creep (Kuznir & Bott 1977; Meissner & Strehlau 1982). It is easy to imagine that under high-temperature conditions and with an appropriate stress pattern, lower crustal and upper mantle material can be laterally transferred very quickly and possibly partly incorporated into convection cells. This may explain the small crustal depths in the Variscides and the surprisingly high proportion of sialic material. Hence, thin, predominantly sialic platelets presumably formed the input of the Variscan orogeny. Platelets were stacked together during the compressional phases, guided by a mobile and ductile lower crust and upper mantle which moved according to global and regional stresses connected with the shift of the supercontinents Africa in the south and North America-Eurasia in the north. The assumption of shallow underplating of warm crustal and mantle material below the adjacent continents seems more reasonable than postulating a hypothetical A-subduction, a process which seems impossible in view of

the missing negative buoyancy of the warm material. The picture evolving from the geological and geophysical reasoning mentioned above is shown in Fig. 7(A, B, C) which is based on reconstructions and structural maps by Ziegler (1978), Lorenz (1976) and Weber (1981). Figure 7(A) gives a simplified picture of the tectonic situation in Europe in the Middle and Upper Devonian. The shift of the tectonic activity is to the Taunus-Hunsr/ick area and the whole MGH is connected with at least some 100 km of crustal shortening (Weber 1981) along some prominent overthrust faults, e.g. the Hunsr/ick border fault in its juvenile state. This tectonic event around 350 Myr ago is depicted in Fig. 7(B). As mentioned before, the continued movement of the tectonic activity to the north finally reached the area of the NVDF where the collision with the LBM led to major thin-skinned (and possibly also thick-skinned) overthrusts and to a formation of the tectonic front sub-parallel to the southern and southwestern margin of the London-Brabant and Welsh massifs; see Fig. 7(C).

Conclusions The shifting and interstacking of predominantly sialic platelets in a shallow marine environment under relatively high temperatures are considered the essential preconditions for the Variscan orogenies in central and Western Europe. Such platelets are by no means rare in today's oceans or shelves and cover up to 10% of the oceanic areas (Nur & Ben-Avraham 1982). As research continues it becomes clearer that not merely subduction but also collision between continental plates or platelets leads to an orogeny. Whether large oceanic areas were present between platelets as in the case of the Appalachians or shallow seas as in the case of the Variscides is considered a matter of secondary importance. Also the sequence of collisions moving towards the ocean in case of the Appalachians but moving towards the continent

N W German Variscides in the case of the Variscides m a y play an important but u n k n o w n role. Certainly m u c h m o r e i m p o r t a n t is the difference in t e m p e r a t u r e and viscosity of the crust in these two areas. H i g h t e m p e r a t u r e s p r o v i d e d the p r e c o n d i t i o n s for the g e n e r a t i o n of the thin, sialic platelets, for their interstacking and for u n d e r p l a t i n g of the w a r m u n d e r l y i n g material. A l t h o u g h apparently no large o c e a n i c sutures d e v e l o p e d , the a b o v e - m e n t i o n e d c o n c e p t for the Variscan o r o g e n i e s certainly fits into the f r a m e w o r k of m o d e r n plate tectonics.

31

ACKNOWLEDGMENTS: Thanks are due to our colleague H. Murawski, Frankfurt, for his initiative and steady encouragement of the combined geologicalgeophysical studies in the Rhenish Massif. Further discussions with K. Weber, G6ttingen and W. Frisch, Tiibingen, have helped us in developing our arguments. The seismic studies in the Rhenish Massif were supported by the Deutsche Forschungsgemeinschaft (German Research Association), and fieldwork was carried out by PRAKLA-SEISMOS, Hannover. Institut ffir Geophysik Publication No. 246.

References AHORNER, L. & MURAWSKI, H. 1975. Erdbebentfitigkeit und geologischer Werdegang der Hunsriick-Siidrand-St6rung. Z. dt. geol. Ges. 126, 63-82. AHRENDT, H. J., HUNZIKER, C. & WEBER, K. 1978. K/Ar-Altersbestimmungen an schwachmetamorphen Gesteinen des Rheinischen Schiefergebirges. Z. dt. geol. Ges. 129, 229-47. AMPFERER, O. 1906. Uber das Bewegungsbild von Faltengebirgen, Austria. Jb. geol. Bundesanst. Wien, 56, 539-622. BARTELSEN, H. 1970. Deutung der seismischen Weitwinkelmessungen in der Rheinischen Masse unter Verwendung neuer Kontroll- und Auswerteverfahren. Diploma Thesis. Institut fiir Geophysik, Frankfurt. 90 pp. , LUESCHEN, E., KREY, Th., MEISSNER, R., SCHMOLL, H. & WALTER, Ch. 1982. The combined reflection-refraction investigation of the Urach geothermal anomaly. In: HAENEL, R. (ed.) The Urach Geothermal Project. Schweizerbart'sche Verlagsbuchhandlung, Stuttgart. BLESS, M. J. M., BOUCKAERT, J. & PAPROTH, E. 1980. Environmental aspects of some PrePermian deposits in NW Europe. Meded. Rijks geol. Dienst. 32, 3-13. CLISMENT, J. 1963. R6sultats pr61iminaires des campagnes g6ophysiques de reconnaissance dans les permis de recherches 'Arras et Avesnes' de l'Association Shell Francaise-P.C.R.B.SAFREP-Objectifs du forage profond Jeumont-Marpent No. 1. Annls Soc. g~ol. N. 83, (2)7-42. COOK, F. A., BROWN, L. D. & OLIVER, J. E. 1980. The southern Appalachians and the growth of continents. Scient. Am. 243, 124-38. ETHERIDGE, M. A. & WILKIE, J. C. 1979. Grainsize reduction, grain boundary sliding and the flow strength of mylonites. Tectonophys. 58, 159-78. FRISCH, W. 1977. Plate-tectonic evolution of the Eastern Alps. Acta geol. Acad. Sci. hung. 21, 223-8. GASS, J. G. 1982. Ophiolite: Ozeankruste an Land. Spektrum Wissenschafien. 10/82, 98-107. GIESE, P. 1978. Die Krustenstruktur des Varistikums und das Problem der Krustenverkiirzung. Z. dt. geol. Ges. 129, 513-20.

GLOCKE, A. & MEISSNER, R. 1976. Near-vertical reflections recorded at the wide-angle profile in the Rhenish Massif. In: G[ESE, P., PRODEHL, C. & STEIN, A. (eds) Explosion Seismology in Central Europe--data and results, 252-6. SpringerVerlag, Berlin. HERRMANN, A. G. • WEDEPOHL, K. H. 1970. Untersuchungen an spilitischen Gesteinen der variskischen Geosynkline in Nordwestdeutschland. Contr. Miner. Petrol. 19, 255-74. KENOLTY, M., CHADWICK,R. A., BLUNDELL,D. J. & BACON, M. 1981. Deep seismic reflection survey across the Variscan Front of southern England. Nature, 262, 374-7. KUSHIRO, I. 1981. Viscosity, density, and structure of silicate melts at high pressures, and their petrological application. In: HARGRAVE (ed.) Physics of Magmatic Processes, Princeton University Press. KUZNIR, N. J. & BOTT, M. H. P. 1977. Stress concentration in the upper lithosphere caused by underlying viscoelastic creep. Tectonophys. 43, 247-56. LORENZ, V. 1976. Formation of Hercynian subplates, possible causes and consequences. Nature, 262, 374-7. MEISSNER, R. 1967. Zum Aufbau der Erdkruste. Beitr. Geophys. 76, 211-54, 295-314. 1981. Passive margin development. A consequence of specific convection patterns in a variable viscosity upper viscosity upper mantle. Oceanologica Acta, 115-21. Actes 26 e Congr6s International de Geologic des marges continentales, Paris, 7-17 juillet 1980. & VEa~rER, U. 1976. Investigations on isostatic balance in different parts of Eurasia, based on seismic and gravity data. In: GIESE, P., PRODEHL, C. & STEIN, A. (eds) Explosion Seismology in Central Europe--data and results, 396-400. Springer-Verlag, Berlin. & -1979. Relationship between the seismic quality factor Q and the effective viscosity. J. Geophys. 45, 147-58. , BARTELSEN,H. & MURAWSKI,H. 1980. Seismic reflection and refraction studies for investigating fault zones along the Geotraverse Rhenoherzynikum. Tectonophys. 64, 59-84. -

-

-

-

32

R. M e i s s n e r e t al.

- &~ 1981. Thin-skinned tectonics in the northern Rhenish Massif. Nature, 290, 339-401. , SPRINGER, M., MURAWSKI, H., BARTELSEN, H., FLUH, E. R. & DORSCHNER, H. 1983. Combined seismic reflection-refraction investigations in the Rhenish shield and their relation to recent tectonic movements. In: FUCHS, K. & MURAWSKI, H. (eds) Plateau Uplift, 276-87. Springer-Verlag, Berlin. -& STREHLAU, J. 1982. Limits of stresses in continental crusts and their relation to the depth-frequency distribution of shallow earthquakes. Tectonics, 1, 73-89. MURAWSK[, H. 1976. Raumproblem und Bewegungsablauf an listrischen F1/ichen, insbesondere an Tiefenst6rungen. Neues Jb. Geol. Paldont. Mh. 209-20. NUR, A . & BEN-AVRAHAM, Z. 1982. Oceanic plateaus, the fragmentation of continents and mountain building. J. geophys. Res. 87, 3644-61.

OLIVER, J. 1980. Exploring the basement of the North American Continent. Am. Scient. 68, 676-83. - 1982. Probing the structure of the deep continental crust. Science, 216, 689-95. SPRINGER, M. 1982. Auswertung reflexionsseismischer Messungen im Hohen Venn. Diploma Thesis. Institut f/Jr Geophysik, Kiel. 140 pp. TEICHMULLER, M. & TEICHMCrLLER, R. 1979. Ein Inkohlungsprofil entlang der linksrheinischen Geotraverse yon Schleiden nach Aachen und die lnkohlung in der Nord-Siid-Zone der Eifel. Fortschr. Geol. Rheinld Westf 27, 323-55. WEBER, K. 1981. The structural development of the Rheinische Schiefergebirge. In: ZWART, H. J. & DORNSlEPEN, U. F. (eds) The Variscan Orogen in Europe. Geologie Mijnb. 60, 149-59. ZIEGLER, P. A. 1978. North-Western Europe: tectonics and basin development. Geologie Mijnb. 57, 589-626.

R. MErSSNE~, M. SPRINGER & E. FLIAH, Institut ffir Geophysik, Universitfit Kiel, Federal Republic of Germany.

Late events in the tectonic history of the Saxothuringian zone Wolfgang Franke SUMMARY: The Saxothuringian zone lies between the Moldanubian block (largely consolidated in late Precambrian time) to the south and the Rhenohercynian zone to the north. It is characterized by exotic blocks of relatively high-grade metamorphic rocks set among very low-grade Palaeozoic sequences. These 'Zwischengebirge' (Miinchberg, Wildenfels, Frankenberg) were formerly interpreted as metamorphic 'diapirs'. Recent investigations have led to a revival of the nappe concepts previously proposed by Suess, Wurm, Kossmat and others. The Miinchberg complex is a pile of later Proterozoic to early Palaeozoic volcanic and sedimentary rocks, some now at advanced states of metamorphism, in which both stratigraphic sequences and metamorphic grades appear in inverted order. These rocks rest upon a Carboniferous wildflysch, which, in its turn, rests upon an autochthonous Devonian and, locally, a Lower Carboniferous sequence. The flysch material, like the nappes above, was derived from sources in the SE. Special features of the sedimentary facies, the tectonic deformation, and the state of very low-grade metamorphism, combine with the evidence of a well-developed thrust at the base of the wildflysch sequence to suggest that this sequence should be treated as the lowest tectonic unit in the Miinchberg pile of nappes. Tectonic deformation of the Palaeozoic sequence began with the production of tight to isoclinal, recumbent, NW-facing folds, accompanied and outlasted by subhorizontal thrusting. It was at this time that the nappe-like tectonic units (already in their metamorphic state) were emplaced. An F 2 refolding produced open, upright to SE-facing folds. A study of illite crystallinity has indicated the significance of a transverse zone, which was the locus of enhanced heat flow throughout the time of deformation, and has confirmed that a metamorphic inversion was introduced when the (relatively strongly metamorphic) wildflysch was thrust over the autochthonous Devonian and Carboniferous. The Saxothuringian zone shows the closest approach to an alpino-type character found in the northern part of the Variscides. Basin development, deformation and metamorphism are best explained in terms of a model based on horizontal tectonism.

The Saxothuringian zone (Kossmat 1927) lies between the external Rhenohercynian zone and the central Variscan crystalline rocks (Moldanubian zone, see Fig. 1). It has been a controversial area ever since geologists began, more than 100 years ago, to give attention to its problems. It is an area in which Palaeozoic rocks of a variety of sedimentary facies and at a variety of metamorphic g r a d e s o c c u r in intimate spatial association. But the chief stimulants to discussion have always been the exotic masses of gneissic rocks near Mfinchberg (West Germany), Wildenfels and Frankenberg (East Germany) and in the Eulengebirge (the Sowie Gory in the West Sudeten, Poland; see the black patches in Fig. 1). Earlier in this century, these aberrant entities were interpreted as remnants of nappes (Suess 1912; Kossmat 1927; Wurm 1928; and others). Later, many preferred to regard these gneisses as high-level representatives of an old basement, around which there had been special developments of facies in the Palaeozoic sedimentary rocks, and which had behaved in a

diapir-like fashion, squeezing their way upward during Variscan deformation. Recent work in the n e i g h b o u r h o o d of the M/.inchberg massif has produced a case for rehabilitating the idea of nappes (Behr et al. 1982: see there a review of the earlier literature). The gneisses of the Miinchberg massif are contained in a late synformal structure (see map, Fig. 2). To the south there is an antiform in the region of the Fichtelgebirge which has rocks ranging in age from Precambrian to Ordovician (and with Silurian and Devonian too along its north-western flank), all of them invaded by post-tectonic granites (Richter & Stettner 1979). The Palaeozoic rocks of the Fichtelgebirge have been affected by greenschist, and locally amphibolite facies, regional metamorphism (Mielke et al. 1979). To the NW and NE of the Mfinchberg complex, Palaeozoic rocks (Ordovician to Carboniferous) occur in relatively wide areas of outcrop. These are, in contrast to the rocks of the Fichtelgebirge, weakly metamorphic (only locally in greenschist 33

W. Franke

34

FIG. 1. Structural map of the European Variscides, during the time of synorogenic sedimentation (Devonian and Carboniferous). Exotic Saxothuringian massifs: MM (Mfinchberg), WM (Wildenfels), FM (Frankenberg), SG (Sowie Gory); after Franke & Engel 1983. facies). The Mfinchberg gneisses and the Palaeozoic rocks around them lie in a dishshaped structure. Within this configuration, the following units may be distinguished in order from top to bottom (which is also the order from the centre outward, see Fig. 2 and Table 1): Gneisses: in the central part of the complex there is a group of predominantly basic metamorphic rocks (the 'Hangendserie' or upper series), around whose margins one finds paragneisses and acidic orthogneisses (the 'Liegendserie' or lower series. 'Hangendserie' and 'Liegendserie' are probably separate tectonic units. On the evidence of radiometric dating, the protoliths of the paragneisses were laid down in a range of time from late Precambrian

to Cambro-Ordovician (S611ner, et al. 1981). Gabbroic or basaltic rocks, now in the eclogite facies, were first introduced in the Cambrian (525 Ma: Gebauer & Gr/inenfelder 1979). All of the radiometric data now available consistently suggest an age of metamorphism for the gneisses at around 380 Ma (Devonian). Greenschist facies rocks: the gneisses are bounded on the NE, SE and SW by outcrops of a 'Randamphibolit' (marginal amphibolite) and of a 'Phyllit-Prasinit-Serie'. This latter unit includes bimodal volcanics (basalts, keratophyres and related pyroclastic rocks) and sedimentary rocks (phyllites, sandstones and a few carbonates). The phyllite-prasinite-series is taken to be Ordovician on the basis of lithological comparisons with dated (fossiliferous)

TABLE 1. Sequence and composition o f the tectonic units in the Mfinchberg pile of nappes. Age and lithology o f the protoliths are given in brackets. Horizontal lines mark tectonic boundaries. "Hangend-Serie" banded hornblende-gneiss, amphibolite, sediments and volcanics, gabbro)

paragneiss,

tectonic insertions of eclogite

(early Palaeozoic

"Liegend-Serie" paragneiss, including granulitic varieties (U.Precambrian or early Palaeozoic) and orthogneiss (Ordovician ~ranite) tectonic insertions of serpentinite "Rand-A~phibolit" & "Phyllit-Prasinit-Serie" (greenschist facies) bimodal metavolcanics and metasediments (probably all early Palaeozoic); insertions of serpentinite Allochthonous Palaeozoic rocks ("Bavarian Facies")

"Randschiefer" bimodal volcanics and sediments, with Ordovician fossils Silurian and Devonian radiolarian cherts Carboniferous clastics

Autochthonous Palaeozoic rocks

(wildfl~sch)

("Thuringian Facies")

Tectonic history o f the S a x o t h u r i n g i a n z o n e Ordovician in the unit below. The greenschist facies rocks (and, locally, the 'Liegendserie') also contain tectonic insertions of serpentinite. Fossiliferous Palaeozoic rocks in 'Bavarian' facies: around the outer limits of the outcrop of the metamorphic rocks (and, by interpretation, lower in the pile of nappes) volcanics and fossiliferous sedimentary rocks appear. These belong to a facies that is quite clearly distinct from both the Fichtelgebirge Palaeozoic and the other, relatively widespread, Palaeozoic rocks in the areas to the NW and the NE of the Mfinchberg Complex. It has always been the practice to refer to this special local development as the 'Bavarian' facies and to set it in contrast to the 'Thuringian' facies in the areas around. Adjacent to the margin of outcrop of greenschist facies rocks (and, by interpretation, below them) there is found first the 'Randschiefer-Serie' (marginal slates). These are the metamorphosed equivalents of the greenschist facies rocks above, and they have produced Ordovician fossils. Geochemical studies of the volcanic rocks in the Randschiefer (Wirth 1978) have revealed that the basalts are essentially alkali basalts (and also continental tholeiites) and that the keratophyres, which are trachytic in composition, are differentiates from an alkali basaltic parent magma. Below the Randschiefer, there is a discontinuous belt of outcrop of radiolarian cherts of Silurian and Devonian age, which form tectonic insertions between the Randschiefer above and the Lower Carboniferous below. The Bavarian Lower Carboniferous outcrops relatively widely on the NW and NE of the Mtinchberg and is the lowest unit in the Miinchberg synform. It has usually (and also in Behr et a[. 1983) been interpreted as following in normal upward sequence from the Thuringian rocks below. The Bavarian Carboniferous differs from the Thuringian principally in containing frequent developments of conglomeratic rocks; in its content of lensoid bodies of Lower Carboniferous shallow marine limestones (often identified as 'Kohlenkalk', or Carboniferous Limestone); and in the innumerable phacoids of preCarboniferous rocks it contains. These lensoid bodies of rock are in most cases treated in the earlier literature as upthrust slices within a 'Frankenw/ilder Schuppenzone' (Frankenwald imbricate zone). An exhaustive comparison of the Bavarian and the Thuringian facies would go well beyond the scope of this present article. Details are to be found in thorough examinations of the subject by Wurm (1961), von Gaertner et al. (1968, in English) and, for the Lower Carboniferous, in Gandl & Mansourian (1978).

35

One can say, as a first approximation, that the Bavarian facies was deposited in what were throughout deeper water conditions. This is especially clear in the Devonian, where the Bavarian regime is almost entirely represented by radiolarian cherts. Pelagic sediments do, of course, also appear in the Thuringian facies (cephalopod limestones, shales, some sandstone turbidites) and the basaltic rocks in the Adorf Stufe (low Upper Devonian) might be thought to fit well enough with any suggestion of deep basinal conditions. However, the development of embryonic coral/ stromatoporoid reefs on some of the topographic highs produced by the volcanism shows that the water depth was by no means abyssal. And in any case, there is available Wirth's (1978) report that the basalts in the Thuringian Devonian are continental tholeiites and alkali basalts. In the Lower Carboniferous in Thuringian facies, greywackes, shales, and occasional conglomerates are present. It has been understood for some time (Lambelet et al. 1967) that these include turbidites. In the Bavarian Lower Carboniferous, with its conglomerates and exotic blocks, turbiditic characteristics and the presence of slumped masses had been recognized even earlier (Greiling 1966). Many of the clastic accumulates, and the Carboniferous Limestone bodies, have, however, been interpreted as shallow water deposits and have been attributed to sedimentation in shallow marine milieux around an early Carboniferous island taken to be represented by the Mfinchberg gneissic rocks. According to Behr et al. (1982), the clastics in the Bavarian Lower Carboniferous can be attributed to a deep-sea fan. The exotic blocks of Carboniferous Limestone and of older rocks are regarded as sedimentary klippen. The Bavarian Carboniferous is an example of a wildflysch which was fed with material derived from a pile of nappes in the process of emplacement during the time of sedimentation and in which rocks of Bavarian facies had been caught up. The clastics in the Thuringian Carboniferous are relatively distal and represent an external region of the original flysch basin. In the opinion of Behr et al. (1982) the following stand as arguments in favour of the fact that the Miinchberg complex is allochthonous: (i) The presence of a variety of rocks of early Palaeozoic age, at drastically different metamorphic grades (eclogite to very low grade), now encountered in close spatial association with one another and in inverse sequence.

36

W. Franke

FIG. 2. Geological map of the region around the Miinchberg gneiss complex.

Tectonic history o f the Saxothuringian zone (ii) The powerful development of fabrics (metamorphic and mylonitic foliations: Vollbrecht 1981; Behr et al. 1982) in the Mfinchberg crystalline rocks, whose disposition was essentially horizontal before development of the Mfinchberg synform. (iii) The existence of a deep-sea clastic fan, much better explained as a trench accumulate than as a record of deposition around a small, autochthonous island of gneissic rocks. Wurm (1928) has already suggested the likely site of the root-zone of the nappes to have been at the 'Erbendorf Line', which is the boundary between the Fichtelgebirge and the crystalline rocks of the Moldanubian zone on its southern side. Slices of ultrabasic rocks occur along the 'Erbendorf Line'. The suggestion that a pile of nappes may have overridden the Fichtelgebirge as they were transported northwestward may be taken to be in good accord with the relatively high grade of regional metamorphism (see above) in that overridden region. The re-interpretation offered by Behr et al. (1982) did, however, leave open two particular problems: (i) The relationship of the Bavarian Carboniferous to other rocks around: the Bavarian Carboniferous seems to follow the Thuringian Devonian in normal upward succession, yet at its NW limit (see Fig. 2) it is thrust over Thuringian Carboniferous and must therefore have been to some extent involved in the process of nappe-transport. (ii) The widespread occurrence in the Palaeozoic rocks of folds facing SE and of thrusts overriding in that direction. These have always been seen to be a considerable hindrance to any proposal of nappes driving north-westward. In what follows, deformation and low-grade metamorphism in the Bavarian and Thuringian Carboniferous are briefly described. It is suggested that this evidence can, after all, be accommodated in a proposal of major nappe development. Deformation

Though the tectonic history of the SaxothuringJan zone covers most of the Devonian and Carboniferous, the present study, for a number of reasons, is focused upon the deformation seen in Carboniferous rocks: (i) The older and more metamorphosed rocks which came to be part of the

37

Mfinchberg pile of nappes brought the main features of their deformed state and of their metamorphism with them. It is the Carboniferous rocks, that best reflect the final stages of nappe-emplacement and subsequent, local phases of deformation. (ii) The well-developed anisotropy of the Carboniferous flysch and its wide distribution recommend these rocks as a suitable vehicle for regional comparisons of structure. In the Carboniferous around the Mfinchberg, three main phases of deformation can be distinguished: (i) Originally flat-lying, tight folds with associated cleavage, shear-surfaces and thrusts (all NW facing). (ii) Open, essentially upright folds with some associated thrusts (SE facing) (iii) Various local developments of later fabrics. D a deformation

What is grouped together as the D x deformation is in itself polyphase and shows repeated production of folding, shearing and thrusting. All of the structural surfaces are disposed subparallel to bedding and appear therefore to have been horizontal in their original layout. The major thrusts, i.e. the surfaces on which the nappes travelled and which are now the boundaries between the several tectonic units in the stacked-up Mfinchberg complex--also belong in the D 1 set of structures.

The lower and upper limits" o f the Bavarian Carboniferous In the neighbourhood of the Miinchberg complex the Bavarian Carboniferous is almost everywhere in contact with the basaltic volcanics (Adorf Stufe: low Upper Devonian) of the Thuringian facies-association. It is only in the NW that Bavarian is thrust over Thuringian Carboniferous (see Fig. 2). It is fortunate that the base of the Bavarian Carboniferous is exposed in numerous quarries in the basalts. In every case a strongly developed, beddingparallel thrust is present. The thrust surface in many cases is immediately above a mass of spilites, although it can, elsewhere, lie as much as some tens of metres above the upper surface of the basalts. In these latter cases, one encounters, below the thrust surface, tuffites, tuffaceous greywackes, shales (some with carbonate nodules) and cephalopod limestones which belong in the upper parts of the Adorf Stufe or (occasionally) low in the Famennian. An excep-

38

W. Franke

tional case within one of the tiny Thuringian windows east of the M/inchberg (Fig. 2) shows a sub-thrust sequence which ranges up into black shales of Tournaisian age. Usually, however, the Famennian and some unknown part of the Lower Carboniferous are cut out at the thrust. The thickness of the thrust-zone itself is of the order of a few tens of metres and may in some individual cases be as much as a few hundred metres. At the structural base a mylonite a few centimetres thick is present. There follows above this a tectonic m61ange with a shaly matrix containing a mix of shale materials, some identifiably Devonian, some Carboniferous, but often reduced to a state in which the two are not distinguishable one from the other. 'Floating' within such a m61ange are lensoid, sheared-out bodies of more competent rocks (Ordovician quartzites, Devonian limestones and vulcanites, Lower Carboniferous greywackes and conglomerates). These phacoids have maximum dimensions usually in the range from a few centimetres to a few metres (see Fig. 3). In some of the more extensive exposures, however, one can observe that larger isolated masses (dimensions up to a few hundreds of metres) of the Thuringian Devonian from below occur in the thrust zone. By these means repetitions at mappable scale become evident (see Fig. 4). In the area in the NW where there is widespread outcrop of the Bavarian Carboniferous (see Fig. 2) with numerous enclaves of Thuringian facies, such planed-off 'shavings' of Thuringian volcanic and sedimentary rocks can be recognized at distances of a few hundred metres from the basal

thrust. Deformation in the clastic rocks near the thrusts has been dominantly cataclastic with little sign of recrystallization. Strongly sheared Lower Carboniferous follows above the m61ange zone. It contains numerous exotics derived from the Bavarian Palaeozoic, but at this level there are no longer any slices of material derived from the Thuringian Palaeozoic. Slickensides and drag-folds (see Fig. 4) indicate a sense of tectonic transport in which the overriding body is taken northwestward. In the explanation of the official 1:25,000 geological map, the basal m61ange is in some areas treated as a stratigraphic unit ('K/ihbergSchichten'). The contact between the Thuringian Devonian and the Bavarian Carboniferous has, until now, usually been taken to be a deformed stratigraphic contact. The presence of Ordovician quartzites in the 'KfihbergSchichten' as well as the details of tectonic fabrics developed at the boundary, do, however, make it clear that this 'Devonian-Carboniferous boundary' is a tectonic feature. The Thuringian enclaves within the area of outcrop of Bavarian Carboniferous (see Fig. 2) should be i n t e r p r e t e d - - a s was already done in the works of Wurm ( 1 9 2 8 ) ~ a s tectonic windows. The strips of outcrop of Devonian at the N W limit of the Bavarian Carboniferous (see Fig. 2) belong in a strongly thrust-out anticline whose overturned NW limb contains Carboniferous in Thuringian facies. This anticline represents a ramp which takes t h e basal thrust of the Bavarian Carboniferous from its position at the

SW

NE

<'=s""-~

FIG. 3. Basal overthrust, beneath Bavarian Lower Carboniferous, with tectonic m61ange. 1-4: phacoids. 1: lithology unknown (phacoid inaccessible), 2: greywacke (Bavarian Carboniferous), 3: spilite (Thuringian Devonian), 4: flaser limestone (Thuringian Devonian). Heiss quarry east of Stadtsteinach, at the margin of the largest tectonic window exposing Thuringian facies in the area NW of the gneiss-complex (see Fig. 2 for location).

Tectonic history o f the Saxothuringian zone

NW

I

39

~

i

SE

201 rn

/

FIG. 4. Bavarian Carboniferous basal overthrust, with development of tectonic 'planings' from Thuringian facies. Schicker quarry, Bad Berneck, south of the southern extremity of the gneiss-complex.

upper limit of the basaltic rocks to a higher position within the Carboniferous. NW of the anticline axis (near the western limit of Palaeozoic outcrop) Bavarian Carboniferous is in structural contact with the Thuringian Carboniferous below at a horizontal thrust plane (see Figs 2&7). The upper boundary of the Bavarian Carboniferous, too, is at an important thrust-zone with intensive tectonic admixings of Carboniferous with material from the older rocks above. Along the boundary there are numerous bodies, large and small, of radiolarian cherts from the Silurian and Devonian in Bavarian facies. Due to competency contrasts with the surrounding rocks, and to the presence of thin clay seams between the chert beds, these cherts acted as a weak horizon, which predisposed the formation of the important thrust-zone between the wildflysch and the overlying, older units. It is possible to see in map relationships a suggestion that tatters of the higher units are inserted into the outcropping Carboniferous. One occasionally finds within the Carboniferous detached blocks which represent Bavarian facies rocks and therefore are not slices planed off from units structurally below, and which have no association with mudflows or slumps. At the boundaries of such bodies signs of tectonic 'working' are especially well-developed (mylonitization, cataclasis). They are probably

clippings derived from the overlying unit which have been inserted into the flysch below. As regards the question 'olisthostrome versus tectonic m61ange', it can be stated that the 'Frankenwald imbricate zone' owes its chaotic character not only to the frequent occurrence of mudflows and sedimentary klippen, but also to tectonic insertions from the Thuringian rocks below and from the superjacent Bavarian units.

Fabrics developed at shear surfaces In the Carboniferous in the Thuringian facies, fold-structures are usually intact (see above). The Bavarian Carboniferous, in contrast, is intensively sheared. It is only in the cases of a few compact packets of greywacke that beds can be followed for distances exceeding 10 m. Otherwise, the greywackes and conglomerates occur as tectonically defined bodies, rhomboid in cross-section and so orientated that their equatorial planes lie parallel to bedding. These rhomboids have their limits set by the bedding and by structural surfaces (shear surfaces, often disposed parallel to cleavage). It is still often possible to recognize that such rhomboids were derived from one bed of greywacke (see Fig. 5) even though the beds may only be traceable over a few metres. Coherent stratigraphic association of the rhomboids can be observed in ranges of exposed sections from 1 m to tens of metres. Normal bedding

40

W. Franke

attitudes are predominant. This clearly indicates that the lensoid or rhomboid forms are not part of the deformed state of what had earlier been a mudflow, because in such a case the contained bodies of greywacke material would show no consistent stratigraphic relationships to one another. Much of the deformation of the competent units can be attributed to a flattening process with the formation of boudinage and also rhomboids between high angle conjugate shears (Fig. 5). However, further separation of these units has probably been achieved by beddingparallel shears (thrusts). The characteristic rhombic form of these bodies of relatively competent rocks is evident at a wide range of scales: millimetre to centimetre scale rhomboids developed in silt or fine sandstone bandings, decimetre to metre scale rhomboids of thick greywackes observable in single exposures, and at map scale, delimitation of individual structural entities within the flysch. All these shear surfaces are, as a general rule, younger than the folding and cleavage. They are a tectonic feature characteristic of the Bavarian facies. Their separation varies according to the material involved from a few millimetres to the decimetre scale. The cumulative movement on the total of such surfaces must be significantly large. The strong shearing that is so characteristic of the Bavarian Carboniferous is almost immediately no longer in evidence in the rocks below

the basal thrust. In the Thuringian facies shear surfaces at a low angle to or parallel to bedding are, by comparison, very uncommon and are restricted to particular horizons. It should, however, be noted that the volcanosedimentary sequence of the Ordovician 'Randschiefer' shows strong shearing comparable with that found in the underlying Bavarian Carboniferous.

Folding and cleavage Fold closures are relatively rarely seen in the Bavarian Carboniferous. They are NW facing close or isoclinal folds whose hinges are isolated by shear surfaces and whose axial surfaces are disposed more or less parallel to bedding. Inverted limbs are, as a rule, very short. These folds were probably generated as drag-folds at thrust surfaces. The present layout of the axial surfaces of these folds is controlled by the shallow, late, dish-shaped structure of the M/inchberg region. The folds in every case face NW. This applies also in the narrow strip of Carboniferous outcrop that intervenes between the Mfinchberg complex and the Fichtelgebirge. The folds there face downward but, nevertheless, clearly north-westward. If there were any substance in earlier suggestions that the Miinchberg gneisses had arisen in a diapir-like structure during Variscan deformation, one would expect to find SE-facing in the structures in the Carboniferous

FIG. 5. Greywackes and shales in the Bavarian Carboniferous dissected by conjugate shear surfaces and bedding-parallel thrusts. Arrows indicate stratigraphic upward direction. Railway cutting at B~irenh~iuser, NW of the gneiss-complex.

Tectonic history o f the Saxothuringian zone rocks of this small area. The NW-facing folds are instead interpreted as evidence of movement of the Mfinchberg structural pile northwestwards away from the Fichtelgebirge region. Cleavage is usually only weakly developed. In the Carboniferous, slaty cleavage is confined to the neighbourhood of the enhanced heatflow region of the Frankenwald transverse zone (see below). Even thin (centimetre scale) greywacke beds in the Carboniferous are not cleaved. Slaty cleavage can be seen in some exotic blocks (slices produced by thrusting, or possibly sedimentary klippen?) as well as in most of the Ordovician 'Randschiefer' which lies structurally above the Bavarian Carboniferous. These fabrics must have been produced before emplacement of the exotic blocks and the 'Randschiefer' tectonic unit. Occasionally the cleavage in the Carboniferous cuts both limbs of a fold. Such folds are therefore pre-cleavage folds. On very rare occasions one finds folds which exhibit the structural style of D 1 folds, but face south-eastwards (e.g. below the thrust in Fig. 5). Such cases are probably best explained by the formation of a NWfacing D 1 fold in beds which had previously been inverted either by slumping or during an earlier stage of the D 1 deformation. There are also rare examples of F 1 folds of quartz veins whose limbs are parallel to thrust-surfaces. All these associations of structures suggest that folding, cleavage and thrusting belonging to the first phase of deformation have each been (possibly on several occasions) repeated, and that the shearing outlasted any production of folding and cleavage. It is possible to find quite frequent examples of tight to isoclinal, in some cases flat-lying folds, in the Thuringian Carboniferous NW of the overthrust Bavarian facies. Further NW from there the folds become more open and tend to be more upright. In the Carboniferous to the NW of the Berga anticline (see Fig. 2) there is, so far, no clear evidence of D 1 structures. The F1 folding probably dies out toward the more external (north-western) part of the flysch basin. In some cases the F 1 fold axes are rotated counter-clockwise away from their normal ( N E - S W ) strike. The normal north-wegtward facing then becomes W- or even SW-facing. Often it is only certain individual folds within one exposure that are so rotated. The rotation can take place without any immediate association with fault-surfaces. Cases of the kind can be found in the Bavarian Devonian and Carboniferous, but are also frequent in the Devonian and Ordovician in Thuringian facies.

41

D 2 a n d later d e f o r m a t i o n s

A second phase of folding deforms the horizontally disposed fabrics produced during the first generation of folding. The D 2 deformation produced in the Carboniferous rocks a set of relatively open, upright to south-facing foldstructures. The second cleavage is a spaced cleavage or crenulation cleavage, and is in general more weakly developed than the first cleavage. The superposition of fabrics is clearly documented in refolded folds (Fig. 6), the presence of two cleavages, and also by the fact that open, south-facing folds deform bedding previously inverted. F 2 folds have not been observed in the Palaeozoic rocks adjacent to the NE margin of the gneiss complex. It appears to be the case that, during the F 2 compression, this area lay in the pressure shadow below a north-eastern continuation of the rigid gneissic rocks, which has since been eroded. During the F 2 deformation the Miinchberg synform was, if not fully formed, then at least first clearly defined. Further uplift of the adjacent (on the south) Fichtelgebirge region probably took place during emplacement of the late granites (Upper Carboniferous and Permian). The strong 'tectonic condensation' of the stratigraphy and of the metamorphic sequence in the area intervening between the M6nchberg complex and the Fichtelgebirge is not due to any process of folding, but was brought about by the operation of synthetic normal faults on the NW flank of the rising Fichtelgebirge. These dip more steeply ( 7 0 - 9 0 ~) than the bedding and the F 1 planar fabrics (ca. 1 0 - 5 0 ~ Fig. 7). F 2 effects account for some of the windows of Thuringian facies that emerge (in F 2 anticlines) within the NW part of the area of outcrop of Bavarian facies (see Fig. 2). The shallower dip of the basal overthrust at the NW limit of the Bavarian Carboniferous can also be attributed to F 2. In addition, it is probable that the Berga anticline (the larger part of whose outcrop is in the D D R and is therefore, unfortunately, inaccessible) is a major F 2 structure. Post-D 2 structures are only of local importance. A wide-spaced, subhorizontal crenulation cleavage can probably be attributed to gravitational collapse under the load imposed by the superjacent rocks (cf. for example, Weber 1978). There are, in addition, a variety of kink-band systems, some arranged parallel to strike, others transverse. Because of the generally poor state of exposure it has not yet been

W. Franke

42 NW

lm L

~

. ." ,

.

SE

1

FiG. 6.' F 1 fold (NW-facing), transected by S2 cleavage (SE-facing). Minor F 2 fold at lower right. Arrows indicate stratigraphic upward direction. Thuringian Carboniferous, Lamitz valley, north of the Mtinchberg nappe complex.

possible to develop a regional interpretation of these structures.

Very low-grade metamorphism Stratigraphic and metamorphic inversion is generally typical of the Mtinchberg pile of nappes. A study of illite crystallinity has been undertaken in order to determine if different grades of metamorphism may also be distinguished within the fossiliferous Palaeozoic rocks. Since the observed crystallinity of illites is strongly influenced by the original state of the material (composition and earlier metamorphic history of detrital micas), it was decided to mount the exercise on the basis of comparisons of samples taken exclusively from shales in the Lower Carboniferous flysch. Their composition varies within only fairly narrow limits. The samples were crushed, dispersed to the maximum degree possible with H 2 0 2 and the fraction below 2 #m was then separated in an Atterberg cylinder. Textural preparations of the clay fraction were subjected to X-ray analysis. The crystallinity values give the width of the illite (001) peak at half-height, compared with that of a quartz standard. Lower value readings indicate higher metamorphism. Details of the method are available in Weber (1972). A detailed discussion of the distribution of

crystallinity values would go beyond the scope of this paper. Three overlapping effects, however, can be read from the results and deserve to be mentioned: (i) The Frankenwald transverse zone ('Frankenw/ilder Querzone' in Fig. 2), a horst-zone (Schwan 1956), is clearly identifiable as an area of higher metamorphism (crystallinity values often below 200). This is also reflected in the occurrence of greenschist facies conditions in Devonian spilites south of the 'Blintendorfer Kulm' (Fig. 2; Brand 1980). Enhanced development of cleavage (slaty cleavage) in the transverse zone suggests that heating effects were already in progress during the time of deformation. The alignment of small post-kinematic granites along the transverse zone (Fig. 2) suggests continued heating. The Frankenwald transverse zone appears to be the site of a deepreaching fracture zone, which was associated with higher heat flux during and after the main deformation. For the discussion of further anomalies of crystallinity (see below), readings taken from the neighbourhood of the Frankenwald transverse zone are excluded. (ii) Areal variation of crystatlinity within the outcrop of the Thuringian Car-

Tectonic history o f the Saxothuringian zone NW

Teuschnitz S.

-.-........ , 9-. _

4

"......'" _ . - "

Berga A. Naila S.

43 SE

M0nchberg Massif

o~o ~

'"',,-i.,.,/:"".-, / / ~

:;;.:

",...

++++++

L. Carboniferous in: h-Bavarian F. (230) "~1 (280) Thuringian Facies -~

IBFI

I

(210)

10 kmj

FIG. 7. Diagrammatic cross-section through the Miinchberg pile of nappes and the autochthonous Palaeozoic units toward the NW, close to the SW margin of the Palaeozoic outcrop. The figures give the mean values of illite crystallinity in the different structural domains. HS: 'Hangend-Serie' gneisses, LS: 'Liegend-Serie' gneisses. Ornaments as in Fig. 2. 2 x vertical exaggeration. boniferous are correlated with the tectonic structures (see Figs 2 & 7): the Teuschnitz syncline, which contains the youngest beds, exhibits a clearly lower grade of metamorphism (mean of 300) than the surrounding areas. This is probably due to a smaller stratigraphic overburden. (iii) The mean value of results from Bavarian shales (approximately 230) is clearly lower than the mean from Thuringian shales (approximately 280). This contrast is even more marked, if one considers not the whole Thuringian area, but only the data from the Naila syncline (Figs 2 & 7), immediately to the NW of the Bavarian Carboniferous (mean crystallinity of approximately 290). This change in the crystallinity values takes place immediately at the basal overthrust, on which the Bavarian Carboniferous has overridden the Thuringian. The state of illite crystallinity is therefore pre-tectonic in relation to the overthrusting, and the metamorphic sequence is observed to be inverted.

Conclusions Several lines of evidence make it clear that the Bavarian Carboniferous does not follow the Thuringian Devonian in normal upward stratigraphic progression: (i) The existence of an important overthrust with mylonites and a tectonic mrlange at its basal surface. (ii) The clear sedimentological contrast between the Bavarian and the Carboniferous (see above; and also Behr et al. 1983). (iii) The strong contrast in structural style

between the Bavarian and the Thuringian Carboniferous. (iv) The contrast between the two different levels of very low-grade metamorphism (with an evident inversion). The Bavarian Carboniferous can therefore be regarded as the lowest tectonic unit in the Mfinchberg pile of nappes (see Fig. 7). It has brought in a deviant Carboniferous sedimentary facies, a more advanced state of deformation and a measurably higher degree of metamorphism. The original, proper upward successor of the Thuringian Devonian must have been 'tectonically eroded' during the emplacement of the nappes and carried away north-westward to an area from which it has subsequently been removed by erosion. The appearance of a stratigraphic contact between Bavarian Carboniferous above and Thuringian Devonian below has been brought about by the fact that the basal overthrust has in most places exploited the opportunity provided by the mechanical contrast already available at the upper limit of the competent Thuringian spilite rocks. The pervasive shearing of the Bavarian Carboniferous indicates that nappe-transport was not effected entirely along one single (basal overthrust) surface, but involved innumerable subparallel surfaces. The polyphase character of the D 1 deformation is evidence of the progressive deformation of the Bavarian Carboniferous accomplished during a fairly long episode of tectonic transport. The inversion of the very low-grade metamorphic facies fits well in the general picture of the inverted state of affairs in the Mfinchberg tectonic pile: gneisses and amphibolites rest on greenschist facies rocks, these in turn rest on fossiliferous Palaeozoic, and within the Palaeozoic, Ordovician and Silurian overlie Devonian which rests upon Carboniferous. And

44

W. F r a n k e

one can now add: relatively strongly metamorphic Bavarian Carboniferous occurs above relatively weakly metamorphic ThuringJan Carboniferous. Every possible exercise in inversion seems to have been achieved. Yet the remarkable fact is that bedding, as a general rule, is in a normal attitude. This is in keeping with the suggestion made by Behr et al. (1983) that the thrust surfaces are slightly more steeply inclined than the general bedding and the isograds, so that as nappes coming from south to north were stacked up, older rocks overrode younger ones and more metamorphic rocks were placed on top of less metamorphic ones. The lack of any advanced degree of metamorphism and of any strongly developed internal deformation (no slaty cleavage) in the Bavarian flysch may seem at first sight to contradict the suggestion that this unit has the lowest position at the base of the pile of nappes. The flysch, however, probably still had quite a high water content at the time of the D 1 deformation, and, hence, high hydrostatic pressure, but relatively little deviatoric stress. While it impeded the formation of cleavage and folds, the high pore-water pressure would have done much to ease the transport of the overriding pile of nappes. The Mfinchberg complex, as at present exposed, measures ca. 15 km across the strike. The estimated extent in the same dimension, now making an allowance for post-nappe shortening (the dish-shaped structure) would be of the order of 20 km, and in that pile there would be early Palaeozoic rocks in three different metamorphic states piled on top of one another, and Bavarian Carboniferous at the base of the pile. 'Untelescoping' the pile would lead to a horizontal array laid out over a distance of some 80 kin. There should be added to this a

measure of the distance--approximately 60 kin, which would become 80 km when one makes an allowance to account for post-nappe u p d o m i n g - - f r o m the northern margin of the metamorphic nappes to the putative root-zone at the E r b e n d o r f line south of the Fichtelgebirge. The calculation, at this stage, takes no account of the shortening implied by the insertions of smaller tectonic entities such as the strongly imbricate zone with Devonian cherts which intervenes between the Carboniferous and the Ordovician above, nor of the total shortening involved in the innumerable smaller thrust-surfaces within the nappes. Taken altogether, a transposition of the order of 200 km may apply in the case of the highest tectonic unit (the Mfinchberg gneisses). Significantly higher estimates would not be at all impossible. Though the structural style of the Miinchberg area would fit well in a plate-tectonic concept involving subduction of oceanic crust, the basaltic rocks present have so far produced no indication at all of oceanic affinities. Remnants of oceanic crust might be seen in serpentinite bodies that are caught up along the boundary surfaces of the metamorphic nappes, but their chemical affinities have not yet been examined. It seems also possible, however, that a Saxothuringian ocean did not exist, and that the Saxothuringian zone is an ensialic structure, produced by subduction of continental lithosphere (a process proposed, e.g. by Behr 1978 and Weber 1978). ACKNOWLEDGMENTS: I gratefully acknowledge the invitation to attend the Tectonic Studies Group meeting in Dublin. I am greatly indebted to Dr S. C. Matthews (Uppsala), who (once again) contrived an English rendering of my proposals, and to Dr D. Hutton (Dublin) and two anonymous referees for critical comments.

References BEHR, H.-J. 1978. Subfluenzprozesse im Grundgebirgsstockwerk Mitteluropas. Z. dt. geol. Ges. 129, 291-326. , ENGEL, W. & FRANKE, W. 1982. Variscan wildflysch and nappe-tectonics in the Saxothuringian Zone (NE Bavaria, W. Germany). Am. J. Sci. 282, 1438-70. BRAND, R. 1980. Die niedriggradige Metamorphose einer Diabas-Assoziation im Gebiet yon Berg/Frankenwald. Neues Jb. Miner. Geol. Paliiont. Abh. 139, 1, 82-101. FRANKE, W. & ENCmL, W. 1983. Variscan sedimentary basins on the continent, and relations with SW England. Proc. Ussher Soc. 5, 258-69. VON GAERTNER, H. R., HORSTIG, G. v., STETrNER,

G. & WURM, A. (with contributions from GANDL, J., GREILING,L., LUDWIG,V., RASCHKA, H. & SDZUY,K.) 1968. Saxothuringian in Bavaria. In: Excursions to the Bavarian Margin o f the Bohemian Massif. 1-160.23rd Int. Geol. Congr.,

Prague, 1968, excursion C 34. GANDL, J. & MANSOURIAN,E. 1978. Neue Daten zur Entwicklung des Unterkarbons im Frankenwald (NE Bayern). Z. dr. geol. Ges. 129, 99-108. GEBAUER, D. & GRUNENFELDER,M. 1979. U-Pb zircon and Rb-Sr mineral dating of eclogites and their country rocks. Example: Mfinchberg Gneiss Massif, Northeast Bavaria. Earth planet. Sci. Lett. 42, 35-44. GREIUNG, L. 1966. Sedimentation und Tektonik im

T e c t o n i c h i s t o r y o f the S a x o t h u r i n g i a n z o n e

Pal/iozoikum des Frankenwaldes. Erlanger geol. Abh. 63, 60 pp. KOSSMAT, F. 1927. Gliederung des varistischen Gebirgsbaues. Abh. siichs, geol. Landesamts, 1, 39 pp. LAMBELET, E., MROZEK, H. & SAMTLEBEN,C. 1967. Petrographie, Fazies und Sedimentationsverh/iltnisse des Kulms an der Sfidost-Flanke der Teuschnitzer Mulde (Frankenwald). Mitt. geol. Stlnst. Harnb. 36, 131-68. MIELKE, H., BLUMEt., P. & LANGER, K. 1979. Regional low-pressure metamorphism of low and medium grade in meta-pelites and -psammites of the Fichtelgebirge area, NE Bavaria. Neues Jb. Miner. Geol. Paliiont. Abh. 137, 83-112. RICHTER, P. & STETTNER, G. 1979. Geochemische und petrographische Untersuchungen der Fichtelgebirgsgranite. Geologica Bav. 78, 144 pp. SCHWAN, W. 1956. Die Frankenwfilder Querzone. Abh. dt. Akad. Wiss. 1954, 6, 80 pp. SOLLNER, F., KOHLER, H. & MULLER-SOHNIUS, D. 1981. Rb/Sr-Altersbestimmungen an Gesteinen der Mfinchberger Gneismasse (MM), NEB a y e r n - T e i l 1, Gesamtgesteinsdatierungen. Neues Jb. Miner. Abh. 141, 1, 90-112. SOLLNER, F., KOHLER, H. & MIJLLER-SOHNIUS, D. 1981. Rb/Sr-Altersbestimmungen an Gesteinen

45

der Mfinchberge Gneismasse (MM), NEB a y e r n - T e i l 2, Mineraldatierungen. Neues Jb. Miner. Abh. 142, 178-98. StJESS, F. E. 1912. Vorl/iufige Mitteilung fiber die Miinchberger Deckschollel Sber. Akad. Wiss. Wien, 121, IIa, 253 pp. VOLLBRECHT, A. 1981. Tektogenetische Entwicklung der Mfinchberger Gneismasse (Quarzkorngefiige-Untersuchungen und Mikrothermometrie an Flfissigkeitseinschltissen). GOttinger Arb. Geol. Paliiont. 24, 122 pp. WEBER, K. 1972. Notes on the determination of illite crystallinity. Neues Jb. Miner. Geol. Paliiont. 6, 267-76. 1978. Das Bewegungsbild im Rhenohercynikum-- Abbild einer variszischen Subfluenz. Z. dt. geol. Ges. 129, 249-81. WIRTH, R. 1978. Geochemie und Petrographie der pal/iozoischen Magmatite des Frankenwaldes. Diabase-Keratophyre-Pikrite. Unpublished Dissertation. University of Wiirzburg. 130 pp. WURM, A. 1928. Der Bauplan des variskischen Gebirges in Bayern. Ein Beitrag zum Bewegungsmechanismus des variskischen Gebirges. Neues Jb. Miner. Geol. Paliiont. 60, B, 473-530. 1961. Geologie von Bayern. Gebr. Borntr/iger, Berlin. 555 pp.

WOLFGANG FRANKE, Geologisch-Palfiontologisches lnstitut, Goldschmidt-Str. 3, D-34 G6ttingen, Federal Republic of Germany.

Inverted metamorphic zonation and large-scale thrusting in the Variscan Belt: an example in the French Massif Central J. P. Burg, A. Leyreloup, J. Marchand & Ph. Matte SUMMARY: Half of the northern part of the French Massif Central has been shown to be allochthonous. In the allochthonous rocks, preserved high-pressure granulites represent an early tectonometamorphic event which took place before thrusting. During the southward-directed overthrusting of this high-grade metamorphic slab a plurifacial inverted metamorphism was developed in the underlying sediments. Adjacent areas show different types of metamorphism in the autochthonous and parautochthonous pelites. The climax of the metamorphism is shown to have been reached by the end of the thrusting episode. This was followed by the retrogression of the high-grade rocks and the establishment of a low-pressure metamorphism in both the thrust unit and the autochthonousparautochthonous terranes beneath. We describe this example of inverted metamorphic zonation associated with the large-scale thrusting of an already metamorphosed series.

Two principal domains with different lithostratigraphic, tectonic, metamorphic and mag-~ matic characteristics have been distinguished in the eastern French Massif Central (Burg & Matte 1978) and appear also to be present in the western part of this massif. The southern part of the Massif Central constitutes the autochthonous and parautochthonous domain. The allochthonous domain, which comprises half of the northern part of the Massif Central (Fig. 1) represents the deepest structural level of this Variscan segment (it includes granulite facies rocks, Forestier et al. 1973) and it has overriden the lower grade domain by about 150 kin. In the north, the contact between the two domains is a lithologic unit composed of blastomylonitic gneisses containing sillimanite nodules and, locally, lenses of deformed serpentinites (Burg 1977). In the south the contact zone is represented by mylonites (Faure et al. 1979; Bodinier & Burg 1980-81). From detailed mapping, the isograds of the prograde metamorphism developed in the rocks below the thrust zone appear to be inverted. We review the geology of the Massif Central, describe this inverted metamorphic zonation and discuss the relationships between nappe emplacement and metamorphism.

Lithostratigraphic sequences Autochthonous and parautochthonousdomain This domain comprises three zones separated by thrust faults (Burg & Matte 1978). Whilst the synthetic lithostratigraphical columns for these different areas have been described previously (Burg & Matte 1978, fig. 5), they are summarized below.

In the south (Montagne Noire, southern Cevennes), non-metamorphic or low- to medium-grade rocks are late Precambrian to "Vis6an in age. The details of the fossiliferous series have been described by Thoral (1935) and Geze (1949). In the southern Montagne Noire, the sediments (flysch and limestones) are preserved in the inverted limbs of a pile of large recumbent folds which are overturned to the south (Arthaud 1970). The sequence has suffered a downward increasing metamorphism. The deepest rocks are anatectic gneisses (Schuiling 1960; Demange 1980-81; Thompson & Bard 1982) occurring in a large anticline: the 'Zone axiale' where alkaline orthogneisses are Cambrian (Hamet & Allegre 1972; Ducrot et al. 1979). Syntectonic sedimentation took place during the Vis6an (Engel et al. 1980-81). In the north the geometrically lower rocks appear in the core of large domes (Tulle anticline and in the Rouergue and Velay areas). These are anatectic ortho- and parametamorphites (greywackes and granites) (Demay 1931; Roques 1941; Chenevoy 1964; Collomb 1970) surrounded by a unit similar to the leptynoamphibolitic groups, which are considered to be allochthonous (Floc'h et al. 1977; Nicollet et al. 1979). The leptyno-amphibolitic groups are map units characterized by a bimodal association of mafic and sialic rocks. They include metasediments (greywackes, pyroclastics), felsic rocks have a granitic composition and metamorphosed basic rocks (metagabbros, metabasalts, metadiabases). Some of these rocks have high-pressure mineral assemblages: eclogites (Chenevoy et al. 1969; Collomb 1970; Santallier et al. 1978), and trondjhemites (Nicollet et al. 1979). Both mafic and sialic rocks appear as boudins of variable size (some 47

48

J.P. Burg et al.

FIG. 1. Variscan nappes and thrusts in the Massif Central Frangais. Areas are: (A) Artense, (Ai) Aigurande, (C) C6vennes, (D) Dordogne, (H.A.) Haut-Allier, (L) Limousin, (Ly) Lyonnais, (MN) Montagne Noire, (N) Najac, (R) Rouergue + Albigeois, (S) Sioule, (T) Tulle, (V) Velay, (M) Marvejols.

M e t a m o r p h i c zonation in the M a s s i f Central

tens of centimetres to several kilometres) in paragneisses: their age is unknown. Some of the basic rocks may show tholeiitic affinities (Piboule 1977; Nicollet et al. 1979). If they really are oceanic rocks an important thrust zone should run along the basal contact of these leptyno-amphibolitic outcrops (in the Tulle, Rouergue and Velay areas). However, no clear tectonic evidence has yet been found to confirm such an hypothesis. Above the leptyno-amphibolitic groups, the 'schistes p6riph6riques' (Roques 1971) or 'schistes des C6vennes-Albigeois' form a thick sequence ( > 5000 m) of mainly quartz-rich pelites considered to be essentially lower Palaeozoic. Remains of Cambro-Ordovician acritarchs have been found in the Marvejols area (Baudelot in Briand & Gay 1978) and Ordovician to late Silurian acritarchs and conodonts have been identified in the BasLimousin (Guillot & Doubingez 1971; Guillot & Lefevre 1975). In the eastern and western Massif Central these quartzo-pelitic rocks show increasing metamorphism from south to north. In low-grade rocks the sequence appears as a monotonous alternation of bluish-black slates and greywackes with a few layers of white orthoquartzite and black cherts. Some calcschists are present but true limestones are scarce. Layers of mafic tufts, and in some cases, diabase dykes and pillows (albigeois) have been described (Guillon 1963; Nicolet 1963). Towards the base of the 'schistes p6riph6riques' sequence feldspathic gneisses (porphyroides) are present. They are characterized by porphyroclasts of magmatic blue quartz and albitized microcline. These gneisses are generally interpreted as acid volcanics and/or volcano-sedimentary rocks, but some of them could be granite laccoliths intruded before deformation (Burg & Teyssier 1983). In the Rouergue, they are dated at 514 + 10 Ma (Delbos et al. 1964-65, whole rock Rb/Sr, recalculated with ~l = 1.42x10 -11 yr-1). In the east such 'metarhyolites' are also of Cambrian age ( 5 4 5 -+ 14, whole rock Rb/Sr, CaenVachette 1979). Within the autochthonous and parautochthonous domain three main types of predeformation granitic intrusives can be recognized: (i) Calc-alkaline porphyritic granitic and monzonitic orthogneisses with ages ranging from 406 to 520 Ma (Duthou et al. 1983). (ii) Pink fine grained orthogneisses with Fe-rich biotite. (iii) Fine grained diorites dated at 540 +- 15

49

Ma in the Marvejols area (U/Pb method on zircons, Pin & Lancelot 1978). These ages suggest that a part of the quartzo-pelitic sequence includes upper Precambrian rocks.

Thrust domain (crystalline nappes) The lithostratigraphic succession of the allochthonous rocks may be divided into: (a) a lower, undated, highly metamorphosed part, and (b) an upper low grade or nonmetamorphosed Devono-Carboniferous cover.

The high-grade rocks Two leptyno-amphibolitic groups can be recognized here: (i) The lowermost is similar in composition to that which outcrops in the autochthonous/parautochthonous domain (Tulle, Rouergue and Velay areas). It is a thick series ( > 5000 m) of sillimanite paragneisses (pelites, sandstones and greywackes) containing boudins of amphibolites (5-15% of the formation) in close association with ultramafics (serpentinites) or felsic orthogneisses (derived from porphyritic granites and acid lavas?). The only high-pressure rocks here are represented by eclogites and coronitic gabbros often retrogressed to amphibolite facies. This group is particularly well exposed in the Artense area where low-pressure granulites have been described (Cornen 1980). (ii) The upper leptyno-amphibolitic group is no more than 1000-2000 m thick. The matrix is essentially represented by fel~spathic orthogneisses with some fine grained pelitic gneisses. Twenty per cent of this unit is made up of basic and ultrabasic boudins (amphibolites, eclogites, gabbros, garnet peridotites, Chevenoy et al. 1969; Lasnier 1971, 1977; Coffrant & Piboule 1971; Bonnot & Piboule 1980). Some amphibolites have been shown to have resulted from the retrogression of eclogites and/or mafic high pressure granulitic rocks (pyrigarnites, Lasnier 1977). The main difference between these and the lower leptyno-amphibolites is in the occurrence of sialic granulitic rocks (Marchand 1974). Thin lenses of skarns and marbles with relics of as yet unidentified fossils (algae?) have been found (Forestier et al. 1973).

50

J.P.

B u r g et al.

The age of the rocks present as boudins range from about 550 to about 400 Ma (Pin 1979; Bernard-Griffiths et al. 1980; Duthou et al. 1981; Gebauer et al. 1981; Pin & Lancelot 1982; Suire i982). On the maps, these leptyno-amphibolitic groups as a whole appear fairly continuous. Their boundaries are neither folded nor boudinaged by the major tectonic events and they run parallel to the regional foliation. Moreover, the rocks have affinities with calc-alkaline tholeiitic basic rocks from island arc or back-arc settings (Briand & Piboule 1979) and sialic rocks such as orthogneisses, pelites and marbles. It may be that the leptyno-amphibolitic units were originally a sedimentary m61ange. However we prefer to take the simpler view that they are a tectonic m61ange associated with a major thrust zone which carries these units over lower-grade rocks (Burg & Matte 1978). Shreds of serpentinized harzburgites commonly occur within the thrust zone itself (Burg & Matte 1977) which is characterized by blastomylonitic sillimanite gneisses or typical mylonites (Burg 1977; Faure et al. 1979; Bodinier & Burg 1980-81). This zone has a generally northward-dipping attitude. The serpentinized harzburgites and some related basic rocks have been interpreted as remnants of an ophiolitic suite (Mercier et al. 1982). The anatectic upper sequence is found above the leptyno-amphibolitic groups in large synforms. They are mobilized massive paragneisses generally characterized by large cordierite patches (Chevenoy & Ravier 1971). This sequence includes lenses of high-pressure granulitic material (khondalito-kinzigitic gneisses, eclogites, garnet peridotites). These lenses and relfcts of high-pressure metamorphic minerals lead to the hypothesis that the anatectic gneisses were the result of the melting of a sequence of granulites (Forestier et al. 1973) caused by the under-thrusting of the water-rich pelitic sediments and the subsequent uplifting. The e p & o n a l a n d u n m e t a m o r p h o s e d rocks

These crop out in the northern area. Two units are distinguished. (i) The stratigraphically lowermost epizonal unit consists predominantly of calcalkaline spilitokeratophyric and volcano: sedimentary rocks (Bebien et al. 1977; Pin et al. 1983), considered to be Upper Devonian (Guffroy 1964). There is an unconformity between these epizonal volcanic rocks and the underlying anatexites (Beurrier et al. 1979). (ii) The upper unit is an unmetamorphosed

ignimbritic (Bertaux et al. 1978) and volcanoclastic sequence which unconformably overlies the epizonal sequence (Jung & Raguin 1935). The basal conglomerate contains pebbles of schist and limestone. Limestones which belong to this upper unit provided a mid-Vis6an fauna (Lys, in Echavarri 1966).

Tectonics Definition and extension of the nappes

The prominent tectonic characteristic of the crystalline Massif Central is the presence of a major thrust and nappe regime. This was pointed out for the first time by Demay (1931, 1948) bur the idea was more or less abandoned until recently. The precise location and extension of the Variscan thrusts has only been determined during the last ten years (Burg 1977; Briand & Gay 1978; Burg & Matte 1978). More recent work has resulted in a new structural sketch map (Fig. 1). The location of the main thrusts and nappes in the crystalline Massif Central is based on the following criteria. (i) The superposition of units of different lithologies (see previous section). The overriding slabs consist essentially of the leptyno-amphibolitic groups and the anatexites whereas the underlying series are essentially quartzo-pelitic schists and gneisses (Burg 1977; Burg & Matte 1978). (ii) The superposition of the high-grade rocks over less metamorphosed rocks. The thrust terranes show a plurifacial complex metamorphism and contain boudins of granulites and eclogites. The underlying pelitic series show in turn a more simple prograde metamorphic evolution. (iii) The inversion of the metamorphism in the underlying pelitic series. This feature has been known for a long time (De Launay 1888; Bergeron 1889; Boule 1899-1900; Roques 1941; Demay 1948, 1949) but was first described accurately by Peyretti (1971), Briand (1973a) and Briand & Gay (1978) in the Marvejols area. (iv) The presence in some areas of mylonites (Faure et al. 1979; Bodinier & Burg 1980-81). In these cases the main basal thrust zone is precisely located.

Metamorphic zonation in the Massif Central (v) The presence of shreds of ultramafic rocks (serpentinites) tectonically emplaced near the boundary between the main lithological units (Burg 1977). On the basis of the above criteria, we can distinguish three main units in the eastern Massif Central (Fig. 1).

The Sioule nappe This unit occurs in the north (Grolier 1971). Its small outcrop area is due to the intrusion of the various post-tectonic granitoids between 360 and 320 Ma (Duthou et al. 1983). The minimum amount of overthrusting is 20 km (the observable heave on the map). The Haut-Allier nappe This is the best documented nappe. It extends over 27,000 km 2. The main basal thrust has been folded by gentle synforms and antiforms and the nappe has been eroded as tectonic windows and klippes. The displacement on this thrust may be estimated at around 150 km (Burg & Matte 1977). The Rouergue nappe The existence of a nappe here is controversial because it is suggested only from the superposition of leptyno-amphibolites over anatectic pelitic rocks. Mylonites and inverted metamorphism in the supposed basement have not been observed. In the western Massif Central three units can also be deduced. From their tectonic and metamorphic histories, they appear to be equivalent to the eastern units and the following correlation is proposed: Aigurande Nappe -- Sioule Nappe Haut Limousin N a p p e = Haut Allier Nappe Tulle area = Rouergue area Deformational sequence in the nappes and sense of displacement

Within the nappes The regional foliation and the zones of concentrated ductile strain separating the different tectonic units are folded by large dome and basin structures. They are related to a D 3 deformation (Burg 1977; Burg & Matte 1978). Parasitic mesofolds and minor chevron folds of this generation are common. They show variable orientation but generally face S to SW. Isoclinal folds (F2) are related to the regional foliation. The general N W - S E axial trend within the nappes is parallel to a prominent mineral lineation. These folds, which are

51

developed at all scales, are associated with a crenulation cleavage ($2) which is more or less parallel to a first cleavage ($1) in the F 2 fold limbs and on a regional scale. The D 1 deformation is associated with a widespread synmetamorphic foliation which affects the pre-upper Devonian rocks. The orientation and style of the folds corresponding to this phase are unknown. Isoclinal folds preserved in granulitic boudins might represent an older phase or may be contemporaneous with D a. Anatexis occurred during the D 3 deformation and has often obliterated the microstructures of the D 1 and D 2 phases. Along the base of the nappes lithologically variable mylonites (leptynites, augen gneisses, pelitic gneisses) locally occur in intense narrow zones ( 1 0 - 3 0 m thick) with clear-cut boundaries parallel to the regional foliation of the country rocks and to the internal layering of the mylonites. Within these zones of intense ductile deformation the general N W - S E trend of the regional lineation L 2 gives way to a prominent N - S lineation of mineral streaks, grains and aggregates. This lineation, which is in places parallel to the axes of minor isoclinal folds which deform an earlier foliation, is consistent in orientation along the entire exposed length of t h e Contact zone. Structural and textural criteria indicate that the lineation is D 2 in age. It is interpreted as a tectonic transport azimuth related to the N - S thrusting of the overlying rocks. However shear sense indicators (sigmoidal micas, asymmetric pressure shadows, rotated clasts, shear bands etc.) have not been found: hence the direction of overthrusting is ill-documented. Pseudotachylite veins which formed prior to th~ D 3 folding provide evidence for late N - S movements.

Deformational sequence in the underlying series The underlying pelitic series extends over a very large area (25,000 km2). The most prominent tectonic features is a flat lying or slightly northward-dipping composite $1_2 foliation. This foliation is related to superposed isoclinal folding (F 1 and F2) and it occurs over a rock thickness of 10 km. F 1 folds are rare, tight, isoclinal folds with variable trends. In some areas (Rouergue, Cevennes) the S0/S 1 intersections trend N - S to N E - S W and lie parallel to a prominent stretching lineation which is mainly visible in the orthogneisses (Mattauer & Etchecopar 1977). F 2 folds are generally isoclinal with rounded hinges and are relatively consistent in trend ( E - W to N W - S E ) . They are associated with a

J . P . Burg et al.

52

well-developed crenulation cleavage (52) which folds S 1 but is, in general, nearly parallel to it. The apparent vergence of F 2 is variable (probably indicating major fold structures). However, the vergence is predominantly southward within the low-grade rocks. Close to the thrust zone, it is northward over some tens of metres, and this may be part of an inverted F 2 limb developed during southwards overthrusting. In the areas where F 2 folds are well developed the most prominent lineation is E - W and results from the intersection of $1 and S 2 giving, in some orthogneisses, an apparent E - W elongation. The $1_2 foliation is folded by gentle, steeply inclined, folds trending from N W - S E to E - W over much of the area but N - S in southern Rouergue. The F 3 folding occurred prior to the emplacement of the Margeride granite dated at 323 + 12 Ma (Rb/Sr whole rock, Couturi6 et al. 1979).

north of the different index minerals Bi, Gt, St, And, Ky, Sill and FK took place mainly during the D 2 deformation event (see mineral abbreviations in Table 1). The standard approach of deformation crystallization relationships (e.g. Vernon 1978) shows, however, that temperature and pressure reached peak values during the interphase D 2 - D 3 (Fig. 3). Because of the upright attitude of the structures which developed later than the climax of the metamorphism in the autochthonous and parautochthonous units, we infer that the mapped zonation cannnot be due to any such late tectonic event. Moreover, the metamorphism we have studied is by far the more important one in the region and has suffered little late retrogression which might arguably account for a part of this zonation. Therefore we argue that the isograds present a primary relationship. Taking into account the mean northwam dip of the series, the metamorphic zonation must be inverted and may be ascribed to the southward translation of the Haut-Allier Nappe (Burg & Matte 1977).

Inverted plurifacial metamorphism More than 4000 thin sections from the autochthonous, parautochthonous and allochthonous pelitic rocks have been used to characterize the metamorphism associated with the HautAllier nappe. Mineral assemblages associated with the penetrative $1_2 foliation (related to the nappe displacement) are shown in Fig. 2. No significant angle can be observed between the deduced discontinuous reaction isograds and the traces of the regional foliation. Moreover both are folded by the large F 3 structures which again suggests that the isograd surfaces are parallel or close to the $1_2 planes. This leads us to conclude that the appearance from south to

Mineral assemblages Very low, low and medium-grade rocks outcrop essentially south of the Margeride granite; high grade and a more or less anatectic series outcrop only above the thrust zone or north of the granite. Two areas are thus geographically distinguished.

In the south There is a continuous prograde change from epizone to lower amphibolite facies conditions in good outcrop from S (or SW) to N (or NW) across the area. The successive metamorphic mineral assemblages are:

(A) Chl. Mu. Alb. Gr. Op. Isograd Bi (in)

(B) Chl. Mu. Bi. Alb. Gr. Op. Gt (in) (C)

Chl. Mu. Bi. Gt. PI (An 15-20). Gr. Op. St (in)

(D) Mu. _+ Bi. ___Gt. St +_ Chloid. P1 (An 15-20). Gr. Rut. And (in) (El) Mu. Bi. _ Gt. St. + A n d + Ky. PI. Ru. Op. Ky (in) (FI) Mu. Bi. 2 Gt. St. Ky. PI. Ru. Op. St (out) (G) Mu. Bi. Gt. Ky. PI. Ru. Op. Sill (in) (H) Mu. Bi. +_ Gt +- A n d +- Ky + Sill +_ Cd. P1. Ru. Op.

FIG. 2. Sketch map and section of the Haut-Allier Nappe and associated discontinuous reaction isograds. K: local occurrence of kyanite in the And-St-Gt zones. Some data from Peyretti (1971) Cheze (I975), Joubert (1978), Briand & Gay (1978) and Suite (1982) have been used. Short single dashes with open arrowheads indicate generalized orientation of regional foliation.

J . P . Burg e t al.

54

TABLE 1. Mineral abbreviations Alb : albite And : andalusite Bi : biotite Cd : cordierite Chl : chlorite Chlo~d : choritoi'd FK : K feldspar Gt : garnet Gr : graphite Ky : kyanite

Mu : muscovite Op : opaque Pi : pinite P1 : plagioclase (An20-25) Ru : rutile Ser : sericite Sill : sillimanite Sph : sphene St : staurolite Tou: tourmaline

In the Andalu~ite zone, the recurrence of parageneses C and D without chloritoid can be attributed to bulk rock chemistry. Quartz and tourmaline are represented in all these parageneses. Chloritoid is a good temperature indicator as it appears with staurolite (Briand 1973b) close to the St(in)isograd. The garnet and staurolite zones are generally too thin to be distinguished at the scale of Fig. 2 where only the medium-grade zone is shown; small blades of kyanite have been described (Poulain 1972) and found by us in some places within the And-St-Gt zones (Fig. 2). The Ky zone proper can only be defined just below the thrust zone (Peyretti 1971; Briand 1973a; Suire 1982; this work). This mineral is also in equilibrium within the mylonites of the Marvejols area. There is good textural evidence that sillimanite grew on biotite in the thrust zone and on andalusite within the nappe. Sillimanite is associated with kyanite

along the contact zone (Briand 1973a) and is persistent up to the geometrically uppermost outcrops. Metastable andalusite and kyanite are also observed respectively at the western and eastern ends of the thrusted leptyno-amphibolitic group in the Marvejols area but the occurrences are too small to be shown on Fig. 2. Rutile constitutes solid inclusions in staurolite or garnet or kyanite. The disappearance of staurolite coincides with the thrust zone. Neither clear prograde formations or destabilization of kyanite and staurolite, not coexistent St.-Sill have been observed.

In the north The same mineral assemblages seen in the south are recognized here in the same order along a section from west to east. In the Desges tectonic window, the lowest grade rocks show: (E2) Mu. Bi. Gt. St. And. Sill. PI. Ru. Op. (F2) Mu. Bi. Gt. St. And. Ky. Sill. P1. Ru. Op. (This mineral assemblage represents the kyanite zone.) In both E 2 and F 2 staurolite may form inclusions in garnet. In the Truy6re and Celoux windows highergrade rocks are encountered, which are characterized by the incipient Mu (out) reaction. (L) Mu. Bi. Gt. St. Sill. P1. Ru. Op. (J) Relic Mu. Bi. Gt. Sill. FK. P1. Ru. Op. (K) Bi. Gt. Sill FK • Cd. P1. Op. The mineral assemblage L replaces G which has not been seen in this area.

FIG. 3. Crystallization-deformation relationships. For mineral abbreviations see Table 1.

Metamorphic zonation in the Massif Central The K assemblage is only represented in the mylonites of the Truy6re Valley. Above these rocks Mu. Bi. Gt. Sill +_ Cd. bearing anatexites (H assemblage without Ky and And.) constitute most of the nappe. Sillimanite, which contributes to the formation of the $1_2 foliation, grew on biotite and while sillimanite may have developed on post-D 3 andalusite (Fig. 3), as observed in the south.

Retrograde evolution of the pelitic granulites The pelitic granulites outcrop in the northern Haut-Allier (Marchand 1974) and in the Lyonnais area (Davoine 1975), i.e. in the most internal parts of the nappe. The same granulites have been described in the Sioule Nappe (Gentilhomme 1975). They occur as pods or lenses some metres wide surrounded by cordieritebearing migmatites (H paragenesis) and are associated with the basic and ultrabasic highpressure granulites (Lasnier 1977). Their primary high-pressure paragenesis is (Marchand 1974): (L) Bil.Gt 1. relic St. Ky. mesoperthitic FK.PI.Ru.Gr. This mineral association could not have been generated by the prograde metamorphism of the Haut-Allier region. Textural evidence shows that kyanite formed from staurolite. These rocks have suffered a syn-D 2 retrogression through several steps indicating a decrease of P.T. conditions (Fig. 3). They are:

55

Type of metamorphism and estimation of P.T. conditions during thrusting (Fig. 4) According to Miyashiro (1973), the enumerated mineral assemblages belong to an intermediate low-pressure metamorphism for the lower-grade rocks of the southern region and to an intermediate-pressure type (Barrovian) just below the thrust in the Marvejols area. This can be illustrated by a piezothermic array (Richardson & England 1979) with a variable slope (Fig. 4A). A P.T. array corresponding only to an intermediate low-pressure metamorphism can be traced for the Haut-Allier region, taking into account coexistent St.-Sill in the high grade rocks (Fig. 4B). In decreasing PH20 and increasing Pco 2 the gradient deduced for the southern region could have generated the L granulitic paragenesis which only occur in the north where the proposed metamorphic gradient cannot produce L rocks. This suggests that the change in slope of the piezothermic array is an effect of the thrusting. Below the Haut-Limousin Nappe (which is for us the western equivalent to the Haut-Allier Nappe, see above) Santallier et al. (1978) have described the successive appearance of Bi, Gt, St, Ky, Sill, and KF during the D 1 deformation event. The normal isograds appear here also to be parallel or close to the foliation (Fig. 5) and can be attributed to an inverted metamorphism in the western part of Limousin. This is comparable to the situation in the eastern Massif Ky---> Sill

(M) Bi.2Gt.1 -+ relic Ky. Sill. FK. PI. Ru. Op. Gt.1 ~ Mu. (N) relic Bi.2 + Gt.1 Sill. Cd. FK. P1. Ru. Op. FK. --->Mu. (O) Mu. Bi.2. relic Gtl. Gt2. Sill +_ relic FK. pl. Ru. Op. and a post-D2 evolution: (P) Chl. Mu. Bi.2 relic Gt.l_2. And --- relic Sill.

AI2SiO5 --~ And.

Gtl, Gt2, Bil, Bi2 are successive phases of crystallization. The metamorphic contrast between the allochthonous granulites and the autochthonous-parautochthonous pelites requires that the higher-grade rocks of the former group were metamorphosed (probably somewhere to the north) prior to their emplacement above the underlying sediments, a metamorphic event which accordingly predated the D 2 deformation (Burg 1977). Nowhere in the metamorphic history of the autochthonous cover is there any indication of such an early high-grade event (D 1 or pre-D1).

Central (this work). The noticeable difference is that andalusite and coexistent staurolitesillimanite have not been found in Limousin and thus the piezothermic array is there of intermediate pressure or intermediate highpressure type (Fig. 4A): a type of metamorphism which is consistent with the formation of L rocks.

Late metamorphic evolution The crystallization of syn- to post-D 3 porphyroblasts is a characteristic feature of these

56

J.P.

B u r g et al.

FIG. 5. Schematic map of the Northwestern Massif Central showing the Limousin Nappe and the metamorphic zonation. Ornament as in Fig. 2. Trace of isograds essentially after Santallier et al. (1978) and Autran & Guillot (1975).

FIG. 4. (A, B) Piezo-thermic arrays for pelitic rocks in the Massif Central. Equilibria: (1): Bi. + Mu. (Nitsch & Brown in Winkler 1974). (2): Gt. (in) unpublished Gt. Bi. geothermometer. (3): Chloid. + St. (Richardson 1968). (4) Chl.§ St. (Hoschek 1969). (5) St.+Mu ---> Ky.+Bi. (Hoschek 1969) (6) St.§ Gt.+Ky. (Richardson 1968) (7) Mu (out) (Storre 1972) (8) Polymorphic AI2SiO 5 triple point (Richardson et al. 1969). G = evolution of the granulites reported with P.H20 2 kB. White arrows: late evolution of the metamorphism. =

areas. Porphroblasts of chlorite, muscovite, biotite, garnet, andalusite, cordierite and a very late sillimanite cut SI_ 2 and were thus developed at the end of the tectonometamorphic history (Fig. 3). This second metamorphic stage seems also to be inverted as Chl., Mu. and Bi. crystallization took place in the south and Gt., And. and Sill. porphyroblasts are found below and above the thrust zone in the medium to high-grade rocks in the north. In

the north and south, this second metamorphic event is exclusively of low-pressure type (Fig. 4A, B) and could be related to syn-kinematic intrusions of granitoids in the thrusted series. Pinite is a common retrograde alteration products of staurolite and garnet and kyanite undergoes retrogradation to sericite. Chlorite may overgrow earlier biotite. In the northernmost area late sericite and chlorite flakes have developed (Forestier 1963). The trace of their first appearance cross-cuts the isograds and the thrust contact. These very low to low-grade minerals are attributed to the mantle gneiss dome of Velay in the east (Fig. 1). Contact metamorphism due to the emplacement of the 323 Ma Margeride granite (Couturi6 et al. 1979) gave rise to the alteration of staurolite into andalusite and hercynite. In the very low-grade and low-grade series, andalusite is characteristic of the hornfelsed rocks.

Discussion and summary On a structural sketch map which includes the whole Massif Central a major thrust unit can be

Metamorphic zonation in the Massif Central distinguished: the Haut-Allier/Haut-Limousin Nappe (Fig. 1). In the eastern Massif Central this unit roots in the Monts due Lyonnais and in the Artense area. The leptyno-amphibolitic groups and overlying anatexites of the HautAllier and Marvejols regions are allochthonous. Since only one contact zone can be followed in the field this excludes the possiblity that the Marvejols area is a separate thrust unit. The klippes of Decazeville (Burg & Matte 1977) and possibly Najac (Bodinier & Burg 1980-81) may belong to this large nappe. Microtectonic data show that the thrusting is essentially contemporaneous with the regional D 2 deformation event. Within the internal parts of the nappe pre-D 2 high-pressure, high-temperature granulites and eclogites are preserved (Forestier et al. 1973; Marchand 1974; Lasnier 1977). This implies that the high-pressure sequence formed elsewhere and the rocks were tectonically emplaced into their present position. This high-pressure event is considered to be 400-380 Ma in age (Matte & Burg 1981). During thrusting of the metamorphic nappe a plurifacial inverted metamorphism was developed in the underlying sediments and attained a climax by the end of the thrust emplacement episode. In addition adjacent areas may show different types of metamorphism: (i) In the external parts of the belt, the low-grade rocks have suffered an intermediate low-pressure metamorphism. (ii) Along the thrust zone tightening of the isograds is compatible with an intermediate-presure type metamorphism in the Marvejols area and with an intermediate low pressure metamorphism in the more internal zones (e.g. Haut-Allier). (iii) In the Limousin, metamorphism is of intermediate pressure or intermediate high-pressure type (see above). We have noticed the apparent parallelism between the 81_ 2 foliation, the discontinuous reactions isograds and the thrust zone. An angle should exist, however small, between the isograds and the base of the nappe. We feel that since, for example, sillimanite is present below the thrust in the north, but is absent in the south (Figs 2 & 5) this obliquity may be real. However, a similar cutout could be the result of late cataclasic movements on the thrust in the south, as in Decazeville (Fig. 2), where anatexites directly overlie low-grade rocks (Burg & Matte 1977). The late evolution of this inverted metamorphism shows an increase in tempera-

57

ture conditions (low-pressure type) probably due to the overlying hot pile above less metamorphosed sediments and possibly to syntectonic granite emplacement in the nappe. We assume that after thrusting, volatiles given off during the metamorphism of the underlying series can be expected to escape upward into the overriding material causing retrograde metamorphism of the relic high-pressure granulites and late anatexis. The tectono-metamorphic record in the Massif Central shows that the geological history must involve significant crustal thickening most readily explained by postulating thrust emplacement of a thick slab of rocks: the Haut-Allier/Haut-Limousin Nappe. These rocks were metamorphosed prior to their thrusting over sediments of an autochthonousparautochthonous domain. Additional thrust slices such as the Sioule-Aigurande unit may have been emplaced over the HautAllier/Haut-Limousin Nappe eventually burying it to depths of 25-30 km. Late diaphthoresis is partly due to uplift but can be considered as a very late stage of the continuous geological history which began with this crustal thickening. Burying the base of the lower crust may have initiated its partial melting and the formation by diapirism of the Velay mantle gneiss dome which completed the tectono-metamorphic evolution at about 300 Ma (Rb/Sr whole rock, Caen-Vachette et al. 1982). This model of tectonic thickening of the crust and inverted metamorphism is similar to that proposed for the development of the Caledonides (e.g. Andreasson & Lagerblad 1980), the Appalachians (e.g. Crawford & Mark 1982), the Alps (e.g. Oxburgh & Turcotte 1974) and the Himalayas (e.g. Le Fort 1975). Several quantitative models of overthrusting have been suggested to illustrate the distributions of metamorphic assemblages that might be expected in such tectonic situations (Oxburgh & Turcotte 1974; Bird & Toksoz 1975; Graham & England 1976; Thompson 1981). They can be adapted to the inverted metamorphism observed in the Massif Central. The metamorphic gradient, however, appears here much steeper than those expected from the models. This is essentially because in the Massif Central the thrust slab was metamorphosed, prior to emplacement, to higher temperatures than those used in the model calculations. In conclusion, the distribution of assemblages of metamorphic minerals in the pelitic schists of

58

J.P. Burg et al.

the Massif C e n t r a l w h i c h lie below the H a u t A l l i e r / H a u t - L i m o u s i n N a p p e can be e x p l a i n e d by m e t a m o r p h i s m a c c o m p a n y i n g the s o u t h w a r d e m p l a c e m e n t of thrust slices on to the c o n t i n e n tal shelf of s o u t h e r n France. This a p p e a r s to s u p p o r t g e o d y n a m i c m o d e l s which imply plate tectonics and o b d u c t i o n processes for the form a t i o n of the Variscan Belt of W e s t e r n E u r o p e ( M a t t e & B u r g 1981).

ACKNOWLEDGMENTS: This work has been carried out with the financial support of the B.R.G.M. (detailed mapping of the Massif Central), and of the C.N.R.S. (ATP. Geodynamique no. 4519 and LA. no. 266).

References ANDREASSON, P. G. & LAGERBLAD, B. 1980. Occurrence and significance of inverted metamorphic gradients in the western Scandinavian Caledonides. J. geol. Soc. London 137, 219-30. ARTHAUD, F. 1970. Etude tectonique et microtectonique compar6e de deux domaines hercyniens: les nappes de la Montagne Noire (France) et l'anticlinorium de l'Iglesiente (Sardaigne). Pub. U S T E L A . Sdrie G~ol. Struct. 1, 175 pp. AUTRAN, A. & GUILLOT, P. L. 1975. L'6volution orog6nique et m6tamorphique du Limousin au Pal6ozoi'que (Massif Central frangais). C. r. hebd. Sdanc. Acad. Sci., Paris, 280D, 1649-52. BEBIEN, J., ROCCI, G., FLOYD, P. A., JUTEAU, Th. & SAGON, J. P. 1977. Le volcanisme D6vonodinatien 616ment d6terminant dans las reconstitution due cadre g6otectonique de l'Europe moyenne varisque. Coll. int. C.N.R.S. 'La Chafne varisque d'Europe moyenne et occidentale'. 243, 275-91. BERGERON, J. 1889. Etude g6ologique du massif ancien situ6 au sud du Massif Central. Th6seParis. Ann. Sci. gdol. 22, 362 pp. BERNARD-GRIFFITHS, J., LASNIER, B., MARCHAND, J. V1DAL, PH. 1980. Approche par la m6thode Rb/Sr de l'6tude de granulites acides en HautAllier (Massif Central franqais). 8dine Rdun. Ann. Sci. Terre. Marseille, 41 pp. BERTAUX, J., GAGNY, C. & RUBIELLO, M. F. 1978. Note pr61iminaire sur l'organisation des formations volcaniques et volcano-s6dimentaires du Pal6ozoi'que sup6rieur de la feuille de Roanne au 1/50.000. C. r. somm. Sdanc. Soc. g~ol. Fr. 6, 289-92. BEURRIER, M., PIBOULE, M., CHIRON, J. C. & GAY, M. 1979. Relations de la s6rie de la Br6venne avec celle du Lyonnais. 7dine Rdun. Ann. Sci. Terre. Lyon, 52. BIRD, P. & TOKSOZ, M. N. 1975. Thermal and mechanical model of continent-continent convergence zones. J. geophys. Res. 80, 4405-16. BODINIER, J. L. & BURG, J. P. 1980-81. Evolution m6tamorphique et tectonique des s6ries cristallophylliennes de Rouergue occidental: mise en 6vidence d'un chevauchement dans la r6gion de Najac (Aveyron). Bull. Bur. Rech. G~ol. Min. S6rie 2, Section I (4), 315-39. BONNOT, H. & PIBOULE, M. 1980. Mise en 6vidence d'une dualit6 d'origine des ultrabasites ~t grenat

du Limousin et recherche de la signification des p6ridotites d'origine mantellique dans le Massif Central fran~ais. C. r. hebd. S~anc. Acad. Sci., Paris, 291D, 129-32. BOULE, M. 1899-1900. G6ologie des environs d'Aurillac. Bull. Serv. Carte gdol. Fr. XI, 279-90. BRIAND, B. 1973a. Lithostratigraphie et mOtamorphisme de la sOrie cristallophyllienne de Marvejols. Bull. Bur. Rech. Gdol. Min. 26me sOrie, I(4), 183-98. 1973b. Le chlorito~de de la s6rie de Marvejols (VallOe du Lot). Bull. Soc. fr. Mindr. Cristallogr. 96, 155-7. -& GAY, M. 1978. La s6rie inverse de SaintGeniez-d'Olt: evolution m6tamorphique et structurale. Bull. Bur. Rech. GOol. Min. 3, 167-86. & PIBOULE, M. 1979. Les m6tabasites de la sOrie de Marvejols (Massif Central): T6moins d'un magmatisme thol6i'tique d'arri6re arc cambroordivicien? Bull. Bur. Rech GOol. Min. Section I(2), 131-71. BURG, J. P. 1977. Tectonique et microtectonique des s6ries cristallophylliennes du Haut-Allier et de la Vall6e de la Truy6re. ThOse 3dme cycle. Montpellier. 79 pp. -& MAaTE, Ph. 1977b. La klippe de la Bessenoits (Decazeville, Aveyron). Un nouvel argument en faveur de l'existence d'un chevauchement majeur vers le Sud dans le Massif Central. C. r. Somm. Sdanc. Soc. g~oL Fr. 6, 325-9. -& MAaqE, Ph. 1978. A cross section through the French Massif Central and the scope of its Variscan geodynamic evolution. Z. dt. geol. Ges. 129, 429-60. - & TEYSSIER, Ch. 1983. Contribution ~ l'6tude tectonique et microtectonique des s6ries cristallophylliennes du Rouergue Oriental (Massif Central franqais): La d6formation des laccolithes syntectoniques type Pinet. Bull. Bur. Rech. G~ol. Min. CAEN-VACHErrE, M. 1979. Age Cambrien des rhyolites transform6es en leptynites dans la s6rie m6tamorphique du Pilat (Massif Central frangais). C. r. hebd. S~anc. Acad. Sci., Paris, 289D, 997-1000. , COUTURIE, J. P. & DtDmR, J. 1982. Ages radiom6triques des granites anatectique et tardimigmatique du Velay (Massif Central

Metamorphic

z o n a t i o n in t h e M a s s i f C e n t r a l

fran~ais). C. r. hebd. SOanc. Acad. Sci., Paris, 294, S6rie II, 135-8. CHENEVOY, M., 1964. Pr6cisions nouvelles sur les terrains m&amorphiques du Mont-Pilat (Massif Central) et leur histoire cristallog6nique. Bull. Soc. gOol. Fr. 7, 55-63. , COFERANT, D. & PIBOULE, M. 1969. Horizons 6clogitiques en Limousin, Massif Central franqais. C. r. Acad. Sci., Paris, 268D, 5-8. - & RAVIER, J. 1971. Caract6res g6n6raux des m6tamorphismes du Massif Central. In: J. Jung Symposium: GOologie GOomorphologie et Strucure Profbnde du Massif Central francais, 109-32. Plein Air Service. ClermontFerrand. CHEZE, Y. 1975. Etude g6ologique de la Chataigneraie au Nord d'Entraigues (Aveyron): p6trographie, structure et m6tallog6nie. Th~se 3~me cycle. Clermont-Ferrand. 144 pp. COFFRANT, D. & PIBOULE, M. 1971. Les 6clogites et roches associ6es des massifs basiques de SaintJoseph (Monts du Lyonnais, Massif Central fran~ais). Bull. Soc. gOol. Fr. 7, 283-91. COLLOMB, P. 1970. Etude g6ologique du Rouergue cristallin. Mem. Serv. Expl. Carte gOol. det. Fr. Paris. 419 pp. CORNEN, G. 1980. P6trologie des gneiss a g6drite d'Ensalers (Massif Central franqais). Bull. MinOr. 103, 478-90. COUTURII~, J. P., VACHETTE-CAEN, M. & VIALETTE, Y. 1979. Age namurien d'un laccolite granitique diff6renci6 par gravit6: le granite de la Margeride (Massif Central fran~ais). C. r. hebd. SOanc. Acad. Sci., Paris, 289D, 449-52. CRAWEORD, M. L. & MARK, L. E. 1982. Evidence from metamorphic rocks for overthrusting. Pennsylvania Piedmont, U.S.A. Can. Min. 2 0 , 333-47. DAVOINE, P. 1975. Leptynites ~ m6soperthite, grenat, disth6ne, spinelle et dumorti6rite dans les Monts du Lyonnais: Etude pr61iminaire. C. r. Somm. SOanc. Soc. COol. Fr. XVII (2,3), 67-9, DE LAUNAY, L. 1888. Compte-rendu de l'excursion du 24 aofit de Chfiteauneuf fi Manzat, au gout de Tazenat, fi Enval et fi Riom. Bull. Soc. gOol. Fr. 16, 1087-93. DELBOS, L., LASSERRE, M. & ROQUES, M. 1964-65. G6ochronologie et r6tromorphose dans la s6rie cristallophyllienne du Rouergue (Massif Central frangais). Sciences Terre, 10 (3-4), 329-42. DEMANGE,, M. 1980-81. Le m6tamorphisme m6sozonal progressif des roches p61itiques sur le flanc nord du massif de l'Agofit. (Montagne Noire). Bull. Bur. Rech. GOol. Min. 26me s6rie, Section 1 (4), 269-91. DEMAY, A. 1931. Les nappes c6venoles. MOrn. Serv. Expl. Carte gOol. dot. Fr. Paris, 320 pp. 1948. Tectonique antestephanienne du Massif Central. MOm Serv. Expl. Carte. gOol. dot. Fr. Paris, 259 pp. - 1949. Sur la tectonique ant6st6phanienne, probablement hercynienne, du Limousin septentrional. C. r. hebd. SOanc. Acad. Sci., Paris, 228, 1599-601. DUCROT, J., LANCELOT, J. R. & RHLLE, J. L. 1979.

59

Datation en Montagne Noire d'un t6moin d'une phase majeure d'amincissement crustal caract6ristique de l'Europe pr6varisque. Bull. Soc. gOol. Fr. 21(4), 501-5. DUTHOU, J. L., PIBOULE, M., GAY, M. & DUFOUR, E. 1981. Datations radiom6triques Rb-Sr sur les orthogranulites des Monts du Lyonnais (Massif Central franqais). C. r. hebd. SOanc. Acad. Sci., Paris, 292, s6rie lI, 749-52. - - . , CANTAGREL, J. M., DIDIER, J. & VIALETTE, Y. 1983. Paleozoic granitoids from the French Massif Central: age and origin studied by 87Rb-sYsr system. In press. ECHAVARRI, A. 1966. Etude p6trographique des tufs anthracif6res et des roches associ6es au Sud de la r6gion de Roanne (Loire). ThOse 3Ome cycle. Paris. 165 pp. ENGEL, W., FEIST, R. & FRANKE, W. 1980-81. Le carbonif6re ante-st6phanien de la Montagne Noire: Rapports entre mise en place des nappes et s6dimentation. Bull. Bur. Rech. GOol. Min. 26me s6rie, Section I (4), 341-89. FAURE, M., PIN, Ch. & MAILHI~, D. 1979. Les roches mylonitiques associ6es au charriage du groupe leptyno-amphibolique sur les schistes du Lot darts la r6gion de Marvejols (Lozere, Massif Central fran~ais). C. r. hebd. Sdanc. Acad. Sci., Paris, 288D, 1267-70. FLOC'H, J. P., SANTALLIER, D., GROLIER, J. & GUlLLOT, P. L. 1977. Donn6es r6centes sur la g6ologie du Bas-Limousin. C. r. 102~me Congr~s. Nat. Soc. Savantes. Fasc. II, 147-58. FORESTmR, F. H. 1963. M6tamorphisme hercynien et ant6hercynien dans le bassin du Haut-Allier (Massif Central fran~ais). Bull. Serv. Carte g~ol. Fr. 271-59. --, LASNIER, B., LEYRELOUP, A. & MARCHAND, J. 1973. Vues nouvelles sur la catazone dans le Massif Central fran~ais, et le Massif Armoricain de l'affieurement au Moho. Bull. Soc. gOol. Fr. 15, 562-77. GEBAUER, D., BERNARD-GRIFFITHS, J. & GRI3NENFELDER, M. 1981. U-Pb zircon and monazite dating of a Mafic-Ultra mafic complex and its country rocks. Example: Sauviat sur Vige, French Central Massif. Contr. Miner. Petrol. 76, 292-300. GENTILHOMME, P. 1975. Leytynites et quartzites de Brouilly-Cesset. Rev. scient. Bourbonnais. 82-101. GEZE, B. 1949. Etude g6ologique de la Montagne Noire et les C6vennes m6ridionales. MOrn. Soc. gOol. Fr. 29-62. GRAHAM, C. H. & ENGLAND, P. C. 1976. Thermal regime and r6gional metamorphism in the vicinity of overthrust faults and example of shear hetting and inverted metamorphic zonation from southern California. Earth planet. Sci. Lett. 31(1), 142-52. GROLmR, J. 1971. Contribution fi l'6tude g6ologique des s6ries cristallophylliennes inverses du Massif Central franqais: La s6rie de la Sioule (PuT de D6me, Allier), M~m. Bur. Rech. GOol. Min. 64, 163 pp. GUFFROV, . . J. 1964. Sur l'existence du Giv6tien dans le

J.P. Burg et al.

60

Morvan C. r. hebd. SOanc. Acad. Sci., Paris, 2619-20. GUmLON, J. H. 1963. Etude g6ologique et m6tallog6nique de l'Albigeois; la r6gion d'AlbanTr6bas (Tam). ThOse 30me cycle. Paris, 74 pp. GUILLOT, P. L. & DOUBINGER, J. 1971. D6couverte d'Acritarches dans les schistes s6riciteux de G6nis (Dordogne). C. r. hebd. SOanc. Acad. Sci., Paris, 272D, 2763-4. & LEFEVRE, J. 1975. D6couverte de conodontes dans le calcaire ~ entroques de G6nis en Dordogne (Bas-Limousin). C. r. hebd. SOanc. Acad. Sci., Paris, 2 8 0 D , 1529-30. HAMET, J. & ALLEGRE, C. J. 1972. Age des orthogneiss de la zone axiale de la Montagne Noire (France) par la m6thode 87Rb-86Sr. Contr. Miner. Petrol. 34, 251-7. HOSCHEK, G. 1969. The stability of staurolite and cloritoid and their significance in metamorphism of pelitic rocks. Contr. Miner. Petrol. 22, 208-32. JOU~ERT, M. 1978. Etude p6trographique, structurale et m6tallog6nique de la Chataigneraie (secteur du Veinazes, Cantal) Massif Central fran~ais. ThOse 30me cycle, Clermont-Ferrand. 206 pp. JUNG, G. & RAGUIN, E. 1935. Discordance du Vis6en sur le socle cristallophyllien entre Balbigny, N6ronde et Violay (Loire). C. r. Somm. SOanc. Soc. gOol. Fr. 16, 247. LASNIER, B. 1971. Les p6ridotites et pyrox6nolites grenat du Bois des Feuilles (Monts du Lyonnais) (France). Contr. Miner. Petrol. 34, 29-42. - 1977. Persistance d'une s6rie granulitique au coeur du Massif Central franqais (Haut-Allier). Les termes basiques, ultrabasiques et carbonat6s. ThOse d'Etat. Nantes. 351 pp. LE FORT, P. 1975. Himalayas: the collided Range. Present knowledge of the continental arc. Am. J. Sci. 275A, 1-45. MARCHAND, J. 1974. Persistance d'une s6rie granulitique au coeur du Massif Central franqais (Haut Allier). Les termes acides. ThOse 30me cycle. Nantes. 267 pp. MATTAUER, M. & ETCHECOPAR, A. 1977. Argument en faveur de chevauchements de type Himalayen dans la chaine hercynienne du Massif Central franqais. Ecologie et GOologie de l'Hirnalaya, Coll. Int. C.N.R.S. 268, 261-7. MAaaE, Ph. & BURG, J. P. 1981. Sutures, thrusts and nappes in the Variscan Arc of western Europe: plate tectonic implications. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 8, Blackwell Scientific Publications, Oxford. MERCIER, J.-C. C., P o z z o DI BORGO, M., FRISON, J. Y. • GIRARDEAU, J. 1982. Les associations basiques et ultrabasiques du Bas-Limousin: restes d'un complexe ophiolitique d6membr6 d'une fraicheur remarquable. 90me Rdun. Ann. Sci. Terre, Paris, 430. MIYASHIRO, A. 1973. Metamorphism and Metamorphic Belt. Allen & Unwin, London. 479 PP. NICOLET, B. 1963. Etude g6ologique et m&allog6nique de l'Albigeois. La r6gion d'Alban-St Jean de Jeannes. ThOse 30me cycle, Paris. 89 pp. 258,

NICOLLET, G., LEYRELOUP, A. & DuPuY, C. 1979. Petrogenesis of high pressure trondhhemitic layers in eclogites and amphibolites from Southern Massif Central, France. In: Trondhjemites, Dacites and Related Rocks. 434-63. Elsevier, New York. OXBURGH, E. R. & TURCOTTE, D. L. 1974. Thermal gradients and regional metamorphism in overthrust terrains with special reference to the Eastern Alps. Schweiz. miner, petrog. Mitt. 54 (2.3), 641-62. PEYRETTI, G. 1971. Etude g6ologique des formations cristallophyliennes ~ l'ouest de Marvejols (Loz6re). ThOse 30me cycle. Lyon. 133 pp. PIBOULE, M. 1977. Utilisation de l'analyse factorielle discriminante pour la reconnaissance de la nature des magmas parents des amphibolites. Application ~ quelques m6tabasites due Rouergue et du Limousin (Massif Central Franqais). Bull. Soc. gdol. Fr. 19, 1133-43. PIN, CH. 1979. Age h 482 Ma des roches orthod6riv6es du groupe leptyno-amphibolique de Marvejols (Loz~re, Massif Central Franqais) d6termin6 par la mEthode U-Pb sur zircons. C. r. hebd. Sdanc. Acad. Sci., Paris, 288D, 291-4. -& LANCELOT, J. R. 1978. Un exemple de magmatisme Cambrien dans le Massif Central: les m6tadiorites quartziques intrusives dans la s6rie du Lot. Bull. Soc. gdol. Fr. 7, 203-8. - & -1982. U-Pb dating of an Early Paleozoic bimodal magmatism in the French Massif Central and of its further metamorphic evolution. Contr. Miner. Petrol. 79, 1-12. --, PETERLONGO, J. M. & DUPUY, C. L. 1983. R6partition des terres rares dans les roches volcaniques basiques d6vonodinantiennes du Nord-Est du Massif Central. Bull. Soc. gOol. Fr. In press. POULAIN, D. 1972. Les micaschistes des environs de Saint Geniez d'Olt (Aveyron). ThOse 30me cycle. Paris VI, 78 pp. RICHARDSON, S. W. 1968. Staurolite stability in a part of the system Fe-AI-Si-O-H. J. Petrol. 9, 468-88. , GILBERT, M. C. & BELL, P. M. 1969. Experimental determination of kyanite-andalusite and andalusite-sillimanite equilibria; the aluminium silicate triple point. Am. J. Sci. 267, 259-72. -& ENGLAND, P. C. 1979. Metamorphic consequences of crustal eclogite production in over thrust orogenic zones. Earth planet. Sci. Lett. 42, 183-90. ROQUES, M. 1941. Les schistes cristallins de la partie Sud-Ouest du massif Central Franqais. Mem. Serv. Expl. Carte gOol. dot. Ft. Paris. 530 pp. - 1971. Structure g6ologique du Massif Central. In: Syrup. J. Jung. GOologie, gOomorphologie et structure pror du .Massif Central Franfais, 17-32. Plein Air Service. Clermont-Ferrand (ed.) SANTALLIER, D., FLOC'H, J. P. & GUILLOT, P. L. 1978. Quelques aspects du m6tamorphisme d~vonien en Bas Limousin (Massif Central Fran~ais). Bull. MinOr. 101, 77-88. SCHUILING, R. 1960. Le d6me gneissique de l'Agofit (Tarn et H6rault). Mem. Soc. gOol. Fr. 91-1, 58 PP.

M e t a m o r p h i c z o n a t i o n in the M a s s i f Central STORRE, B. 1972. Dry melting of muscovite + quartz in the range Ps = 7 Kb to Ps = 20 Kb. Contr. Miner. Petrol. 37, 87-89. SUmE, J. 1982. Signification du groupe leptynoamphibolique de l'Artense (massif Central Franqais). Thdse 3drne cycle. Clermont-Ferrand. 183 PP. Tr~OMPSON, A. B. 198I. The pressure-temperature (P, T) plane viewed by geophysicists and petrologists. Terra Cognita, 1, 11-20. THOMPSON, P. H. & BARD, J. P. 1982. Isograds and mineral assemblages in the eastern axial zone, Montagne Noire (France): implications for

61

temperature gradients and P - T history. Can. J. Earth Sci. 19, 129-43. THORAL, H, 1935. Contribution ~ l'6tude g6ologique des Monts de Lacaunes et des terrains cambriens et ordoviciens de la Montagne Noire. Bull. Serv. Carte gOol. Ft. 192-38, 319-637. VERNON, R. H. 1978. Prohyroblast matrix microstructural relationships in deformed metamorphic rocks. Geol. Rdsch. 67(1), 288-305. WINKLER, H. G. F. 1974. Petrogenesis of Metamorphic Rocks 3rd edn. Springer-Verlag, Berlin, 320 pp.

J. P. BURG & Ph. MATTE, Universit6 des Sciences et Techniques du Languedoc, Departement de G6ologie, Laboratoire associ6 au C.N.R.S. no. 266, Place Eug6ne Bataillon, 34060 Montpellier, France. A. F. LEYRELOUP, Universit6 des Sciences et Techniques du Languedoc, D6partment de G6ologie, Laboratoire de P6trologie, Place Eug6ne Bataillon, 34060 Montpellier, France. J. MARCHAND, Universit6 de nantes, Laboratoire de P6trologie, 2 Rue Houssini6re, 44072 Nantes C6dex, France.

Palaeozoic evolution of the Plateau d'Aigurande (NW Massif Central, France) Jean-Michel Quenardel & Patrick Rolin SUMMARY: The Plateau d'Aigurande represents the north-westernmost part of the French Massif Central. It is overlain by the sediments of the Paris Basin to the north and bounded by the La Marche shear zone to the south. Detailed mapping, mainly from the Creuse Valley region, indicates thrust and nappe tectonics. Beginning at the base of the sequence one can recognize: (1) the Foug6res unit (schists of low to medium grade) intruded by syntectonic leucogranites; (2) the Eguzon unit (medium- to high-grade schists which are partly diaphtoretic in the lower part of the unit); (3) the Gargilesse unit (high-grade schists); and (4) the Migmatitic unit. The metasediments of the Foug6res, Eguzon and Gargilesse units were originally coarse grained clastics. They are interbanded with magmatic units (orthogneisses and amphibolites). The migmatites are derived from greywacke-type rocks intruded by magmatic rocks (orthogneisses and amphibolites). Following a high-pressure (Silurian?) tectonometamorphic event, the piling up of tectonostratigraphic units occurred during two periods of deformation. The main one, probably of late Caledonian (Acadian?) age, was synchronous with or slightly after the climax of metamorphism. The second, of Westphalian age, was accompanied by the emplacement of leucogranitic magma and by retrograde metamorphism. The shear-sense appears to have been from SW to NE during the Acadian phase and from NW to SE during the Westphalian. Structural and lithological studies suggest that the three lowermost units may have been derived from the same palaeogeographic domain while the migmatites have a distinctly different history.

The Plateau d'Aigurande represents the north-westernmost part of the French Massif Central. It is limited to the west by the sediments of 'Le Seuil du Poitou' and to the east by a fault zone called 'Le Sillon Houiller'. Its northern margin is overlain by Liassic sediments of the Paris Basin while its southern border is defined by the well-marked morphological feature of the 'La Marche shear zone' (Fig. 1). The first (1:80,000) geological map of the region was published towards the end of the last century (de Launay 1893). It is still the only available detailed map of the Plateau d'Aigurande. The memoir by Yang Kieh (1932) contains some good descriptions of the crystalline rocks. More recently, Delorme & Emberger (1949), and Bouloton (1974) have interpreted the inverted metamorphic succession of the Plateau d'Aigurande as the preserved inverted limb of a large recumbent fold overturned to the south, with a d6collement between the migmatites and the schists: the conspicuous antiformal structure of the plateau resulting either from a 't6te plongeante' or from later folding. Courty (1952) has found evidence of another thrust within the sequence (Chambon thrust) in the area between Orsennes and the Creuse Valley.

Structural framework The general form of the structure is a large antiform with limbs gently dipping north in the northern part of the area and south to the south of the plateau. The hinge of this structure has been invaded by leucogranitic intrusions (Figs 1, 2 & 3). This antiform is associated with folds of hectometric to kilometric scale (see crosssection) which are best shown by the disposition of amphibolite bands. New geological investigations by the authors and students, carried out in the course of geological survey work, suggest that the inverted succession corresponds to a series of thrust nappes (Rolin & Quenardel 1980; Rolin 1981). We distinguish four superimposed tectonic units each associated with specific lithologies and each of which has had a distinctive tectonometamorphic history. The boundaries between these units are marked by thrust planes (Figs 1 & 2). The Chambon thrust separates the lowermost Foug6res unit from the Eguzon unit. The Gargilesse thrust separates the Eguzon unit from the overlying Gargilesse unit. The base of the uppermost thrust nappe, the Migmatitic unit, is underlain by the Migmatites thrust. These thrust planes are marked by tectonic discordances 63

64

J.-M. Quenardel & P. Rolin

FIG. 1. Map of the structural units of the Plateau d'Aigurande (NW Massif Central, France). which correspond to 'troncatures basales' at the base of the units and 'troncatures sommitales' in the upper part (Ellenberger 1967). The present position of the Gargilesse and the migmatitic units shows that the amplitude of nappe displacement, at least for these two units, reaches several tens of kilometres. A set of late E - W faults is considered to be

related to the La Marche shear zone, an eastward continuation of the southern branch of the South Armorican Shear Zone (Jegouzo 1980). The La Marche Shear Zone represents a major ductile shear zone situated along the limit between a northern continental block (the Plateau d'Aigurande) and a southern one (La Marche complex). This shear zone shows a sinistral

Evolution of the Plateau d'Aigurande

FIG. 2. Structural sketch map of the Plateau d'Aigurande (same legend as Fig. 1). (A), (B) Section lines (Fig. 3). Bedding and main foliation: (1) dip and strike; (2) vertical dip; (3) horizontal dip. Stretching lineation of probable: (4) Acadian age; (5) Westphalian age; (6) Upper Carboniferous age. Shear-sense indicators: (7) Acadian; (8) Westphalian; (9) Upper Carboniferous. Axial traces of late antiform (10); synform (11).

FIG. 3. Schematic structural cross-sections in the Plateau d'Aigurande: (A) along the Petite Creuse river; (B) along the Creuse river. (Legend and location on Fig. 1: black ornament in Eguzon unit = amphibolite; dashed ornament in Gargilesse unit = othogneisses.)

65

J.-M. Quenardel & P. Rolin

66

movement of unknown magnitude and a later dip-slip displacement of 2000-3000 m. The region as a whole still produces seismic activity.

Tectonostratigraphic succession The tectonostratigraphic succession of the Plateau d'Aigurande shows some variations between the eastern, western, northern and southern sides. We shall describe the succession from the Creuse Valley (Fig. 4) where good exposures exist.

The Foug~res unit (lowermost) The Foug~res unit is composed of leucogranites and supracrustal rocks.

The leucogranites The leucogranites which occupy the western central part of the plateau have a rather constant geochemical and mineralogical composition (quartz 35%, K-feldspar + albite ca. 50%, biotite 3-7%, muscovite 5-10%). Grain size is variable (1-8 mm) and porphyritic textures are sometimes developed (Figs 1 & 4). Contact metamorphism, characterized by andalusite and

tourmaline, appears for only a few metres along the northern border of the Crozant massif and to the east of the Crevant massif (Petitpierre 1981). The upper northern part of the Crozant granite contains a well-developed syn-magmatic foliation which bears a stretching lineation. This is correlated with a similar and parallel fabric in the enclosing schists. The schist fabric is related to early movements on the overlying Chambon thrust. This implies that granite emplacement and thrusting are synchronous events (Rolin & Quenardel 1982). Late post-metamorphic movements on this thrust have, however, cataclastically re-deformed the granite. Geochronological data (Rb/Sr whole rock isochron) show that the emplacement of the Crozant and Orsennes massifs occurred during the Westphalian (312 +- 20 Myr, 87Sr/86Sr = 0.7055---0.0060 (Rolin et al. 1982). The Crevant massif intruded the metamorphic rocks at about the same time ( 3 1 2 - 6 Myr, 87Sr/86Sr = 0.7082-+ 0.0015) (Petitpierre & Duthou 1980). The border of the Crevant massif is undeformed (Petitpierre 1981); thus the main movement of the Chambon thrust appears to be synchronous with the intrusion of the Crozant and Orsennes massifs, but slightly earlier than the intrusion of the Crevant massif.

FtG. 4. Tectonostratigraphical succession of the NW Plateau d'Aigurande complex.

Evolution o f the Plateau d'Aigurande

67

The supracrustal rocks

The Chambon orthogneiss

The Foug6res biotite garnet schists show a large variation in mineralogy (quartz 10-70%, albite, biotite, chlorite, muscovite, garnet, tourmaline, sph6ne, apatite 7. Their geochemistry indicates an original K-rich shale composition (Rolin 1981). The Messant acid meta-volcanites which overlie the Foug~res schists are of restricted extent. They contain quartz (60-80%), alkalifeldspar (10%), muscovite (5-20%), biotite and chlorite (5-10%) and sphene. The chemical composition suggests a felsic volcanoclastic origin for these rocks (Rolin, 1981; de La Roche 1965, 1968). Graphitic schists which form the upper part of the sequence are associated with quartzites. They immediately underlie the Chambon thrust.

This is a light pinkish to yellowish rock with a fine discontinuous banding which locally shows a well-marked NE-SW-trending linear (L-S to L) fabric. Its thickness is up to 400 m. It contains quartz (35-40%), K-feldspar (25-30%), oligoclase (25-30%7, muscovite (2-10%) and biotite (1-4%).

The Eguzon unit The estimated thickness of this unit is 2000 m in the northern part of the antiform although it could reach 3000 m in the south-eastern part. The lower part is composed of nodular gneisses. The upper part comprises a variable suite of rocks (metagreywackes, metapelites, orthogneisses and leptyno-amphibolites).

The nodular gneiss formation Their composition is variable. Usually, the matrix is quartz and K-feldspar rich (ca. 85%) with additional biotite and muscovite. The nodular texture is produced by matrixsupported centimetre-size porphyroclasts of feldspar and quartz. Their geochemistry shows a trend from granitic to arkosic composition (Bouloton 1974; Rolin 1981). The nodular gneisses are interbanded with a small number of quartzitic meta-psammite and metapelite layers.

Metagreywackes and metapelites The metapelites occur as a transition between the nodular gneisses and rocks higher in the sequence, they appear again above the leptyno-amphibolites. These dark-coloured rocks are mica-rich. Their texture is lepidoblastic. In addition to quartz (50-60%7 K feldspar (ca. 10%), biotite and chlorite (ca. 25%) and muscovite (10-20%), they contain garnet, sillimanite, kyanite and staurolite. The metagreywackes are fine-grained homogeneous rocks which split into slabs a few centimetres thick due to alternating lithologies. Their mineralogical composition is comparable with that of the metapelites with minor amounts of quartz (40-45%) and biotite (10-15%) but with increased feldspar (oligoclase, 25-45%).

The leptyno-amphibolitic complex This unit shows variable thickness (150-400 m) and composition ('Leptynite' is a light coloured gneiss mainly composed of quartz and alkali-feldspar of either sedimentary or magmatic origin). Petrography and geochemistry (Rolin 1981) allow us to distinguish rocks of sedimentary origin (quartzitic leptynite) or volcano-sedimentary origin (hornblende leptynite) from rocks of magmatic origin (amphibolites of probable alkaline basaltic composition 7. Eclogites (which are not in equilibrium with the country rocks) (Lasnier 1965) and serpentinites occur in small isolated pods within the amphibolites. The Gargilesse unit The most abundant rocks of the Gargilesse unit are coarse grained biotite-sillimaniteschists with a high quartz and feldspar content (ca. 40%). The phyllites are relatively rich in biotite (ca. 15%) and muscovite. Fibrous sillimanite, garnet, relics of kyanite and accessory minerals are also present. Some decimetric to metric-thick layers of amphibolites, quartzites and quartzitic schists are interbanded with the schists. In the lowermost part of the sequence quartzo-feldspathic mobilizates (i.e. newly formed Q - f mineral phases, after Mehnert 1968) appear, giving rise to migmatitic rocks (metatexites). Associated amphibolite and orthogneiss (in the Cerisier-La Mothe area) occur near the overlying migmatites where their thickness reaches 40 m. The Migmatitic unit (uppermost) The Migmatitic unit outcrops along the northern margin of the Plateau d'Aigurande and also in the klippe of Ch6niers to the south (Fig. 1). The migmatites vary from rocks where the palaeosome (parent rock) and the neosome (quartzo-feldspathic mobilizates) are recognizable (metatexites) to nebulitic gneisses (diatexites 7 and granitic anatexites. Agmatites occur in the northern Creuse Valley. The mineralogi-

J.-M. Quenardel & P. Rolin

68

cal composition is: quartz (25-35%, K-feldspar (15-25%, oligoclase (10-30%), biotite (15-25%), muscovite (0-5%), cordierite (0-12%) and also sillimanite and garnet. The rocks show discontinuous layering of gneissic and amphibolitic components. The chemical composition of the gneisses suggests a sedimentary origin (greywackes to pelites, Rolin 1981). The Ceaulmont orthogneiss forms a sheet about 100 m thick near the Creuse Valley, but it wedges out westward. The klippe of Cheniers is composed of anatexites and biotite granite.

Tectonometamorphic evolution The structural and metamorphic evolution of the Plateau d'Aigurande is complex and polyphase. We shall present some major features of this evolution in terms of each of the tcctonostratigraphic units.

The Foug~res unit Bouloton (1974) and Prost (pers. comm.) have drawn attention to two relics of kyanite and sillimanite in the schists of the Foug6res unit. These minerals could be related to earlier periods of high grade metamorphism comparable with the main recrystallization of the Eguzon rocks (see below). The more important metamorphism that we observe in the schists of the Foug6res unit is characteristically of low to medium grade. This is synchronous with the main foliation of the rocks which, as we have shown above, is of Westphalian age. The direction of the mineral lineation ( W N W - E S E ) , which we interpret as a stretching lineation, is parallel to a younger movement lineation. We assume that these two linear structures are related to a continuum of shearing deformation which accompanied the overthrusting of the Eguzon unit on to the Foug6res unit along the Chambon thrust. Small-scale folds and rotation of minerals in the schists immediately below the Chambon thrust indicate a shear-sense from NW to SE (Rolin 1981). However, preliminary studies of preferred lattice orientations of quartz seem to indicate an inverse ( S E - N W ) shear-sense (Cirodde 1981; Lerouge 1981). More detailed investigations are required to resolve this anomaly.

The Eguzon unit Several metamorphic mineral assemblages occur in this unit. The earliest, which is rarely

seen, corresponds to the eclogites. The most widespread assemblage, which occurs in the metapelites and metagreywackes, overprints this. This later metamorphism is high grade (kyanite, sillimanite etc.). The foliation that developed during this stage contains a NE-SW-trending mineral elongation (stretching) lineation. The rotation of minerals and the sense of vergence of small-scale folds (Rolin 1981) are consistent with a north-eastward shear-sense. These high-grade mineral assemblages have been partly retrogressed to chlorite-muscovite grade, particularly in the lower part of the Eguzon unit. The retrogression is associated with a deformation fabric which contains a W N W - E S E stretching lineation. This fabric is correlated with a similar cleavage in the biotite-garnet grade rocks of the underlying Foug6res unit where it is associated there with the syn-metamorphic movements on the Chambon thrust (see above). Thus it would appear that the magmatism and pro-grade metamorphism in the Foug6res unit (which accompanied the overthrusting of the Eguzon unit) retrogressed the earlier high-grade mineral assemblages of the Eguzon rocks. Radiometric data (K/Ar on amphiboles, Cantagrel 1973) indicate a Westphalian age for this late metamorphic event.

The Gargilesse unit Relics of kyanite represent the earliest observable stage of metamorphism in the Gargilesse unit. The Foliation developed during a later high-grade (sillimanite) metamorphism. The metatexites of the lowermost part of the unit show in places a L - S fabric with a weak NE-SW-trending mineral lineation. These structures could be related to the overthrusting of the Gargilesse unit on to the Eguzon unit. A preliminary study of preferred lattice orientation of quartz indicates a NE shear-sense (Schmitt 1982). The absence, in some areas, of the Gargilesse unit between the Eguzon unit and the Migmatitic unit probably indicates that the overthrusting of the Gargilesse unit on to the Eguzon unit took place prior to the thrusting of the Migmatitic unit on to the Eguzon and Gargilesse units.

The Migmatitic unit The earliest recrystallization in the Migmatitic unit (i.e. the metamorphic foliation within the enclaves in the agmatites) is overprinted by the high-grade metamorphism (anatexis). A

J . - M . Q u e n a r d e l & P. R o l i n

weak retrogression of this assemblage occurs in the lower part of the unit where it is associated with a deformation fabric in which the stretching lineation has a N E - S W trend with a S W - N E shear-sense (Lerouge 1981). This late tectonic and metamorphic overprint appears to be related to the overthrusting of the Migmatitic unit on to the lower grade Gargilesse and Eguzon units.

Conclusions The rocks of the Plateau d'Aigurande are the result of a complex geological history. The inverted metamorphic succession (higher grade in the upper part of the sequence) and the existence of 'troncature' suggest typical thrustnappe tectonics. Each structural unit had its own stratigraphic, magmatic and metamorphic evolution before the main overthrusting. The sediments of the Plateau d'Aigurande may be late Precambrian to Lower Cambrian in age since comparable rocks which occur in the BasLimousin area (south of the La Marche shear zone) are cut by Lower Palaeozoic granites (in the range of 4 7 2 - 4 3 2 Myr) (Bernard-Griffiths 1976; Bernard-Griffiths et al. 1977; Guillot 1980).

69

The timing of the first phases of metamorphism is difficult to assess. The eclogites of the Eguzon unit and the first metamorphism of the Migmatitic unit could be related to a Silurian event, by comparison with southern Brittany (Peucat & Cogn6 1977; Peucat et al. 1978). The climax of metamorphism in the Migmatitic unit, as well as in the Gargilesse and Eguzon units, took place just before the thrusting of the Gargilesse and the Migmatitic units (along the Gargilesse and Migmatites thrusts). By comparison with the Limousin (Guillot 1980) this main tectonometamorphic event would correspond to the Acadian phase (Bernard-Griffiths et al. 1977). The latest tectonometamorphic development ( C h a m b o n thrust) is synchronous with the emplacement of leucogranites during the Westphalian. The antiformal structure and faulting of the Plateau are related to a late tectonic overprint. ACKNOWLEDGMENTS:We are grateful to the 'France Profonde' team for helpful discussion and preparation of the manuscript. We thank Dr D. Hutton who kindly reviewed and annotated the English text. Contribution franqaise no. 29 au P.I.C.G. no. 27 'Orog6ne cal6donien'. French contribution no. 29 to I.G.C.P. project no. 27 'Caledonide orogen'.

References BERNARD-GRIFFITHS,J. 1976. Essai sur les figes au DELORME,J. & EMBERGER,A. 1949. La s6rie cristalStrontium dans une s6rie m6tamorphique: ie Bas lophyllienne renvers6e du Plateau d'Aigurande. Revue Sci. nat. Auvergne, 15, 45-82. Limousin. Ann. Sci. Univ. Clermont, 55, 27, 243 pp. ELLENBERGER,F. 1967. Les interf6rences de l'6rosion et de la tectonique tangentielle tertiaire dans le , CANTAGREL,J. M. & DUTHOU, J. L. 1977. Bas-Languedoc (Principalement dans l'Arc de Radiometric evidence for an Acadian tectonometamorphic Event in western Massif CenSaint-Chinian); Notes sur les charriages cisailtral Fran~ais. Contr. Miner. Petrol. 61, 199-212. lants. Revue G~ogr. phys. Gdol. dyn. (2) 9, 2, BOULOTON,J. 1974. Etude g6ologique de la r6gion 87-142. d'Aigurande (NW du Massif Central franqais). GUILLOT, P. L. 1980. La s6rie m6tamorphique du Bas-Limousin: de la vall6e de l'Isle ~ la vall6e de Lithostratigraphie, structure et p6trographie de la s6rie m6tamorphique. Th~se de 3dine cycle. la Corr6ze, le socle en bordure du Bassin AquiClermont-Ferrand. 166 pp. tain. Th~se Doctorat Os Sciences. Universit~ CANTAGREL,J. M. 1973. Signification des figes ?a Orl6ans. 391 pp. l'Argon d6termin6s sur amphiboles dans les JEGOUZO, P. 1980. The South Armorican Shear socles m6tamorphiques anciens. Application au Zone. J. struct. Geol. 2, 39-47. LA ROCHE, H. de 1965. Sur l'existence de plusieurs Massif Central franqais e t ~ l'Aleksod, Sahara facies g6ochimiques dans les schistes pal6ozoialg6rien. Th~se de Doctorat ~s Sciences. ques des Pyr6n6es luchonnaires. Geol. Rdsch. 55, Clermont-Ferrand. 282 pp. CIRODDE, J. L. 1981. Etude g6ologique et structurale 274-301. 1968. Comportement g6ochimique diff6rentiel de la r6gion de Cluis et argument pour un sens de mise en place des nappes ~ l'aide de la fabrique de Na, K, et AI dans les formations volcaniques du quartz, Plateau d'Aigurande (NW du Massif et s6dimentaires. Un guide pour l'6tude des formations m6tamorphiques et plutoniques. Central fran~ais). Unpublished Dipldme C.r. hebd. S~anc. Acad. Sci., Paris, 267D, d'~tudes approfondies. Universit6 Paris-Sud, 39-42. Orsay. 47 pp. COURTu G. 1952. Observations tectoniques sur la LASNIER,B. 1965. Etude p6trographique de la r6gion partie Nord du Plateau d'Aigurande. C.r. Somm. d'Eguzon (Indre). Coupe du versant Nord de l'anticlinal du plateau d'Aigurande ~ zon6ogSdanc. Soc. g~ol. Fr. 312. -

-

70

J.-M. Quenardel & P. Rolin

raphie invers6e. Dipl6me Etudes SupOrieures. Universit6 Poitiers. 105 pp. LAUNAY, L. de, 1893. Carte gdologique 'Aigurande' au 1/80 000. LEROUGE, G. 1981. Etude g6ologique et structurale de la r6gion de Saint Benoit-du-Sault et argument pour un sens de mise en place des nappes l'aide de la fabrique du quartz, Plateau d'Aigurande (NW du Massif Central fran~ais). Unpublished Dipldme d'dtudes approfondies. Universit6 Paris-Sud, Orsay. 50 pp. MEHNERT, K. R. 1968. Migrnatites and the Origin o f Granitic Rocks. Elsevier, Amsterdam. 393 pp. PETITPIERRE, E. 1981. P6trographie, g6ochimie, m6tallog6nie du granite de Crevant et de son contexte mdtamorphique et structural (Plateau d'Aigurande, Massif Central Fran~ais). ThOse 30rne cycle. Clermont-Ferrand. 227 pp. -& DUTHOU,J. L. 1980. Age westphalien par la m6thode Rb/Sr du leucogranite de Crevant, Plateau d'Aigurande (Massif Central fran~ais). C.r. hebd. Sdanc. Acad. Sci., Paris, 291D, 163-6. PEUCAT, J. J. & COGNI~, J. 1977. Geochronology of some blueschists from Ile de Groix (France). Nature, 268, 131-2. , LE MI~TOUR,J. & AUDREN,C. 1978. Arguments g4ochronologiques en faveur de l'existence d'une double ceinture m6tamorphique silurod6vonienne en Bretagne m6ridionale. Bull. Soc. g~ol. Fr. (7), 20, 2, 163-7.

ROLIN, P. 1981. Geologie et structure du Plateau d'Aigurande dans la r6gion d'Eguzon (NW du Massif Central fran~ais). ThOse 30me cycle. Orsay. 229 pp. --, DUTHOU, J. L. & QUENARDEL, J. M. 1982. Datation (Rb/Sr) des leucogranites de Crozant et d'Orsennes. Cons6quences sur l'gtge de la dernibre phase de tectonique tangentielle sur le Plateau d'Aigurande (NW du Massif Central fran~ais). C.r. hebd. SOanc. Acad. Sci., Paris, 294, II, 799-802. --& QUENARDEL,J. M. 1980. Nouvelle interpr6tation du renversement de la s6rie cristallophyllienne du Plateau d'Aigurande (NordOuest du Massif Central, France). C.r. hebd. Sdanc. Acad. Sci., Paris, 290D, 17-20. & ~ 1982. Mod61e de mise en place syntectonique d'un massif de leucogranite hercynien (Crozant-NW du Massif Central fran~ais). C.r. hebd. Sdanc. Acad. Sci., Paris, 294, II, 463-6. SCHMIrr, P. 1982. Etude g6ologique de la coupe de la grande Creuse (Plateau d'Aigurande). Unpublished DiplOme d'dtudes approfondies. Universit4 Paris-Sud, Orsay. 51 pp. YANG K1EH, 1932. Contribution ~ l'6tude g6ologique de la chaine de la Marche et du Plateau d'Aigurande. Mere. Soc. gOol. Fr. 19, 122 pp.

-

-

JEAN-MICHEL QUENARDEL & PATRICK ROLIN, Laboratoire de G6ologie Structurale, Universit6 Paris-Sud, brit. 504, F-91405 Orsay C6dex, France.

Northern margins of the Variscides in the North Atlantic region: comments on the tectonic context of the proMem S. C. Matthews SUMMARY: Some of the misconceptions which have figured in discussions of Variscan geology are analysed. Subduction (but not in the obvious sense), strike-slip movement and sole thrusts may all have a part to play, but none of these alone explains the Variscides. Kossmat's scheme of tectonic zones still deserves attention. So, too, do questions concerning the pre-Variscan geology, especially the prevalence of rift-structures, in central and western Europe. The problem of integrating the pattern of Palaeozoic tectonism in Europe with that in eastern North America is briefly explored. It is suggested that Caledonian and Variscan be treated as parts of one chapter of tectonism and that much of the early tectonic history of central and western Europe be regarded as Pan-African. It is argued that western Europe is unlikely to produce evidence of structural transpositions on the scale recently revealed by seismic reflection profiling in the southern Appalachians. It is also suggested that the idea of a Variscan front is largely illusory.

The Variscan puzzle Solutions of the problem presented by the Variscides have been many, and consistently unsuccessful. Attempts to understand the nature of the problem have appeared less frequently. One practical aspect of the problem concerns the difficulty of handling information emerging from several parts of Europe and presented in several different languages. For this, and for other reasons, it is difficult to recognize PanEuropean traits.of tectonic character in the Variscides. It is not immediately obvious why geologists choose to refer to a Variscan 'fold belt'. One proposal has been adopted widely in Europe (and lately in Great Britain). Kossmat's (1927) scheme of tectonic zones for the Variscan geology of central Europe has survived the tests set during more than 50 years of accumulation of new information and insertion of new ideas, and has consistently been regarded as a meaningful analysis of Variscan tectonism in central Europe. Yet there has never been a clear sight of the course taken by these zones farther west in Europe. Consider the case of south Cornwall, an area which includes the variety of rocks, Precambrian to Upper Devonian, exposed in the Lizard peninsula and adjacent ground. Is south Cornwall to be regarded as a continuation of the Saxothuringian zone, as Stille (1951) proposed? Or, given the presence on the Lizard peninsula of ultrabasic and other intrusives, plus demonstrably older high grade metamorphic rocks, is south Cornwall better regarded as Moldanubian? Do we capture the significance of the Lizard association of rocks by applying any one such label, or by citing any one date? Consider also the case of the old,

pre-Brioverian rocks in north-western France and the Channel Islands. They are much older than anything that is widely detectable in central Europe (Bernardova & Chab 1968; Dornsiepen 1979; van Breemen etal. 1982). How should we hope to account for them in any westward extrapolation from Kossmat's ground? Should we in this case abandon hope of recognizing continuations of the elements of Kossmat's scheme, and rely instead on some less easily testable idea such as deciding that N W France contains an antique microplate or is a microcontinent or a microcraton? Models of the latter type represent an abuse of plate tectonics, quick answers which are of no real service to us in any attempt to understand what Cogn6 (1976), with entirely appropriate use of words, called 'le puzzle Varisque'. Variscan Europe is indeed a puzzle, a number of samples of the European crust each of which has been studied according to the geological traditions of the region in which it occurs. It has never been easy to understand what the evidence from all of these regions, taken together, might mean. The advent of plate tectonics brought readier comment on the overall meaning of Variscan geology than had ever been possible before. No experience of the several parts of the puzzle was now needed, it seemed. The language barriers which had earlier made it difficult to establish a grasp of all of the separate threads in discussions of the geology of Variscan Europe had now, apparently, lost their significance. It was as if plate tectonics had come as a form of Geo-Esperanto, to open up a freer form of communication and raise some promise of a new understanding. A first caveat may have been present in the minds of some observers: Esperantists do not enjoy access to everything 71

72

S. C. Matthews

that is best in the established literatures of Europe. It was, in any case, wise to resist the blandishments of those who emerged in the early 1970s with proposals on consuming plate boundaries. One dares to ask now: did these plate tectonical propositions make anything more clear than it had been before? Dewey & Burke's (1973) remarks on basement reactivation had more merit than anything else in the expressed opinions of the time. The legacy of the 1970s is a number of unsolved (and in some cases, possibly misconceived) problems: is there evidence of a consuming plate margin still to be sought in the Variscides, or evidence of a Rheic or of a Theic ocean? Should one hope to produce a case for recognizing a complete Wilson cycle in the Variscides? Is there reason to regard the Variscides as a peer fold belt of the Caledonides and the Alpides? A useful test of our understanding of Variscan geology comes with this question: how would the geology of central and western Europe appear if all of the post-Variscan cover were stripped off? We might expect to see more granitic rocks (Dornsiepen, 1978, has estimated that granitoids occupy approximately one-quarter of the total area of Variscan outcrop available at present); but would the whole assembly appear more belt-like than it does now? The layout of Variscan geology in Europe, as at present seen, is not conspicuously belt-like. The evidence of Variscan tectonism is exposed in a scatter of regions from southern Ireland, Brittany, Galicia and Portugal in the west to Czechoslovakia, the Eulengebirge and the Holy Cross Mountains in the east. It has been said that the Variscan fold belt extends for some 2000 km across the general strike, from the Variscan front in the north to the South Atlas fault in the south. But since Variscan geology occurs southward of the South Atlas fault, in the Mauretanides, and since it is difficult to attach any clear meaning to the idea of a Variscan front, the figure of 2000 km may not be particularly instructive. In the east, Variscan geology appears to end in the neighbourhood of Tornquist's Line, i.e. approximately at the boundary between the Russian Platform on the east and the more elaborate pattern of European geology on the west. The Caledonian fold belt, it should be noted, takes less account of Tornquist's Line (although there are a number of differences between the British and the Scandinavian Caledonides which are not yet understood: Nicholson 1970) and continues through Scandinavia to enc~ in a large arcuate structure in the Timan-Pechora. A number of features distinguishing Variscan ('Her-

cynotype') geology from geology as seen in the Caledonides and the Alpides ('Alpinotype') were clearly stated by Zwart (1967) who emphasized the abundance of Variscan granites and particularly the fact that metamorphism is dominantly of low pressure-high temperature character, which implies high heat flow through the Variscan crust. We should have in mind here also the fact that, in Europe, Caledonian and Variscan geology are separate in space. This contrasts with eastern North America, where the two coincide, and where authorities such as King (1969) and Rodgers (1970, 1971: see especially his footnote on p. 1172) have not been satisfied that there is a case for treating Caledonian and Variscan as separate tectonic episodes. There have been numerous conferences on Variscan geology during the last 10 years. The theme chosen for the Dublin meeting in September 1982 was 'the northern margins of the Variscides in the North Atlantic region', and was designed to examine the northern marginal zone of the Variscides and in particular the nature of the so-called Variscan front. This revives the long-standing question of the possible identity or identities of such a tectonic feature (Matthews 1974). What should we imagine the 'Variscan front' to be? A southern limit of economic possibility in the minds of oil company geologists as they review Palaeozoic prospects in Europe? A Variscan analogue of the Moine Thrust? Despite the fact that the Variscan front has meant different things to different people at different times the idea of a Variscan front is with us. Innumerable published maps indicate its course across Europe. Would we feel able to attempt to project the boundaries of the central European zones westward into the SW British Isles, the English Channel and NW France if we did not have the example of the Variscan front to guide us? What do we gain by such guidance in hoping to understand, for example, the pre-Mesozoic geology of the Celtic Sea region between Cornwall and Ireland (Gardiner & Sheridan 1981; but compare Matthews 1984a)? How well are we served by such guidance at the western end of Europe, where some propose that the Variscan front carries on towards America--e.g. Rast & Grant (1973, 1 9 7 7 ) - - b u t numerous others identify a curvature in major structure (first proposed by Suess in 1887: who thus gave notice of a problem which no-one has yet satisfactorily solved) which would bring the Variscan tectonic zones in an arc to re-enter Europe in the Iberian peninsula (Kossmat 1921: note an objection to

Northern margins o f the Variscides

73

(2) If one proposes, as has been done in the past (Stille 1951) that the southern part of Cornwall is a SW England representation of a westward continuation of the Saxothuringian zone, the question then arises: how much of the Saxothuringian zone does Cornwall represent? Should one choose to follow Shackleton, Ries & Coward (1982: who f o l l o w - - o n e pres u m e s - - A u t r a n et al. 1980, who do not explain their reasons for representing the Saxothuringian zone in the manner shown on the Tectonic Map of France) in suggesting that the Saxothuringian zone occupies much of the width of the present English Channel, southward to some boundary with the Moldanubian zone (although the fitness of any such boundary for any such role is not made clear by Shackleton et al.)? Or should one follow Read & Watson (1975, fig. 3.2) who give the Rhenohercynian zone as including much of Normandy and show the Saxothuringian zone as occupying all of Brittany? Or should one be guided by van Breeman et al. (1982), whose impression of the layout of Variscan geology in Europe is that of Ellenberger & Tamain (1980) but who insert the idea of a subduction zone running from the NE part of Bohemian Massif, through the Harz Mountains, the Hunsriick and eventually north-westwards through northern France to the southern part of Cornwall, where their suggestion of a structure appears to depend on geochemical information. If one follows Shackleton etal. (1982) the implication is that much of the geology of Kossmat's zones in western Europe? north-western France represents a continuation of the Moldanubian zone. Such a proposition The discussion of Variscan geology in Europe which follows concerns itself first with the ques- would stumble on the fact, already mentioned tion of the applicability of Kossmat's zones in above, that NW France and the Channel Islands western Europe. Certain problems concerning show clear evidence of the survival of rocks distinctly older than, say, 800 Ma and are thereSW England have already been touched upon. Closer examination of the possible affiliations of fore significantly different from anything that is present in the Moldanubian zone in central that region reveals the following difficulties: (1) If, as is generally agreed, the major part of Europe. SW England represents the Rhenohercynian (3) If one prefers to identify south Cornwall, zone, it is helpful to point out that this resemb- with its serpentinite and the pre-serpentinite, Precambrian, high grade lance refers to the stratigraphic evidence alone, presumed and that in terms of structure, and taking into metamorphic rocks, as a western representative account also the presence of a large granite of the Moldanubian zone, the result would be batholith within SW England, there is much attractive to some in that the Saxothuringian more of a resemblance to the style of the Saxo- zone would appear to be absent (note, in Franke (this volume) a map in which the Saxothuringian thuringian zone. And if one remarks, too, that the SW England granite batholith is tin-bearing, zone pinches out, for whatever reason, as it runs does that do further damage to the conven- south-westward from Germany into France). tional view that SW England represents a west- This interpretation might be expected to claim ward continuation of the Rhenohercynian the interest of proponents of large-scale--south zone? Is it imaginable that the area can be Appalachian scale--thrusting in the southwestern region of the British Isles (Shackleton el Rhenohercynian in one sense and Saxothuringal. 1982; Cooper et al., this volume; but see also ian in another?

his proposals in Wegener 1924; Lotze 1945: note kindred proposals by numerous later commentators down to Matte et al., this volume; see Matthews 1984b). Is the observed arc correctly identified as a tightly curved expression of a continuation of Kossmat's (1927) apparently simple set of central European tectonic zones? What is the nature of the I b e r o - A r m o r i c a n arc? What is the age of its first establishment? Why is it taken to lie well to the south of Suess's (1887, 1888) Armorican arc? Is the quasi-concentric Asturian arc which contains Carboniferous rocks in northern Spain (Ribeiro 1974; Ries & Shackleton 1976) an integral part of the same scheme? Why does Lefort (1979) find that in Brittany the arcuate structures are transected by features due to subduction in Devonian time? Do we regard the IberoArmorican arc as being arcuate because of gross Variscan deformation, or as having been in some original, earlier state arcuate? Before proceeding to make proposals on links with the Appalachians (an especially important matter, given recent observations on deep structure there, and given also that the Dublin meeting was the first conference on Variscan geology to have been held in general awareness of the C O C O R P findings) we should first make some pertinent points concerning Variscan geology in Europe and should ask what might be the worth of any Pan-European application of a scheme such as Kossmat's.

74

S. C. Matthews

Matthews 1984a and further argument against such proposals below). ff south Cornwall is taken to be Moldanubian and one thinks fit to assume that any westward continuation of the Saxothuringian zone has been excised, the problem which then emerges is that one has already exhausted one's quota of Variscan tectonic zones within SW England before making any start on the question of the zonal affiliation of Variscan geology in France. (4) Floyd (1982) employs his own impression of the layout of Variscan tectonic zones as a basis for discussion of variations in the geochemical characteristics of Devonian and Carboniferous basic igneous rocks. His opinion that the Saxothuringian zone can be identified in the Chateaulin and Laval synclines in Brittany has little to recommend it. (5) Oliveira, Horn & Paproth (1979) offer a different solution to the problem of tracing the continuation of the Saxothuringian zone in western Europe. They recommend a rotation of the Iberian peninsula through 130 ~, which would put the Beja massif into alignment with the mid-German crystalline rise and would thus produce an arrangement in which evident and long-recognized similarities shown by the Rhenohercynian zone in Germany and the south Portuguese zone in the (now) SW part of the Iberian peninsula would be more easily understood. Oliveira etal.'s (1979) proposal does not, however, help to ease the difficulty of recognizing any continuation of the Saxothuringian zone in NW France. It is not clear how such a proposal would fit with the evidence of Variscan tectonism in North America, nor, within a smaller compass, is it clear how a rotated Iberia would relate to the late Palaeozoic geology of southern Ireland (Matthews, Naylor & Sevastopulo 1983). If anyone were to remark that the suggested 130 ~ rotation produces an interesting pattern, a possible rejoinder would be that rotation through 360 ~ also produces an interesting pattern. A more helpful response might come in the form of this question: why should anyone insist that more or less straight lines, or anything with a belt-like layout, represent the answer to the problem of tracing the interrelationships of the major tectonic features of Variscan Europe? The Saxothuringian zone is especially awkward in all of these attempts to contrive a west European accommodation of Kossmat's zones. It has been something of a Gothic mystery in several recent plate tectonic models of Variscan geology, so much so, indeed, that some have felt it to be the site where their elusive Variscan suture may lurk and where evidence of disap-

pearance of some former part of the European lithosphere may lie. Bromley (1975), for example, employed such a proposal in attempting to develop an explanation of the tin-bearing Variscan granites in Europe. His hypothesis takes no account of the significance of the finely divided tin minerals in the late Precambrian there (see more recent discussions of that evidence in Weinhold 1977 and Wienholz, Baumann & Hofmann 1982). Or again, B a r d et al. (1980) have proposed to link south Cornwall with the Saxothuringian zone in their suggestion of a Lizard-Mtinchberg nappe. In so doing, they have taken little account of anything that is known of Variscan geology in the intervening regions. For example, the trace of their suture lies close to the site of the Saar 1 borehole, where non-metamorphic Middle Devonian has been found resting on a Devonian (380 Ma) granite (Zimmerle 1976; M/iUer 1978). Behr, Engel & Franke (1982, see also Franke, this volume) have, in this light, done European geology some service by providing an up-to-date account (in English) of the geology of the Saxothuringian zone. Its chief characteristics, and those of what can be regarded as kindred tectonic entities in Europe, should be briefly restated here.

The Saxothuringian zone In the Saxothuringian zone (Behr, 1961, 1978; Behr, Walliser & Weber 1980; Behr etal. 1982; Franke, this volume) a special set of Carboniferous rocks called the Bavarian facies, has been thrust over the more normal (Thuringian facies) succession of greywackes and shales. The Bavarian facies contains a variety of clasts: Carboniferous limestone (first deposited in a more shallow environment elsewhere) earlier Palaeozoic rocks, metamorphites. It is itself structurally overlain by non-metamorphic early Palaeozoic rocks, and these have been overriden by early Palaeozoic rocks in amphibolite facies. The highest unit in the structural succession has high grade metamorphic rocks, including granulites and eclogites, some of whose educts can be identified on radiometric evidence as having been, again, of early Palaeozoic age. Peaks of metamorphism (see Behr et al. 1980, fig. 3) are of Ordovician and Devonian age. The former might best be termed the 'Ordovician thermal event' (Zwart 1976; Zwart & Dornsiepen 1978) rather than 'Caledonian'. Nor is it clear that any worthwhile purpose is served by calling the Devonian event 'Acadian'. That term may still be useful in reference to events of early

Northern margins o f the Variscides

75

Devonian age in north-eastern America; but it incorporated in the late range of the Devonian succession in the southern part of what is now does not nowadays conveniently apply to events the Rhenohercynian zone. in central Europe. One notes, and approves of, These major elements in the Palaeozoic Autran & Cognr's (1980) insistence on using structure of central Europe, Moldanubian zone 'Ligrrien' (Pruvost 1949) rather than Acadian and Saxothuringian zone, are clearly identifiin discussing the geology of south Brittany. able in a region which extends SW as far as the A more substantial problem in the Saxoline of the Bayerischer Pfahl, a prominent fracthuringian zone is that of explaining how ture zone which strikes N W - S E and whose conOrdovician rocks were brought to high tinuation to the NW can be seen in the Altenmetamorphic grade during Devonian time. bfiren fault, a structure which was active during There is no basis for any suggestion that Devonian time as is shown in the fact that it Silurian-Devonian cover of the thickness (possoffsets features on the shelf margin at the (then) ibly as much as 20 kin) implied by the northern limit of the Rhenohercynian basiri metamorphic grade might at one time have existed in the region and might then have been (Meischner 1964). In discussing the Saxothuringian zone we are removed to expose the metamorphic rocks to referring to an important feature of the crust in erosion (they are represented among the clasts central Europe which, according to Behr in the Bavarian facies) during Carboniferous (1978), first took on a special character in an time. The high grade metamorphic rocks, episode of rifting, vulcanism and insertion of according to Behr et al. (1982) were brought to intrusions in late Precambrian to early Palaeoshallow depth and emplaced as nappes, now zoic time. Some of the special qualities of that resting on non-metamorphic Palaeozoic original association of material are detectable in (including Carboniferous), by the operation of the kornerupine rocks at Wildenfels in Saxony what Behr (1978) and Weber (1978) have calwhich Schreyer, Abraham & Behr (1975) led a 'subfluence' zone, a delamination (in interpret as metamorphosed evaporite material. Bird's 1978, 1979 sense) which involved subWhat we regard now as the Saxothuringian duction of lithospheric mantle (but not crustal zone is the rift assembly, plus the additions of rocks, and certainly not oceanic crust) and subsequent Palaeozoic time, deformed and returned high grade metamorphics to shallow metamorphosed during what we call the level (see a more explicit reference to the proVariscan orogeny: a synergy of late Precamcess, in English, in Weber 1981). The root zone brian rifting and mid-late Palaeozoic heating. of the nappes, and therefore the trace of the The Saxothuringian zone, with a deepsubfluence zone, is identified at the Erbendoff reaching structure at each boundary, is the best line, long regarded as the boundary between determined entity in the scheme of Variscan the Saxothuringian and Moldanubian zones, tectonic zones for central Europe. The Rhenowhere ultrabasic rocks come to outcrop. Late hercynian zone is therefore firmly delimited at (Carboniferous-Permian) events within the its SE boundary, where it marches with the Saxothuringian zone included emplacement of Saxothuringian zone; but it is not clear that its the Fichtelgebirge granite pluton and the estabNW boundary represents a structure of comlishment, on its northern side, of the shallow parable magnitude. This boundary, taken to synform which now contains the detached masseparate the Rhenohercynian zone from the ses, the Mfinchberg Massif and others, in NE sub-Variscan foredeep, is customarily set where Bavaria and Saxony which are called collecthe Carboniferous succession shows a change tively the 'Zwischengebirge'. from dominantly greywacke character to clasThe northern limit of the Saxothuringian tics of a different association ('F16zleeres' and zone, i.e. its boundary with the Rhenohercynian finally 'paralic') on the NW. According to zone, is regarded as a further subfluence zone Franke et al. (1978) this involves only a quan(Weber 1978). In this case, the suggestion titative change, with paralic successions receives support from geophysical work done developed when accumulation tended to outdo by Giese (1978), who identifies the Rhenosubsidence. There is no evidence to suggest that hercynian-Saxothuringian boundary as proa deep reaching boundary comparable with the ceeding to a depth at which it offsets the Moho. one at the SE border of the Rhenohercynian In the northern part of the Saxothuringian zone exists here. Further, the NW boundary of zone, close to the trace of the SE-dipping subthe Rhenohercynian zone is not entirely effecfluence zone at its northern limit, is the midtive as a limit of cleavage. The evidence of the German Crystalline Rise first identified by M/insterland I borehole (Fuchtbauer 1963; see Kossmat (1927) and long regarded (Brinkmann 1948) as a source of the greywacke material also Teichmiiller, Teichm/iller & Weber 1979)

S. C. Matthews

76

shows that slaty cleavage exists in Upper Palaeozoic rocks at a depth of 4000 m. This occurrence of slaty cleavage at depth in the sub-Variscan foredeep may, like the occurrence of a tin-bearing granite in the 'Rhenohercynian zone' in SW England, serve as a hint that although Kossmat's (1927) zonal analysis still deserves our entire respect, we should nowadays perhaps hope to understand more fully the continuation in depth of the zones defined at outcrop. Given that the Saxothuringian zone is a dominant feature of crustal structure in central Europe, it is difficult to understand all that is implied by the problem of finding a south-westward continuation of its course through the SW part of Germany. One may freely ask whether it exists there, or whether it is broader there, perhaps less foreshortened by late Palaeozoic deformation. In any attempt to identify its course in France, a reasonable first presumption is that it might run toward the Massif Central; but there Burg & Matte (1978) suggest that Variscan tectonism is directed southwards, i.e. in a sense opposite to that seen in central Europe, where Franke (this volume) is prepared to envisage transpositions totalling 200 km of movement land directed northwestwards. The problem of reconciling central European proposals on Variscan tectonism with those raised in France is one of the most difficult tasks in European geology. It is unlikely that any worthwhile solution will come from those who readily suggest Variscan subduction zones (see comment in Matthews 1 9 8 4 a - - a n d in van Breemen et al. 1982). Nor is strike-slip faulting likely to produce a complete answer.

of the rifting. These are not isolated cases. Others, e.g. the Malvern structure in West England, can be added. Matthews (1984b) has argued that the Galician rift structure continues on the north side of the Bay of Biscay as 'l'accident de la Petite Sole' (Guennoc 1978; Lefort & Ribeiro 1980) and runs from there as a graben development, with a Mesozoic history totally different from that experienced by the Galician segment of the original structure, to the Bristol Channel, eventually to link into a triple junction whose northward directed arm is the Malvern structure (in Mesozoic time the site of graben development) and whose southeastward directed arm underlies the Wessex basin in England and the Pays de Bray structure in France as it runs towards Paris. The arcuate pattern of structure is shown in Fig. 1. It should be understood that Fig. 1 attempts to represent only part of the effects of a widespread episode of late Precambrian rifting. The Tayvallich rifting event (Graham 1976), a kindred development, took place near a site where rifting became spreading and the Iapetus Ocean opened: central and western Europe show structures which match in age the beginnings of Caledonian tectonism. Zwart & Dornsiepen (1978: they recommend to students of Variscan geology the study also of pre-Variscan geology in Europe) have noted this broad time-relationship and have suggested that the opening of the Iapetus Ocean may have produced compression in what we regard as the Cadomian tectogene: but whether such an effect explains everything that has been called Cadomian (or Assyntian, or Baikalian, cf. Wood 1974) remains to be seen.

Right-lateral shear

Palaeozoic geology of central and western Europe These problems may be brought clearly into focus by identifying three principal effects in the Palaeozoic geology of central and western Europe:

Late Precambrian rifting The original site of the Saxothuringian zone may be regarded as one case of this kind. The Hesperian aulacogene in NW Iberia (Den Tex 1981; van der Meer Mohr etal. 1981) is another. In both cases it is clear that tectonic activity continued for some time after inception

Arthaud & Matte's (1977) stimulating set of suggestions is inexact. It is, for example, necessary to take account of proposals made by Str6mberg (1976) in the region east and west of the Gulf of Bothnia, i.e. in an area which Arthaud & Matte's (1977) maps leave blank. Secondly, Arthaud & Matte failed to see that their N W - S E fractures are a dominant feature of the SE plate in the Iapetus system. No such trend of fractures is found in the NW plate. Thirdly, Arthaud & Matte (1977) were plainly in error in referring to their fracture system as 'late Hercynian'. A number of the fractures involved were of much earlier establishment. Some (e.g. the Altenbfiren fault mentioned above) were clearly already active during the Devonian. The fracture pattern identified by Arthaud & Matte (1977) is a fracture pattern

Northern margins of the Variscides

x;,

77

\

FIG. 1. Map of the North Atlantic region (le Pichon restoration) to suggest sites of some major fractures (largely after Arthaud & Matte) and, with dotted ornament, possible relationships of rift-like structures established in late Precambrian-early Palaeozoic time. The pattern shown in Europe and NE North America is, very crudely, a 'mid-Palaeozoic' state of development, before Variscan heating. The southern Appalachians, shown as a belt of thrusts, came to this state of development much later in Palaeozoic time.

activated at a time when the Iapetus Ocean was closing. The fractures sketched in Fig. 1 are consistent with the pattern shown in Dewey (1982, Fig. 37). Dewey (1982) regards these as late Devonian to early Permian transform lineaments. Since the system relates to the closure of the Iapetus Ocean, and since it is represented in Scandinavia (Str6mberg 1976), it is likely that Silurian (Bassett, Cherns & Karis 1982) to early Permian is a better statement of the effective range in time. These strike-slip faults do not explain the Variscides. The distribution of Variscan tectonism, as indicated by Devonian and early Carboniferous volcanism, by evolution of granite plutons and by early development of low grade metamorphism, does not match the distribution of the end-Iapetus shear effect. The problem of Variscan heating remains.

the heat source, so that in Devonian time heat was being lost into the relatively young crust of what had formerly been the SE plate, with consequences in terms of metamorphism of Devonian age such as are seen in the Saxothuringian zone. If central and western Europe were at this time overlying a locus of major heat loss, they were exposed to the possibility of evolving further according to constructive plate margin models. We may take it that such a development did proceed, but briefly, and is documented in the evidence now available in south Brittany. Farther east, in central E u r o p e , the evidence is better interpreted in terms of a ~ constructive plate margin manqu6. Tendencies in the chemical characteristics of basic igneous rocks, which some assume to be indications of oceanic character, deserve a more direct interpretation as evidence of mantle affiliation.

Variscan heating Zwart (1967, 1976; see also Zwart & Dornsiepen 1978) has consistently maintained that the principal problem requiring to be explained in Variscan geology is the high heat flow, as indicated by the abundance of granites and the high t e m p e r a t u r e - l o w pressure character of metamorphism. A thin crust is implied. Matthews (1978) suggested that access to the major heat source may be explained as a consequence of the closure of the Iapetus Ocean. One understands that when the Iapetus Ocean was open, in early Palaeozoic time, it had a mid-oceanic ridge, a site at which heat generated deep in the earth's interior could be lost. One presumes that when Iapetus closed the SE plate overode

Comments on the Palaeozoic geology in eastern North America In all three of the arguments pursued above, the suggestion emerges that what we call Variscan geology should be understood to be intimately related to Caledonian geology, from inception of the pattern of arcuate structure in late Precambrian time to acquisition of a major 'Caledonian' source of heating in midPalaeozoic time. We proceed to consider briefly how such arguments bear on the possibility of proposing E u r o p e a n - A m e r i c a n relationships in Palaeozoic geology. Matching western Europe and eastern North American tectonic patterns is as difficult as

78

S. C. Matthews

matching central with west European evidence. However, the problem of choice mentioned at an earlier stage of the discussion--belt-like continuation of Variscan structure westward toward America or a European-American relationship involving arcuate structures--seems to have been settled in favour of the second alternative by Lefort & Haworth's (1979) discovery of large curvatures in the deep structure of both offshore western Europe and offshore eastern Canada. Fig. 1 includes a representation of their findings, and suggests a linkage into North America by adopting Rast, Rast & Kennedy's (1975) observation that belts of structure are set at right angles to one another in the Newfoundland-Labrador region. Given that the writer has no direct experience of the geology of eastern North America it seems prudent to proceed by asking some of the questions which reasonably follow from the interpretation of European evidence advanced here. And given restrictions on space, it seems best to avoid the more banal questions (such as when did the folding/the collision/the 'movements' take place?) and to inquire, instead, into the earlier history of Appalachian tectonism. First, however, this question: is there any profit to be had by trying to identify a Variscan front in North America? The worth of such a concept is already doubtful within Europe (see below). To which American structure or structures should the label be applied? If the western limit of deformation in the northern Appalachians is at the Taconic front, can that, with propriety, be regarded as a Variscan front also? If one thinks to treat the western limit of Alleghanian deformation in the southern Appalachians (Woodward 1957) as a Variscan front (cf. Read & Watson 1975, fig. 3.10) is it permissible to regard an Alleghanian front as a continuation of a Taconic front? Further, given recently acquired understandings of the long-term development of the southern Appalachians, and given long-standing opinion that Caledonian and Variscan effects coincide there (King 1969; Rodgers 1970, 1971), would it be more worthwhile to treat the western limit of deformation as a Caledonian-Variscan front rather than merely a Variscan front? These questions touch on the more substantial problem of the total amount of shortening of the southern Appalachians as compared to the total amount of shortening of the northern Appalachians. Williams (1980) raised this matter in response to a new interpretation of the southern Appalachians advanced by Cook et al. (1979) and Harris & Bayer (1979). He arrived at the suggestion that although structural shor-

tening appears to be greater in the northern Appalachians, Newfoundland structures are essentially rooted. Williams & Hatcher (1982), however, have advanced a view of the northern Appalachians in which the assemblage of tectonic units is assumed to be incoherent: the 'suspect terrane' approach. This model is at variance with Williams's (1979) who matched tectonic units in Newfoundland with tectonic units in the British Isles (cf. Kennedy 1979). Not only did he identify British equivalents (Humber-Hebrides, Dunnage-Dundee, Gander-Greenore, Avalon-Anglesey zones), he set out a series occurring in the same order in the British Isles as in Newfoundland. Figure 1, encouraged by what Lefort & Haworth (1979) have discovered in the offshore regions, makes the suggestion, first that the Palaeozoic geology in NE North America is less chaotic than Williams and Hatcher's (1982) opinions imply, secondly that the Gander and Avalon units may yet be seen to deserve comparison with cases in Europe (although not necessarily the cases Williams 1979 recommends) and thirdly that the Narragansett basin geology in particular (McMaster, de Boer & Collins 1980; Dallmeyer 1982; Mosher & Rast, this volume), with a first set of structures directed north-westward, a second set directed in the opposite sense and 'Variscan' metamorphism and plutonism, may deserve comparison with the Saxothuringian zone (Behr et al. 1982; Franke, this volume). If the comparison holds (fuller information on the early history of the Narragansett geology would serve as a basis for testing its worth) the suggestion of comparability with the Saxothuringian zone should not immediately be greeted as an opportunity to propose a correlative of the Moldanubian zone on one side of the Narragansett basin and a Rhenohercynian zone (plus a sub-Variscan foredeep) on the other. If Fig. 1 is worthwhile, it is not surprising that Schenk (1971) was able to find in the Meguma a resemblance to the geology of NW Africa. That the geology of the northern Appalachians is much deformed is not to be denied. Figure 1 offers the view that the region is less deformed than some might think. It offers an example of a problem which will continue to be discussed during time when more and more seismic reflection data become available: deep-reaching structures, yes, but how much transposition has been effected on them? The southern Appalachians (Hatcher 1978; treated in Fig. 1 as the Appalachians south of an extrapolation of the Kelvin fracture zone--cf. Arthaud & Matte 1977) are greatly deformed as is best indicated by the suggestion

Northern margins o f the Variscides that the crystalline Piedmont has been thrust over a sedimentary sequence which may include rocks originally sited at the NW margin of the Iapetus Ocean (Cook et al. 1979; Harris & Bayer 1979, note also Iverson & Smithson 1983). Radiometric datings and their geological settings within the southern Appalachians are informatively discussed by Dallmeyer (1979). Sinha & Zietz (1982) serve the outside observer well by putting together a great deal of information in their proposal of a 'Hercynian arc' which runs from Maryland to Georgia. One would, nevertheless, wish to be better informed on what might be called 'pre-Hercynian tectonism. What, for example, was the original character of the Carolina slate belt, which Rodgers (1972) suspected to be a southern Appalachian representation of the Avalon 'belt' and which Long (1979) has interpreted as the deformed state of a rift structure first established in late Precambrian to early Palaeozoic time? Or, in what circumstances did the Brevard zone first acquire an identity? If the Brevard zone ends downward in the main sole thrust that underlies the Blue Ridge and Inner Piedmont (Cook et al. 1979) and if thrusting had already begun in Ordovician time (Cook et al. 1979), do we reject any possibility that the Brevard zone may have had a preOrdovician history? It would be interesting to explore further the thought that analogues of some of the early tectonic features of the southern Appalachians are on display, in a less 'deracin6' condition, in central and western Europe. We are faced with the possibility of making a choice between the suggestion that the southern Appalachians is a paradigm, from which we can derive proposals on the likely state of deep crustal structure in Europe, or the alternative suggestion that western Europe is a much less deformed region, which may help us to understand the primitive state of the southern Appalachian structure. Others' questions demand attention. King's (1975): 'What was the configuration of the ancient (pre-Palaeozoic) southern margin of North America?' is an especially good one. In attempting to deal with it (King 1975, figs 1, 2, 3; cf. Dewey 1982, fig. 37) there will be some obligation to take account of the relatively undeformed state of Lower Palaeozoic rocks in Florida, which includes an analogue of the Armorican Quartzite (Rodgers 1970; Lefort & Van der Voo 1981; Smith 1982; Matthews & Ford 1984). Likewise, if comparisons with, say, the long-term history of the Saxothuringian zone may prove to be of service in unravelling the early tectonic history of the Piedmont reg-

79

ion of the Appalachians, it will be wise to bear in mind the fact (often overlooked in discussions of deformation and metamorphism in central Europe) that the Prague syncline, no great distance away from the Saxothuringian zone, contains the non-metamorphic, relatively undeformed early Palaeozoic sequence long known as the 'Barrandian'--see a succinct account of that region in Sch6nenberg & Neugebauer (1981). Lefort & Van der Voo's (1981) reconstruction of tectonic relationships in the North Atlantic region is attractive in several ways. They refer not only to dextral shear, as discus~ sed here in what is regarded as the SE Plate in the Iapetus System, but also a sinistral shear in the NW plate. They address the question of identifying arcuate major structures in the Appalachians (see, in the same connection, Wintsch & Lefort, this volume). One point on which they appear to err is their attribution of late tectonic effects in the southern Appalachians to closure of a 'Theic' Ocean (McKerrow & Ziegler 1972): an ocean which for the most part appears to rely on the need to accommodate a single aberrant late Devonian palaeopole from North Africa (from the Msissi Norite in Morocco: Hailwood 1974). Figure 1 takes account of a wider range of evidence, which suggests early tectonic character in common in North Africa and Europe and makes no suggestion of that separateness of North Africa (or alternatively North Spain--Lefort & Van der Voo do not fully commit themselves to a choice; but see Matthews 1984b) from Europe which Lefort & Van der Voo's (1981) preferred evidence implies. The question of an African association in the Palaeozoic geology of central and western Europe is germane to present arguments and deserves brief discussion.

African connections of Palaeozoic geology The pre-Devonian pattern of European geology shown in Fig. 1 has an African 'allure': arcuate belts of rift-dominated structure enclose distinct domains in central Europe, western Europe and NE North America. Matthews (1984b) has noted that the domain encompassing SW England, West France and much of Iberia includes a striking number of ultrabasic intrusions. The pattern appears to have been established in late Precambrian time, and to have continued in development during the time when these 'domains' were part of the SE plate in the Iapetus scheme. If it is accepted that an appear-

80

S. C. Matthews

ance of resemblance to an African style of regional structure exists, one may ask why this should be so. The likeliest explanation is that what we now call central and western Europe has a continental crust which is relatively young (Jfiger 1977, 1979; Vidal 1977; van Breemen et al. 1982) and which was produced during a late Precambrian-early Palaeozoic reproduction of the circumstances in which crust grew in early Proterozoic time (Windley 1977): thin crust was exposed to relatively high heat flow. These late Precambrian structures in Europe and NE America are variously called 'Cadomian', 'Assyntian', 'Baikalian' (see editorial comment in Murawski 1981) or 'Avalonian'. It is possible that the best collective name for them is 'Pan-African' (Kennedy 1964). One implication of Fig. 1 is that PanAfrican crustal features may be found all of the way northwards to the Iapetus suture. A permissible inference is that in what we call Europe, certain of these Pan-African structures encountered a further accession of heating in mid- to late Palaeozoic time. It is an accident of this kind that causes, say the Saxothuringian zone in central Europe to appear in some respects different from a late Precambrian, Pan-African belt farther south such as, for example, the Damara orogen (Martin & Porada 1977a,b).

Conclusions It is a good working principle that the simplest explanation is to be preferred for as long as it

remains available as an option. No evidence available at present of the North Atlantic region debars the suggestion that only one major oceanic development (Iapetus) was involved in the late Precambrian-late Palaeozoic geology of the region. What has been under discussion here is largly the geology of the SE plate, i.e. predominantly intraplate geology, with an exceptional case to be recognized in south Brittany where a short-lived breaching of continental crust during late Silurian-early Devonian time is recorded. If Caledonian and Variscan are to be regarded as parts of one tectonic scheme--an interpretation consistently encouraged by the arguments set out a b o v e - - t h e n the consequences of Iapetus closure and continent-continent collision, from Scandinavia to the southern Appalachians and from Silurian to Permian time, are as sketched in Fig. 2. Figure 2 represents deformation superimposed on the pattern shown in Fig. 1 (in which, it should be repeated, only the southern Appalachians are shown in their final state). One takes the evidence of dextral shear in the region to suggest a couple, expressed as southeastward directed early thrusting in Scandinavia (Gee 1975, 1980; Str6mberg 1981) and north-westward directed late thrusting where the oblique closure had a late effect in the southern Appalachians. Western Europe, included in the 'anomalously straight' part of a Caledonide reassembly by Phillips, Stillman & Murphy (1976) and identified as carrying the middle elements in Arthaud & Matte's (1977) slightly mistaken (see comments above) proposal of a 'late Hercynian' A p p a l a c h i a n - U r a l shear sys-

FIG. 2. Sketch map of the North Atlantic region (Le Pichon restoration) to suggest: (1) major thrusting, directed south-eastward, in Scandinavia during Silurian/?Devonian time; (2) major thrusting, directed north-westward, in central Europe during Devonian-early Carboniferous time; (3) lack of any thrusting on a comparable scale in western Europe (a 'nodal region') and possibly north-eastern North America also; (4) major thrusting, effective until Carboniferous-Permian time, in the southern Appalachians. The whole assembly, thrusting plus major fractures (see Fig. 1) can be regarded as a couple, active from Silurian to Permian, during progressive south-westward closure of the lapetus Ocean.

Northern margins o f the Variscides tem, appears in Fig. 2 as a nodal region, unlikely to include major thrusting comparable with what is known in either Scandinavia or the southern Appalachians.

Variscan front? The question of a Variscan front, seen against such a background, appears trivial. The idea has arisen in a region where, as comments above suggest, evidence in favour of major mid- or late Palaeozoic thrusting is at a minimum. More than that, the idea is misconceived. It suggests that a line be drawn through four regions:

S W Ireland What has been identified as the Variscan front in SW Ireland is, according to Matthews etal. (1983), reverse faulting on the northern side of a doming produced during Carboniferous time in the course of development of an inversion structure. The heating is not Variscan h e a t i n g - - I r e l a n d is far removed from the main locus of Variscan heating discussed above. The growth of the whole inversion structure, in Matthews et al. (1983) view, was triggered off by a mantle delamination whose inception dates from the time of the collision that closed the Iapetus Ocean. The reverse faulting fades eastward as the doming fades. It does not continue into any structure given the label 'Variscan front' farther east.

Bristol Channel What has been called the Variscan front here (Matthews 1974, 1984b) is the deformed state of one flank or a n o t h e r - - o p i n i o n v a r i e s - - o f a rift structure which was in being (Fig. 1) long before Variscan deformation proceeded. Whichever flank one chooses, this 'Variscan front' does not continue into any structure given the same name farther east. Kenolty et al. (1981), it may be noted, have reported the interesting results of two seismic reflection profiles laid out in an area east of the Mendips. One is entitled to ask why they chose to identify any of the reflectors involved as the Variscan front.

Belgium Belgium has a strikingly well-developed 'Var-

81

iscan front': the Faille due Midi and associated structures. The problem here is that although it has what some would call a Variscan front, Belgium has very little of the geological character that one normally associates with the 'Variscan fold belt' as seen farther east, in Germany for example (Matthews 1984c).

Germany If left to themselves, G e r m a n geologists would probably not feel obliged to invent a Variscan front. If acquiescence in communautaire endeavours were to bring them to make a proposal, they might point to the E n n e p e St6rung (Thome 1970) where the balance of the U p p e r Carboniferous succession changes, or else cite the fact that in more or less the same neighbourhood Ahrendt, Hunziker & Weber's (1978) evidence of metamorphism fades. If the latter criterion were adopted, the boundary would be a limit of heating as indicated by one basis of measurement. Such an elusive boundary would have some of the qualities of an ignis fatuus. Its use in defining a Variscan front might in that case be thought to be largely appropriate. These comments on the tectonic setting of the Variscides in the North Atlantic region treat Caledonian and Variscan as parts of one chapter of tectonism. If they are to any extent valid, there may be a case for redrawing the Tectonic maps of Europe. Meanwhile, Figs 1 and 2 are available as bases for discussion. The regions they represent include some of the longest studied, most closely studied and most discussed features of the earth's crust. Given the present state of our understanding of the problems that arise, it is fitting that Figs 1 and 2 should appear to be crude.

ACKNOWLEDGMENTS: I am grateful to colleagues in several parts of Europe who have introduced me to their local geology and have shown me the ways in which they have learned to think about geology. I am particularly grateful to friends in Dublin, G6ttingen and Uppsala who have made it possible for me to prepare this account of some of my own views. Monica Siewertz (Uppsala) has kindly produced the typescript.

References AHRENDT, H., HUNZIKER,J. C. & WEBER, K. 1978. K/Ar-Altersbestimmung an schwachmetamorphen Gesteinen des Rheinischen Schiefergebirges. Z. dt. geol. Ges. 129, 229-47. ARTHAUD, F. & MATI'E, Ph. 1977. Late Paleozoic

strike slip faulting in southern Europe and northern Africa: result of right-lateral shear zone between the Appalachians and the Urals. Bull. geol. Soc. Am. 88, 1305-20. AUTRAN, A., BRETON, J.-P., CHANTRA[NE,J. CHIRON,

82

S. C. Matthews

J. C., GROS, Y. & ROGER, P. 1980. Introduction tectonics in the crystalline southern la carte tectonique de la France/t 1/1 000 000. Appalachians: COCORP seismic reflection proMdm. Bur. Rech. Gdol. Min. 110, 52 pp. filing of the Blue Ridge and Inner Piedmont. & COGNE, J. 1980. La zone interne de l'orog6ne Geology, 7, 563-7. varisque dans l'Ouest de la France et sa place DALLMEYER, R. D. 1979. 4~ dating: princidans le d6velopement de la chaine hercynienne. ples, techniques and applications in orogenic terIn: COGNE, J. & SLANSKY,M. (eds) Gdologie de ranes. In: JAGER, E. & HUNZIKER, J. C. (eds) l'Europe, du Prdcambrien aux Bassins SddimenLectures in Isotope Geology. 77-104, taires post-Hercyniens. 90-111, Colloque C6, Springer-Verlag, Berlin. 266me ICG, Paris, BRGM, Villeneuve d'Ascq. 1982. 4~ ages from the Narragansett BARD, J. P., BURG, J. P., MATTE, P. & RIBEIRO, A. Basin and southern Rhode Island basement ter1980. La cha~ne hercynienne d'Europe occidenrane: their bearing on the extent and timing of tale en termes de tectonique des plaques. In: Alleghanian tectonothermal events in New EngCOGNI~, J. & SLANSKY, M. (eds) G~ologie de land. Bull. geol. Soc. Am. 93, 1118-30. l'Europe, du Prdcambrien aux Bassins SddimenDEN TEX, E. 1981. Basement evolution in the northtaires post-Hercyniens, 233-46. Colloque C6, ern Hesperian Massif. A preliminary survey of 266me ICG, Paris, BRGM, Villeneuve d'Ascq. results obtained by the Leiden Research Group. BASSETT, M. G., CHERNS, L. & KARIS, L. 1982. The Leidse geol. Med. 52, 1-21. R6de Formation: early Old Red Sandstone DEWEY, J. F. 1982. Plate tectonics and the evolution facies in the Silurian of J~imtland, Sweden. Sver. of the British Isles. J. geol. Soc. London, 139, geol. Unders. Serie C, 793, 1-24. 371-412. BEHR, H.-J. 1961. Beitr~ige zur petrographischen und & BURKE, K. C. A. 1973. Tibetan, Variscan and tektonischen Analyse des s/ichsischen Precambrian basement reactivation: products of Granulitgebirges. Freiberger Forschfi. Cl19, continental collision. J. Geol. 81, 683-92. 1-64. DORNSIEPEN, U. F. 1978. Ein Uberblick fiber die 1978. Subfluenz-Prozesse im Grundgebirgseurop/iischen Varisziden. Z. dt. geol. Ges. 129, Stockwerk Mittel-europas. Z. dt. geol. Ges. 129, 521-42. 283-318. 1979. Rb/Sr whole rock ages within the Euro- - - , ENGEL, W. & FRANKE, W. 1982. Variscan pean Hercynian. Krystalinikum, 14, 33-49. Wildflysch and nappe tectonics in the Saxo- ELLENBERGER, F. & TAMAIN, A. L. G. 1980. Hercythuringian Zone (northeast Bavaria, west Gernian Europe. Episodes, 22-7. many). Am. J. Sci. 282, 1438-70. FLOYD, P. A. 1982. Chemical variation in Hercynian - - 9 WALLISER, O. H. & WEBER, K. 1980. The basalts relative to plate tectonics. J. geol. Soc. development of the Rhenohercynian and SaxoLondon, 139, 505-20. thuringian zones of the mid-European - - , EDER, W., ENGEL, W. & LANGENSTRASSEN,F. Variscides. In: COGNF., J. & SLANSKV, M. (eds) 1978. Main aspects of geosynclinal sedimentaGOologie de l'Europe, du Prdcambrien aux Bastion in the Rhenohercynian Zone. Z. dt. geol. sins SOdimentaires post-Hercyniens, 77-89. ColGes. 129, 201-16. loque C6, 266me I.C.G. Paris. FUCHTBAUER, H. 1963. Petrographische UnterBERNAROOVA, E. & CHAB, J. 1968. Pr~iassyntischer suchungen des Unterkarbons und Devons der Schiefer als klastisches Material in jungBohrung Mfinsterland 1. Fortschr. Geol. proterozoischen Grauwacken im NW-Teil des Rheinld. Westf 11,353-64. Barrandiums. Geologie, 17, 753-75. GARDINER, P. R. R. & SHERIDAN, D. J. R. 1981. BIRD, P. 1978. Initiation of intracontinental subducTectonic framework of the Celtic Sea and adjation in the Himalaya. J. geophys. Res. 83, cent areas with special reference to the location 4975-87. of the Variscan Front. J. struct. Geol. 3, 317-31. 1979. Continental delamination and the ColGEE, D. G. 1975. A tectonic model for the central orado Plateau. J. geophys. Res. 84, 7561-71. part of the Scandinavian Caledonides. Am. J. Sci. 275A, 468-515. BRINKMANNI R. 1948. Die Mitteldeutsche Schwelle. Geol. Rdsch. 36, 56-66. 1980. Basement-cover relationships in the cenBROMLEY, A. 1975. Tin mineralization of Western tral Scandinavian Caledonides. Geol. FOr. Europe: is it related to crustal subduction? Stockh. FOrh. 102, 455-74. Trans. lnstn min. Metall. 84, B28-30. GIESE, P. 1978. Die Krustenstruktur des Varistikums BURG, J. P. & MATTE, P. 1978. A cross section und das Problem der Krustenverkiirzung. Z. dt. through the French Massif Central and the scope geol. Ges. 129, 513-20. of its Variscan Geodynamic Evolution. Z. dt. GRAHAM, C. M. 1976. Petrochemistry and tectonic geol. Ges. 129, 429-60. significance of Dalradian metabasaltic rocks of COGNE, J. 1976. La chaine hercynienne ouestthe SW Scottish Highlands. J. geol. Soc. London, europ6ene correspond-elle/l un orog6ne par col132, 61-84. lision? Propositions pour une interpr6tation 9 GUENNOC,P. i 978. Structure et 6volution g6ologique g6odynamique globale. In: Ecologie et Gdologie de la pente continentale d'un secteur de l'Atlande l'Hirnalaya. Colloques int. Cent. natn. Rech. tique nord-est: de la terrasse de M6riadzek scient. 268, 111-29. l'6peron de Goban. ThOse, 3dine cycle. Brest, COOK, A., DENNIS, S., BROWN, L., KAUFMAN, S., 95 pp. OLIVER, J. & HATCHER, R. 1979. Thin-skinned HAILWOOD, E. A. 1974. Palaeomagnetism of the -

-

-

-

-

-

-

-

N o r t h e r n m a r g i n s o f the Variscides Msissi Norite (Morocco) and the Palaeozoic reconstruction of Gondwanaland. Earth planet. Sci. Lett. 23, 376-86. HARRIS, L. D. & BAYER, K. 1979. Sequential development of the Appalachian orogen above a master decollement--a hypothesis. Geology, 7, 568-72. HATCHER, R. D. (Jr) 1978. Synthesis of the southern and central Appalachians, U.S.A. Pap. geol. Surv. Can. 78-13, 149-57. IVERSON, W. P. & SMITHSON, S. B. 1983. Reprocessing and reinterpretation of C O C O R P Southern Appalachian profiles. Earth planet. Sci. Lett. 62, 75-90. JAGER, E. 1977. The evolution of the Central and West European continent. In: COGNE, J. (ed.) La Chafne Varisque d'Europe Moyenne et Occidentale, 227-39. CNRS, Rennes. 1979. Evolution of the European continent. In: JAGER, E. & HUNZIKER, J. C. (eds) Lectures in Isotope Geology, 222-4. Springer-Verlag, Berlin. KENNEDY, M. J. 1979. The continuation of the Canadian Appalachians into the Caledonides of Britain and Ireland. In: HARRIS, A. L., HOLLAND, C. H. & LEAKE, B. E. (eds) The Caledonides of the Bristol Isles--reviewed, 3-18. Geological Society of London. KENNEDY, W. Q. 1964. The structural differentiation of Africa in the Pan-African (_+ 500 m.y.) tectonic episode. Eighth Ann. Rep. Scientific Results, Session 1962-1963, 48-9. Research Institute of African Geology, University of Leeds. KENOLTY, N., CHADWICK,R . A., BLUNDELL, D. J. & BACON, M. 1981. Deep seismic reflection survey across the Variscan Front of southern England. Nature, 293, 451-3. KING, P. B. 1969. The tectonics of North America-a discussion to accompany the Tectonic Map of North America. Scale 1 : 5,000,000. Prof. Pap. U.S. geol. Surv. 628, 94 pp. 1975. Ancient southern margin of North America. Geology, 2, 732-4. KOSSMAT, F. 1921. Die mediterranen Kettengebirge in ihrer Beziehung zum Gleichgewichtszustande der Erdrinde. Abh. math. Phys. KI. siichs. Akad. Wiss. 38(2), 1-62. 1927. Gliederung des variszischen Gebirgsbaues. Abh. siichs, geol. Landesamts, 1, 39 pp. LEFORT, J. P. 1979. The Ibero-Armorican Arc and the Hercynian orogeny in western Europe. Geology, 7, 384-8. -& HAWORIH, R. T. 1979. The age and origin of the deepest correlative structures recognized off Canada and Europe. 7ectonophys. 59, 13950. & RIBEIRO, A. 1980. La faille Porto-BadajozCordue a-t-elle controll6 l'ouverture de l'oc6an sud-armoricain? Bull. Soc. gdol. Fr. (76me s6r.), 22, 455-62. -• VAN DER VOO, R. 1981. A kinematic model for the collision and complete saturing between Gondwanaland and Laurussia in the Carboniferous. J. Geol. 89, 537-50.

83

LONG, L. T. 1979. The Carolina slate belt-evidence of a continental rift zone. Geology, 7, 180-4. LOTZE, F. 1945. Einige Probleme der Iberischen Meseta: zur Gliederung der Varisciden der Iberischen Meseta. Geotekt. Forsch. 6, 1-14, 78-92. MARTIN, H. & PORADA, H. 1977a. The intracratonic branch of the Damara orogen in southwest Africa. I. Discussion of geodynamic models. Precambrian Res. 5, 311-38. -& ~ 1977b. The intracratonic branch of the Damara orogen in southwest Africa. II. Discussion of relationships with the Pan-African mobile belt system. Precambrian Res. 5, 339-57. MATrHEWS, S. C. 1974. Exmoor Thrust? Variscan Front? Proc. Ussher Soc. 3, 82-94. 1978. Caledonian connexions of Variscan tectonism. Z. dt. geol. Ges. 129, 423-8. 1984a. Pre-Mesozoic geology of the Celtic Sea and adjacent regions. J. Earth Sci. R. Dublin Soc. (in press). 1984b. Questions concerning the Palaeozoic geology of northern Spain Neues Jb. Geol. Paliiont. Abh. (in press). 1984c. Tectonic antecedence of the Variscan geology of Belgium. Bull. Soc. beige GdoL (Delmer-Legrand Volume) . (in press). & FORD, I. H. 1984. Armorican, Hercynian, Variscan and certain other foreign terms. Proc. Ussher Soc. (in press). - - . , NAYLOR, D. & SEVASTOPULO, G. D. 1983. Palaeozoic sedimentary sequence as a reflection of deep structure in southwest Ireland. Sediment. Geol. 34, 83-95. MCMASTER, R. L., DE BOER, J. & COLLINS, B. P. 1980. Tectonic development of southern Narragansett Bay and offshore Rhode Island. Geology, 8, 496-500. MCKERROW, W. S. & ZIEGLER, A. M. 1972. Palaeozoic oceans. Nature, 240, 92-4. MEISCHNER, K. D. 1964. Allodapische Kalke, Turbidite in Riff-nahen Sedimentations-Becken. In: BOUMA, A. H. & BROUWER, A. (eds) Develop. Sedimentol. 3, 156-91. Mt3LLER, E. 1978. Ergebnisse der Forschungsbohrungen Diippenweiler/Saar. Nachr. dt. geol. Ges. 19, 32-3. MURAWSKI, H. 1981. Problems of the Variscides in Central Europe. Geotectonics, 15, 479-91. NICHOLSON, R. 1979. Caledonian correlations: Britain and Scandinavia. In: HARRIS, A. L., HOLLAND, C. H. & LEAKE, B. E. (eds) The Caledonides of the British Isles--reviewed, 3-18. Geological Society of London. OLIVEIRA, J. T., HORN, M. & PAPROTH, E. 1979. Preliminary note on the stratigraphy of the Baixo Alentejo Flysch Group, Carboniferous of Southern Portugal and on the palaeogeographic development, compared to corresponding units in Northwest Germany. Comunr Servs. geol. Port. 65, 151-68. PHILLIPS, W. E. A., STILLMAN, C. J. & MURPHY, T. 1976. A Caledonian plate tectonic model. J. geol. Soc. London, 132, 579-609. PRUVOST, P. 1949. Les mers et les terres de Bretagne -

-

-

-

84

S. C. M a t t h e w s

aux temps pal6ozoiques. Annls Hdbert Haug, 7, 345-60. RAST, N. & GRANT, R. 1973. Transatlantic correlation of the Variscan-Appalachian Orogeny. Am. J. Sci. 273, 572-9. & 1977. Variscan-Appalachian and Alleghanian deformation in the Northern Appalachians. In: COGNt~, J. (ed.) La Chaine Varisque d'Europe Moyenne et Occidentale, 583-6, CNRS, Rennes. ~, RAST, D. E. & KENNEDY, M. J. 1975. Plate tectonic significance of repeated tectonic trends in eastern Canada. Nature, 258, 61-2. READ, H. H. & WATSON, J. V. 1975. Introduction to Geology. Vol. 2: Earth History. Part 1: Later Stages o f Earth History, Macmillan, London. 371 pp. RIBE1RO, A. 1974. Contribution h l'6tude tectonique de Tras-os-Montes oriental. Mems Servs geol. Port. 24, 177 pp. R1ES, A. C. & SHACKLETON,R. M. 1976. Patterns of strain variation in arcuate fold belts. Phil. Trans. R. Soc. A, 283, 281-8. RODGERS, J. W. 1970. The tectonics o f the Appalachians. Wiley, New York. 271 pp. 1971. The Taconic Orogeny. Bull. geol. Soc. Am. 82, 1141-78. 1972. Latest Precambrian (post-Grenville) rocks of the Appalachian region. Am. J. Sci. 272, 507-20. SCHENK, P. W. 1971. Southeastern Atlantic Canada, northwestern Africa and continental drift. Can. J. Earth. Sci. 8, 1218- 51. SCHONENBERG, R. & NEUGEBAUER, J. 1981. Einfiihrung in die Geologie Europas. Rombach, Freiberg. 340 pp. SCHREYER, W., ABRAHAM, K. & BEHR, H. J. 1975. Sapphirine and associated minerals from the kornerupine rock of Waldheim, Saxony. Neues Jb. Miner. Geol. Paldont. Abh. 126, 1-27. SHACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in SW England. J. geol. Soc. London, 139, 533-41. SINHA, A. K. & ZIETZ, I. 1982. Geophysical and geochemical evidence for a Hercynian magmatic arc, Maryland to Georgia. Geology, 10, 593-6. SMITH, D. L. 1982. Review of the tectonic history of the Florida basement. Tectonophys. 88, 1-22. STtLLE, H. 1951. Das mitteleurop~iische variszische Grundgebirge im Bilde des gesamteurop/iischen. Beih. geol. Jb. 2, 138 pp. STROMBERG, A. G. B. 1976. A pattern of tectonic zones in the western part of the East European platform. Geol. FOr. Stockh. FOrh. 98, 227-43. 1981. The European Caledonides and the Tornquist Lineament. Geol. FOr. Stockh. FOrh 103, 167-71. SUESS, E. 1887. Ober unterbrochene Gebirgsfaltung Sber. preuss. Akad. Wiss. 94, 111-7. 1888. Das Antlitz der Erde, 2er Band. Tempsky, Freytag. 703 pp. ~['EICHMOLLER, R., TEICHMOLLER, R. & WEBER, K. 1979. Inkohlung und Illit-Kristallinit~it. Verg-

leichende Untersuchungen im Mesozoikum und Pal~iozoikum von Westfalen. Fortschr. Geol. Rheinld Westf 27, 201-76. THOME, K. N. 1970. Die Bedeutung der EnnepeSt6rung ffir die Sedimentations und Faltungsgeschichte des Rheinischen Schiefergebirges. Fortschr. Geol. Rheinld Westf 17, 757-808. VAN BREEMEN, O., AFTALION, M., BOWLS, D. R., DUDEK, A., MiSAR, Z., POVONDRA,P. & VRANA, S. 1982. Geochronological studies of the Bohemian massif, Czechoslovakia, and their significance in the evolution of Central Europe. Trans. R. Soc. Edinb., Earth Sci. 73, 89-108. VAN DER MEER MOHR, C. G., KUIJPER, R. J., VAN CALSTEREN, P. W. C. & DEN TEX, E. 1981. The Hesperian Massif: from lapetus aulacogen to ensialic orogen. A model for its development. Geol. Rdsch. 79, 459-72. VIDAL, P. 1977. Limitations isotopiques/l l'gge et l'evolution de la crofite continentale en Europe moyenne et occidentale. In: COGNE, J. (ed.) La Chatne Varisque d'Europe Moyenne et Occidentale. C.N.R.S., Rennes. WEBZR, K. 1978. Das Bewegungsbild im Rhenoherzynikum--Abbild einer varistischen Subfluenz. Z. dt. geol. Ges. 129, 249-81. 1981. The structural development of the Rheinische Schiefergebirge. Geologie Mijnb. 60, 149-59. WEGENER, A. 1924. The Origins of Continents and Oceans (translation of the German 3rd edn). Methuen, London. 212 pp. WEJNHOLD, G. 1977. Zur pr/ivaristichen Vererzung im Erzgebirgskristallin aus der Sicht seiner lithofaziellen und geotektonisch-magmatischen Entwicklung whhrend der assyntischkaledonischen ,~ra. Freiberger ForschH)2. C 320, 53 pp. WIENHOLZ,R., BAUMANN,L. & HOFMANN, J. 1982. Einige neue Erkenntnisse zum geologischen Bau und zur Lagerst/ittenbildung im Erzgebirge. Z. angew. Geol. 28, 418-26. WILLIAMS, H. 1979. Geological development of the Northern Appalachians: its bearing on the evolution of the British Isles. In: BOWLS, D. R. & LEAKE, B. E. (eds) Crustal Evolution in Northwest Britain and Adjacent Regions, 1-21, Seel House Press, Liverpool. 1980. Comments and replies on 'Thin-skinned tectonics in the crystalline southern Appalachians; COCORP seismic-reflection profiling of the Blue Ridge and Piedmont' and 'Sequential development of the Appalachian orogen above a master decollement--a hypothesis'. Geology, 8, 211-2. & HATCHER, R. D. (Jr) 1982. Suspect terranes and accretlonary history of the Appalachian orogen. Geology, 10, 530-6. WINDLEY, B. F. 1977. The Evolving Continents. Wiley, London. 385 pp. WOOD, D. S. 1974. Ophiolites, m61anges, blueschists and ignimbrites; early Caledonian subduction in Wales? In: Doaq', R. H. (Jr) & SHAVER,R. H. (eds) Modern and Ancient Geosynclinal

N o r t h e r n m a r g i n s o f the Variscides Sedimentation; Problems of Palinspastic Restoration. Spec. Publ. Soc. econ. Paleont. Miner., Tulsa, 19, 334-44. WOODWARD,H. P. 1957. Chronology of Appalachian folding. Bull. Am. Ass. Petrol. Geol. 41, 2312-27. ZIMMERLE, W. 1976. Petrographische Beschreibung und Deutung der erbohrten Schichten. In: Die Tiefbohrung Saar 1. Geol. Jb. Reihe A, 27, 291-305.

85

ZWARX, H. J. 1967. The duality of orogenic belts. Geologie Mijnb. 46, 283-309. 1976. Regional metamorphism in the Variscan orogeny of Europe. Nova Acta Leopoldina, 45, (Kossmat-Symposion), 224, 361-7. & DORNSlEPEN, U. F. 1978. The tectonic framework of Central and Western Europe. Geologie Mijnb. 57, 627-54.

S. C. MXVIHEWS(deceased 5 May 1983), Enheten f6r Paleobiologi, Box 564, S-751 22 Uppsala, Sweden.

An interpretation of the Variscan tectonics of SW Britain M. P. Coward & S. Smallwood SUMMARY: The structure of SW Britain is interpreted in terms of thin-skinned tectonics, producing a NNW verging fold and thrust zone. In south Wales the thrusts involve high level imbricate stacks showing some 20-40% shortening with a shallow floor thrust in Namurian-Westphalian strata, cut to the south by deeper level folds and thrusts with a basal decoupling zone in or below the Lower Palaeozoic rocks. In Devon and Cornwall the Upper Palaeozoic rocks show a gently dipping, widespread and locally intense cleavage related to the progressive development of large thrust sheets and associated inclined to recumbent folds. In north Cornwall these northward verging structures have been redeformed and often completely overprinted by a large south verging backthrust. Some of this backthrust movement may be offset, by a lateral ramp, to a much lower decoupling zone, producing a major backfold in south Devon. The thrust transport direction is to the NNW as determined from mineral lineations, maximum extension directions, fold and thrust traces and tear faults. Fold hinges are generally normal to the transport direction but are locally rotated into the NNW trend in more intensely deformed zones. In north Devon and south Wales, the folds are oblique suggesting a component of differential movement. Considering SW Britain and southern Ireland the fold trends are arcuate and there is often extension parallel to the fold axes. These phenomena are interpreted as due to the obliquity of the thrusting and the Variscan front. Pre-tectonic facies changes along the front hinder lateral fault propagation giving sticking points at lateral tips which become poles of rotation. Deformation affected only the upper crust; there was no major crustal thickening during Variscan tectonics though the crust has been locally thinned probably by a factor of 2 before onset of compression in Devonian times. Deformation was diachronous with a slow displacement rate. Sedimentation occurred in an advancing series of fore-deeps ahead of the deformed and thickened zone. Estimates of shortening across the thin-skinned zone are about 50%, that is 150 km across SW England. The crust beneath the thin skin must extend back beneath the English Channel and northern France. The Variscan structures of SW Britain (south Wales, Devon and Cornwall) have been interpreted in terms of a thin-skinned fold and thrust zone (Shackleton, Ries & Coward 1982) and as the northern margin of a strike-slip orogen (Badham 1982). On a larger, European, scale, they form part of the external (RhenoHercynian) zone which can be traced f r o m southern Ireland through Britain into northern France, Belgium and central G e r m a n y (Autran et al. 1980). Throughout this zone, the rocks show only low grade (greenschist or lower grade) metamorphism; the higher grade rocks with syntectonic granites occur to the south, in the internal crystalline zones (cf. Autran et al. 1980). A granite batholith underlies the Variscides of SW England, but this is late to posttectonic, probably Permian in age (Dodson & Rex 1971). The sediments involved in the Variscan tectonics of SW Britain range in age from Devonian to late Carboniferous. Olistoliths in the Devonian flysch deposits of south Cornwall contain shelf facies Ordovician, Silurian and Lower Devonian sediments, comparable with some of the Lower Palaeozoic sediments of north France (Leveridge 1974; Barnes, Andrews & Badham 1979). Apart from these

minor examples in Cornwall no other Lower Palaeozoic rocks are involved in the thrust tectonics of S W England. In Devon and Cornwall there are no thrust sheets of basement rocks as in the southern Appalachians, although in south Wales, at the margin of the belt, Lower Palaeozoic rocks occur in the cores of major anticlines. There are major sedimentary facies changes in Devonian and Carboniferous sediments. In SW Cornwall, the Devonian strata form a thick sequence of flysch-like deposits with local olistostromes (Barnes et al. 1979) but in south Devon and central Cornwall, Devonian sediments are of a shallow marine shelf facies with sandstones, shales and reef limestones (House et al. 1977). In north Devon and south Wales the Devonian strata include thick red terrestrial sandstones and shales (the Old Red Sandstone facies). The Carboniferous rocks show a similar change from flysch-like deposits in the south to shallow water limestones, sandstones and coal deposits in south Wales. Dewey (1982) and Leeder (1982) apply a crustal stretching model to explain the thick sequence of sediments in north D e v o n and estimate a stretching factor (/?) of about 2. In south Pembrokeshire, 89

90

M. P. C o w a r d & S. S m a l l w o o d

~S

p

I 50 km m~

~

arcuate trend probably more movement in west

& high ~9

fold facing direction thrust antiform

.fl~ synform

FIG. 1. (a) Map of SW Britain and southern Ireland showing the main structures of the Variscan front, the main zones of backthrusts and back-folds, the thrusts and folds of south Wales and their equivalents in southern Ireland, the lines of section shown in Fig. 6 and the locations of Figs 2 and 5, after Naylor et al. (1983), Hobson & Sanderson (1983), Kellaway & Hancock (1983). (b) Sketch of the arcuate form of structures related to the Variscan front (see text). (c) The position of the Variscan front across southern Britain and Ireland, after Dunning (1977) and Wallace (1983).

Carboniferous rocks show unconformities and onlap/offlap sequences, accompanied by widespread facies changes (Jenkins 1962; Sullivan 1965). The major thrusts marking the Variscan front follow these facies changes and hence presumably the boundary to the sedimentary basin. Facies variations mean that correlation across SW Britain is difficult, especially where palaeontological criteria are missing. The exact ages of many of the Devonian rocks of south Cornwall are unknown. Similarly the strata do not possess a widespread layer-cake stratigraphy and hence cross-sections are more difficult to restore. Within small areas, such as the south Pembrokeshire thrust zone, the strata may be considered approximately layer-cake allowing the construction of local balanced crosssections. The structures of SW Britain show many of the features of a thin-skinned foreland fold and thrust zone. There are no major mountains, no evidence of thickened crust (Holder & Bott 1971; Autran et al. 1980) and no high grade metamorphic rocks. The structure is characterized by the presence of both fore- and back-

thrusts (Fig. 1) and a range in fold facing directions (Sanderson & D e a r m a n 1973; Hancock 1973). The approximate position of the tectonic boundary to the Variscan, the Variscan front (Dunning 1977), is shown in Fig. l(c). Throughout most of southern Britain it has a W N W trend but is largely obscured by Mesozoic cover rock. A n orogenic front may be defined as the limit of the main zone of deformation and may be the outcrop trace of the lowermost thrusts as in Pembrokeshire (Fig. 1) or may be the limit of folding above a deeper decoupling zone where this zone does not intersect the topographic surface. This latter type occurs in southern Ireland (Fig. 1), where Variscan folding extends much further to the north and west than in the rest of Britain. The Hercynian structures on the margins of the belt in Pembrokeshire in south Wales are described first. This is followed by a summary of the structural geology of Devon and Cornwall and then a regional synthesis. Thrust term i n o l o g y follows that of recent papers by Elliott & Johnson (1980) and Butler (1982).

V a r i s c a n tectonics o f S W Britain

The northern margin of the Variseides: south Pembrokeshire This thrust belt is composed of two parts (Figs 2 and 3). To the north there is an imbricate zone developed in Upper Carboniferous rocks, above a shallow floor thrust which dips gently south. This zone is bounded to the south by the Johnston and Ritec thrusts (Fig. 2), south of which there are large-scale folds in Devonian and Lower Carboniferous rocks, with Lower Palaeozoic inliers in the fold cores (Hancock 1973; Hancock, Dunne & Tringham 1983). There is also a change in facing direction of the folds; the axial planes of the folds fan through the vertical, so that in the north, near the Johnston and Ritec thrusts, the folds face northwards but in the south, they face southwards (Hancock 1973; Hancock, Dunne & Tringham 1981). The east coast section

A cross-section along the east coast of Pembrokeshire is shown in Fig. 3(a). To the south of the Ritec thrust, large-scale folds in the Old Red Sandstone and Lower Carboniferous rocks are upright and have a wavelength of 1-2 km with locally parasitic structures on the limbs (see also Hancock et al. 1981). Cross-sections through the structures suggest a relatively deep decoupling level, at a few kilometres depth. The shales carry a weak cleavage and reduction spots in the red shales show nearly oblate strains, flattened in this cleavage. We consider these folds to have formed by a layer parallel

91

shortening and buckling process, rather than by ramp climb at depth. Further deformation caused the northern limb of one of the anticlines to fail along the Ritec thrust and the folded rocks to be thrust northwards some 500 m over mid- to late Carboniferous strata. This displacement may increase slightly to the west (Hancock et al. 1981) but the fault cannot be reliably traced to west of Milford Haven (Fig. 2). To the north, Namurian and Westphalian sediments are folded and form an imbricate stack (Fig. 3a); restored crosssections indicate about 45% shortening between Tenby and Saundersfoot. Major thrusts cut the section at Monkstone and Saundersfoot. A possible deeper level of decoupling is indicated by thrusts which outcrop in the Red Roses disturbance belt. The depth to the floor thrust increases to the south, although to restore the section the Ritec thrust must cut through this floor thrust to the Carboniferous imbricates. The imbricate thrusts are considered to have formed in piggyback fashion but these are cut by major out of sequence thrusts carrying lower level rocks in hanging walls (Fig.

3a). The west coast section

A synoptic cross-section along the west coast from St Ann's Head to north of Broad Haven is shown in Fig. 3(b). In the south, folds with a wavelength of about 1 km and an axial cleavage face south in Old Red Sandstone rocks. Lower Palaeozoic rocks occur in anticlinal fold cores and also on the hanging wall of the Musselwick

FiG. 2. Map of south Pembrokeshire showing the main fold and thrust zones. Thrusts are shown with teeth on the hanging wall. For more details see Hancock et al. (1981, 1983). Lines VW and XY are shown in Fig. 3.

92

M. P. C o w a r d & S. S m a l l w o o d

FIG. 3. (a) Simplified balanced cross-section along the east Pembroke coast, line VW of Fig. 2. ORS = Old Red Sandstone (Devonian). Horizontal scale = vertical scale. (b) Simplified synoptic section along the west Pembrokeshire coast; data are projected on to line XY. See also Hancock et al. (1983). Horizontal scale = vertical scale.

fault. Large-scale thrusts are rare, though some small thrusts grow out of synclinal fold cores, as at St A n n ' s Head. To the north, the large-scale folds are more upright and northward facing near the Benton fault and Johnston thrust. This thrust carries Precambrian rocks on its hanging wall and truncates structures in the U p p e r Carboniferous strata of the footwall. The Benton fault is a normal fault in the hanging wall of the Johnston thrust (Fig. 3b). North of the Johnston thrust, the Upper Carboniferous strata are cut by numerous imbricate thrusts. The floor thrust to this set outcrops near the base of the Westphalian at Settling Nose (Fig. 3b), but increases in depth to the south. Restored cross-sections indicate 25% shortening between the Johnston thrust and Settling Nose, but there is 30% shortening in a short section at Broad Haven. General

A piggyback sequence of thrust development is interpreted for the imbricate thrusts; major exceptions being the Ritec and Johnston thrusts. These faults have too steep an angle of ramp for them to have formed in sequence with the shallow imbricate faults (Fig. 4). They must form from a deeper level decoupling zone beneath south Pembrokeshire, uplifting and folding the earlier high level decoupling zone at the base of the imbricated Upper Carboniferous. This high level decoupling zone has

since been eroded in south Pembrokeshire. The folding presumably developed by some sticking process on the lower decoupling zone (Fig. 4c) producing layer parallel shortening, buckling and cleavage. Later high-level thrusts grow out of these fold cores; the thrust at St A n n ' s H e a d is a back-thrust in a regional sense, growing out of a south-facing syncline. Further deformation caused fore-thrusts to develop from the basal decoupling zone, the surface expressions of these being the Johnston and Ritec thrusts. The increase in fold wavelength to the south may be due to non-layer-cake stratigraphy above the lower decoupling zone. Silurian, Devonian and Lower Carboniferous rocks all thicken substantially to the south (Hancock et al. 1981) and the sticking on the lower decoupling zone may be due to thickness and facies changes. The Namurian and Westphalian rocks of the upper imbricate belt are fairly constant in thickness and the earlier high-level thrusts were able to extend much further northwards. However thrust climb in these imbricates is lithologically controlled, with ramps developed in thick channel sandstone facies. Exact correlations between the west coast and east coast sections are difficult. The thrusts may involve substantial lateral climb and differential displacement, for example, the Johnston and Ritec thrusts are arranged in an en-echelon manner and must join in a lateral ramp, or more likely, die out in lateral tips as shown in Fig. 2. Similarly the folds are not continuous across south Pembrokeshire.

Variscan tectonics o f S W Britain

A

A"

B

b

c

STICKING PO/NT

Fx6. 4. (a) and (b) illustrate the difference between a piggyback thrust sequence involving a deep level thrust (a) and that of a deep level thrust forming out of sequence (b). In (a) A'B restores to AB. In (b) it is impossible to r e s t o r e A'B to AB if the deep level thrust formed early. (c) The suggested thrust sequence for Pembrokeshire involving a high level imbricate zone folded above a low level decoupling zone. The lower thrust sticks; shortening leads to high strains, buckling and eventual failure, causing the new thrust to cut through the higher level zone, as in (b).

The Upper Carboniferous rocks of south Pembrokeshire show relatively little cleavage development, though south of the Johnston and Ritec thrusts the sandstones locally show well developed pressure solution stripes and often a grain shape fabric, while the shales show a slight slaty cleavage. The south Pembrokeshire fold belt is cut by conjugate wrench faults, suggesting a maximum compressive stress normal to the fold axial planes and extension along the strike of the belt (Anderson 1951). These faults are largely confined to the hanging walls of the Johnston and Ritec thrusts and suggest extension along the strike of the belt during thrusting. In SW Pembrokeshire, at Marloes, locally developed, conjugate brittle-ductile shear zones and associated cleavage indicate a locally compressive stress along the strike of the belt (Knipe & White 1979). Any model for thrust development in Pembrokeshire must allow for this complex strain history and probable nonplane strain finite deformation.

93

Variscan tectonics of Devon and Cornwall SW England has been divided into several tectonic zones based on the attitude of small-scale structures (Dearman 1969; Sanderson & Dearman 1973; Hobson & Sanderson 1983) and the local polyphase deformation correlated across these zones. Recently Shackleton et al. (1982), Rattey & Sanderson (1982) and Coward & McClay (1983) have modified these concepts and described the structures in terms of lowangle shear zones and thrust tectonics. Thus Coward & McClay (1983) considered the schistosity and folds in south Devon to be related to the development of large thrust sheets and locally as in Torbay and south of Dartmouth (Fig. 5), several phases of cleavage and folds were produced during the progressive development of the thrust stack. The thrusts developed in piggyback fashion and early thrusts were folded and carried forwards on lower thrusts (Fig. 6b). Thrusting east of Dartmoor (Waters 1970) is the northward continuation of this zone and a similar section of thrusts and related folds is described from the Plymouth area (Chapman, Fry & Heavey, this volume). In SW Cornwall, the positions of individual thrusts are less certain, as shown in Fig. 5. These thrusts are considered to lie beneath and on the foreland side of a thrust sheet carrying igneous and higher grade metamorphic rocks. In south Cornwall the Lizard complex represents some form of ophiolite, of Lower Devonian age, thrust northwards over the Devonian sediments (Sanders 1955; Styles & Kirby 1980) and the structures are characterized by gently dipping recumbent folds and schistosity (see Fig. 6a, c and Rattey & Sanderson 1982; Hobson & Sanderson 1983). These gently dipping structures necessitate the presence of a low angle shear zone o r decoupling zone at depth and the structures in Fig. 6(c) are shown branching from this. The exact depth of this decoupling zone is uncertain. Note if any steeper ramps had formed at depth or there had been thickening of the lower crust, this would cause changes in the dip of the schistosity and probably have uplifted higher grade rocks. In north Devon the structures are characterized by upright chevron folds and a weak cleavage and form a large synclionrium of Upper Carboniferous Culm facies rocks (Fig. 6a, c). These upright folds are flanked to the north and south by inclined to recumbently folded Lower Carboniferous to Devonian rocks, the structural facing directions being outwards from the synclinorium (Dearman 1969; Sanderson & Dear-

M. P. C o w a r d .& S. S m a l l w o o d

94

e

--~ normal fault ..... tear

1----.~-D

/ c ~-.y

..i. minerallineation

,1
wi#~;~o~th

Tintagel~ ~ Polzea~.jk"~ ~one

1N

Bodmin

/

'~-JJ

ofnaope~arlm~176 .~ ; ~i i

',', ~2 ~

C~'ro,bay

FIG. 5. Map of SW England showing the main thrust traces and major normal faults. For the structure of the complex zone of nappes between the Bodmin and Dartmoor granites, see Isaac el al. (1982). R = Rusey fault zone, C/D Cornwall/Devon boundary. After Hobson & Sanderson (1983), Coward & McClay (1982), Chapman et al. (this volume), and Leveridge & Holder (pers. comm. 1982). PEMBROKE FOLDS & THRUSTS _~

CULM SYNCLINORUM HIGH STRAIN ZONE NORMA L FAULTS FACING CONFRONTATION

~

-<-,~C" " . . . - ' ~ ' x . ~ :

. . :: . ~ . : . . ~

. . . . . . .

"-

""

~

-.

~--_~

LIZARD ~ -

+

+-T~

~ ~,..-~..~..~-C-~--~

| ;::

CARBONIFEROUS

- "

GRANITE

I.( .+; -. 7.+ +'" . ; .' +-+ . ;+ "-+ 7 + ;+ ; ' ;+ +-,

o

FOLD FACING DIRECTION

9t 0 k m

. . . . . .

=~SE o~ CRUSr~_

NORTH

SOUTH

S~VONNAPPE ~LE

(~)

MENDIPS TO S DEVON ---~

(~)

PEMBROKE 70 S CORNWALL HIGH LEVEL IMBRICATES

~)

S IRELAND

BLIND TDRUSi S - -

L IZARD

~

-

-

-

-

~

~

--a

=--

....

--4

:=:--~'~--. . . . .

~-

~

--

N

MAJOR BACK THRUST

?,~_~lOWER Z)ECOUPLING ZONE DISPLACEMENT INCREASES TO WEST

,

1Okra

- . . . . .

FIG. 6. (a) Synoptic cross-section through the Variscan of SW Britain, with data projected on to line PO of Fig. 1. Data for SW England partly from Hobson & Sanderson (1983), Shackleton et al. (1982), data from south Pembrokeshire from Hancock et al. (1981, 1983), Brooks, Mechie & Llewellyn (1983). Crustal thickness from Holder & Bott (1971). (b) and (c) Suggested crosssections through thrust structures of lines PQ and MN (of Fig. 1), data from above references and Coward & McClay (1983). (d) Schematic section through southern Ireland to show the correlation of structures with those of Wales.

o

Variscan tectonics o f S W Britain

man 1973). In the Ilfracombe area, on the north side of the synclinorium, the inclined folds are associated with a well-developed pressure solution cleavage and X : Z strains of over 10:1 (Shackleton et al. 1982). South of Widemouth, on the south side of the synclinorium, the folds become more inclined and then recumbent with increase in strain, to become asymptotic to a major decoupling zone, the Tintagel high strain zone of Sanderson (1979) and Rattey & Sanderson (1982). This zone is about 25 km wide and contains deformed agglomerates and fossils showing X : Z strain ratios generally over 5:1, with narrow zones where the ratios are up to 25:1 (Hobson 1971). Shackleton et al. (1982) estimate a displacement of over 20 km for this zone. The Culm synclinorium presumably developed above the same decoupling zone which forms the base to the shears and thrusts in south Devon and south Cornwall (Fig. 6c). Much of the movement on the south Cornwall structures must have been taken up by this backthrusting between Tintagel and Widemouth (Figs 6c, 7). According to Roberts & Sanderson (1971) and Hobson & Sanderson (1975, 1983) the southern limit of these south facing structures occurs at Polzeath (Fig. 5), south of which the structures are only north facing, although Shackleton et al. (1982) have questioned the exact position of this confrontation. Inland, the Bodmin Moor granite obscures the confrontation zone, but to the east of this, Isaac, Turner & Stewart (1982) describe a complex stack of thrust sheets. They report however, only northward verging thrusts and low-angle exten-

A /

9

~

S TICK ING

~

1

.~---._._

BRANCH LINE

B POINT

~ ....~..

b

"~-'~"r"~ a

FIG. 7. Sketches to show the origin of a 'pop-up' structure. In (a) folding occurs behind the sticking point and failure occurs either by means of a forethrust or a backthrust. Displacement on the backthrust (ab), equals the displacement on the original floor thrust and also equals the displacement of the backthrust branch line. The hinterland drives in as a wedge beneath the backthrust.

95

sional detachment sheets. Isaac et al. (1982) consider the Rusey fault (Freshney 1965), bounding the southern margin of the Culm synclinorium, to be the northern limit of underthrusting. Similar low angle normal faults occur on the coast near Tintagel (Fig. 6a and Freshney 1965; Sanderson 1979). Their significance is discussed below. There is evidence for folding above a deep backthrust in south Devon, where, south of Darmouth, the folds are upright to overturned, facing south. (Shackleton et al. 1982). The upright cleavage is believed to be the same age as the recumbent cleavage to the north, but steepened by a large back f01d as shown in Fig. 6(b). A new cleavage is associated with this back-fold and this increases in intensity southwards and locally obliterates the earlier structures.

The thrust tectonics in general The thrust transport direction The thrust transport direction may be determined from three lines of evidence: (a) the dominant principal extension directions in more deformed rocks; (b) the general trend of folds and thrusts; (c) the trend of tear fauits. The maximum extension directions (X) determined from strain measurements and some mineral elongation directions in south Devon and Cornwall are shown in Fig. 5. Throughout south Devon and south Cornwall, these plunge to approximately 150 ~ suggesting a NNW directed tectonic transport. Similar NNW or SSE plunging lineations occur in the region of back-folding near Tintagel, though in the area S W of Polzeath the mineral lineations have a more anomalous E - W trend. Other anomalous E - W lineations occur south of Torcross in south Devon. Slickensides and fibre growths on faults show a dominant 150 ~ trend in south Devon, though in north Devon and south Pembrokeshire, the slickensides are often more N-S, perpendicular to fold hinges. In south Devon the folds and thrusts trace approximately NE to E N E on the map (Fig. 5), normal to the proposed transport direction, though locally in more deformed rocks there are oblique fold hinges and some bedding/cleavage intersection lineations trend N W - S E almost parallel to thrust movement (Coward & McClay 1983). In Cornwall, north of the Lizard and in the Tintagel area, the majority of fold hinges originally trended normal to the thrust transport

M. P. C o w a r d & S. S m a l l w o o d

96

direction (Hobson & Sanderson 1983) but in south Cornwall, the fold hinges show complex patterns and Rattey & Sanderson (1982) map out alternating zones of: (a) folds originally trending normal to the transport direction, now occurring as lobate or sheath-like folds (see Cobbold & Quinquis 1980) where the hinges have been rotated into the shear direction and (b) zones of dominantly oblique folds where the hinges are not sheath-like but the folds all face to the S W. Rattey & Sanderson (1982) suggest that these oblique folds developed in zones of localized differential sinistral movement where the NE part of the overlying Lizard thrust sheet moved further than the SW part, causing differential movement in underlying thrust sheets. On a regional scale, the backthrusts of the Tintagel and Torcross areas are oblique to the general transport direction. The E - W grain of the structures in Devon and Cornwall is due to these oblique folds and thrusts (Coward & McClay 1983). This obliquity suggests a regional component of differential displacement of dextral sense to the thrust zone as a whole (see Fig. 8). The folds and thrusts in south Pembrokeshire and the Gower also trend E - W , oblique to the thrust transport direction in SW England and also slightly oblique to the boundary of Variscan deformation: the Variscan front. The movement direction for these structures, as suggested from fibres and slickensides, elongation directions in shear zones and the conjugate fault system, is to the north, approximately normal to the fold hinges. This suggests a rotation in the movement direction from the regional NNW trend. This change in movement direction may occur even in one sheet as will be discussed below. It is useful to compare the structures of south Pembrokeshire with those of southern Ireland where Devonian and Carboniferous rocks are similarly deformed into upright folds and are locally cut by thrusts (Cooper et al., this volume). SHEAR COUPLE

BRANCH LINE OBLIQUE TO LATERAL TIP

OBLIQUE FOLD TRACES

FIG. 8. Sketch plan showing variable displacement on a thrust, which has grown from a branch line and the trend of the oblique folds near the oblique to lateral tip.

However the Irish structures have a more E N E trend than those in south Pembrokeshire and trace out a broad arc (Fig. 1). As the structures in south Pembrokeshire are slightly oblique to the margin of the Variscan front, it is suggested that they represent the lateral tip regions of the thrusts, where displacements die out and the faults stick. The arcuate shape of the structures in Ireland may therefore reflect differential displacement on low level thrusts in Devonian strata or below (Fig. lb). The arcuate thrusts may represent the bow of Elliott's bow and arrow rule for differential thrust displacement (Elliott 1976). The structures of SW England are intersected by a series of NNW trending upright tear faults (Dearman 1963). Some of the movement occurred during Mesozoic to Tertiary times; the Tertiary Bovey sediments occupy a NNW trending fault valley east of Dartmoor. However it seems to be too much of a coincidence that these structures follow exactly the dominant trend of Variscan mineral lineations and extension directions and it is likely that they are reactivated tear faults developed during Variscan thrusting. Strains related to thrust development Strain ratios in south Devon and Cornwall are moderate in intensity with X : Z ratios generally between 2.5:1 and 6.5:1 (Coward & McClay 1983). This is in agreement with the deformation intensity suggested by the strong cleavage development in the slates. Coward & McClay (1983) record X : Z strains of over 23:1 for volcanics south of Dartmouth and there are mylonitic textures and high strains in some of the limestones near Torquay. Obviously deformation is heterogeneous, related to the proximity of the rocks to major thrusts and shears and the position of the rocks in major folds. Similar strain variations have been recorded from the Tintagel high strain zone (Rattey & Sanderson 1982). The majority of strains seem to plot on the oblate side of the K = 1 line (see Coward & McClay 1983 for data). This may be due to volume loss associated with plane-strain deformation during thrusting (Ramsay & Wood 1973), but as shown by Coward & Kim (1981) and Coward & Potts (1983), oblate strains may be produced by the superimposition of components of differential movement during thrusting. There may be real extension along the thrust belt. Evidence for this occurs in south Devon, south of Torcross (Fig. 5), where pyrites and fossil orthocones show fibre-fills in an E N E direction. This change in orientation of the strain axes is similar to that on the NW coast of Cornwall, SW of Polzeath (Fig. 5). In both cases the

Variscan tectonics o f S W Britain ENE extension is small. In the rest of south Cornwall and south Devon, even where the strain is low, no evidence has been recorded of such extensions parallel to the trend of the belt. However the conjugate fault system in south Pembrokeshire does suggest E - W extension. Many models of thrust and shear zones assume a bulk plane strain geometry, with no change in length along the shear plane, normal to the shear direction (Coward 1976; Ramsay 1980). However as shown in Fig. 9(a), such extensions may occur in thrust zones above culiminations formed by the local stacking of thrust sheets (Butler 1982). Likewise, where a fault dies out at a lateral tip, then this tip acts as a pole of rotation for the overthrust sheet. Normally this pole will migrate as the thrust grows. However if the thrust displacement rate is large relative to the migration rate of the lateral tip, then the thrust sheet will either develop oblique trending folds in the zone of differential movement (Fig. 9b) or will suffer extension parallel to the thrust trace, normal to the movement direction (Fig. 9c). Thus the extensional strains in south Pembrokeshire may suggest a rotation of the Pembrokeshire thrust sheets about a lateral tip east of the

97

Gower, where the thrusts die out against the Variscan front. Hancock et al. (1983) suggest that about 10% of axial extension was accompanied by about 5 ~ of clockwise rotation of the thrust sheets. Evidence for a larger rotational component is given by the palaeomagnetic work of McCelland-Brown (1983), which indicates a 40 ~ clockwise swing of Carboniferous palaeomagnetic poles in rocks above the Johnston and Ritec thrusts.

Chronology of Variscan deformation K-Ar mineral ages by Dodson & Rex (1971) suggest a diachroneity of deformation across SW England. Ages of 365-345 Ma have been obtained from the slates of south Cornwall, indicating that cooling of micas occurred before the end of the Devonian. There is evidence for some uplift of south Devon and Cornwall in Devonian times; many of the Lower Devonian sediments of south Cornwall (the Gramscatho Beds and Mylor Slates) are poorly sorted greywackes with localized olisotromes, which presumably

', E ~ E N S / ON

THRUST

$••

.OBL IQUE FOLDS

I ",, \~, , ~

f

i

I

I

,~

7 ~

EXTENSION

EX TENSION ,~

STICKING

POINT

POLE 0,~/-

ROTATIONAT

LATERAL TIP

FIG. 9. Three methods of producing extension, normal or oblique to the transport direction. (a) Section through a culmination; the transport direction is here normal to the section. The rocks above the culmination either thin or need extra material from behind this section. (b) Plan of a thrust tip zone, showing oblique extension due to shear produced by differential movement. (c) Plan of a thrust tip showing extension along the thrust belt due to rotation round the lateral tip.

98

M. P. C o w a r d & S. S m a l l w o o d

developed from an uplifted mass to the south (Hobson & Sanderson 1983). Similarly Holwill (1966) records pebbles of deformed Givetian limestone in the Upper Devonian slates of Torbay. The majority of Devonian-Carboniferous slates of south Devon and Cornwall, however, give ages of 340-320 Ma. From stratigraphic evidence, Isaac et al. (1982) consider many of the thrust sheets west of Dartmoor to have formed during the Middle Carboniferous and the Namurian-Westphalian sediments of the Culm synclinorium to be flysch deposits from these thrusts. Late Carboniferous metamorphic ages of 2 9 0 - 2 7 0 M a have been obtained from sediments of the Culm synclinorium and late ages of 310-295 Ma have been obtained from the back-folded areas of south Devon (Dodson & Rex 1971). Thrusting in south Pembrokeshire and the Mendips must be later than the Westphalian sediments. These ages support the structural evidence for the piggyback motion of the thrusts, in that the higher level thrusts in the south formed first and that deformation gradually spread to the NNW. The slates of Cornwall were deformed in Devonian times to develop a widespread penetrative cleavage and then these rocks were carried to the NNW on lower level shears. The mineral ages of the slates were not reset except where the rocks were crenulated by the backthrusting. Shackleton et al. (1982) make a rough estimate of shortening across Devon and Cornwall of over 150 km. This, together with an estimate of the shortening across the Bristol Channel region and the shortening in south Pembrokeshire, suggests a timeaveraged displacement rate in the order of magnitude of 200 km in about 90 My, that is less than 0.25 c m y r -1, and a forward fault propagation rate of over 350 km in 90 My, that is about 0.4 cm yr-1. These rates are slow compared to other orogenic belts. There are few significant structural unconformities in SW England (see Shackleton et al. 1982), suggesting that deformation was submarine throughout the Variscan and that crustal thickening was insufficient to produce major mountain belts, though the Devonian olistromes and Devonian to Carboniferous flysch deposits suggest unstable slopes and foredeep development during deformation. This foredeep migrated northwards from what is now south Cornwall in Devonian times to north Devon in the Middle to Upper Carboniferous. Many of the rapid facies changes in the Devonian, such as the development of Upper

Devonian swells and basins (House et al. 1977) may be due to the migration of a foredeep and peripheral bulge.

Late Hercynian extensional faults Normal faults, trending E - W and dipping to the north at moderate angles occur between Tintagel and Widemouth (Shackleton et al. 1982; Sanderson 1979) locally reorientating the earlier folds in the tilted fault blocks (Fig. 6a). To the south many of the faults are listric and are well exposed in the cliffs from south of Polzeath to Tintagel. In the Tintagel area they curve to join low-angle shears which reorientate and redeform the earlier fabric. In the Tintagel high strain zone it is sometimes difficult to separate the strains related to this extensional phase from strains related to earlier backthrusting; some of the boudinage and the local development of extensional shear bands may be related to this late extensional phase. Isaac et al. (1982) report similar low-angle extensional faults at the southern boundary of the Culm synclinorium between Bodmin Moor and Dartmoor. Their schematic sections suggest over 30% local extension during the development of these structures. Dearman & Butcher (1959) record similar normal faults cutting southward-facing folds on the northern margin of Dartmoor and Moore (1975) suggests that the mineralization surrounding the granites was associated with a N - S extensional regime. In SW Cornwall, Rattey (1980) and Rattey & Sanderson (1982) describe minor extensional structures related to thrust sheet emplacement; the Lizard ophiolite possibly moved forwards by some process of gravity spreading, squeezing out material beneath and in front of it. However it is unlikely that gravity spreading produced the late normal faults at the southern margin of the Culm synclinorium and the Tintagel area; they are more likely related to a late phase of Variscan crustal thinning, possibly related to granite intrusion (Freshney 1965; Isaac et al. 1982) but probably analogous to the extensional faulting of the Basin and Range Province of the western USA (Wernicke & Burchfiel 1982). Some extension may have accompanied the early stages of opening of the English Channel; many of the fault-bounded Mesozoic basins in southern England began in Permian times (Stonley 1982).

Variscan tectonics o f S W Britain

Relation to Variscan tectonics of mainland Europe

A discussion of the European Variscides and the plate tectonic setting is beyond the scope of this paper, but it is worth commenting on the relationship between Variscan thrust tectonics in Britain and mainland Europe. The major tectonic units and the major shear zones and thrust zones of western Europe are shown in Fig. 10. The thrust zone of SW Britain represents a thin-skinned detachment, similar to that described from south Belgium and the Rhenish massif of west Germany (Meissner, Barelsen & Murawski 1981; Weber 1981; Giese 1983). The Rhenish massif is only a small part of a major NW verging thrust complex; all the structures in the internal German Variscides as well as those of the external zone, show NW verging thrusts and shears. Thus in the Moldanubian crystalline zone, the structures are dominated by a low angle schistosity and gneissic banding with thrust sheets of eclogite and granulite facies material (Behr, Engel & Franke 1980). The tectonic transport direction is to the NW as determined from mineral lineations and the axes of major and minor sheath folds (authors' unpublished observations). Throughout the external (Saxo-Thuringian and RhenoHercynian) zones the thrust movement direction was to the NW (Behr et al. 1980; Weber 1981). There are strike-slip faults and mylonite zones in Bavaria, but those developed late, after the main thrust de-

~

~

SHEAR ZONES

"k%~,

. ~

.

'.

.

~

!

,

-.~

~,;x =..J"

~ f CAOOMIAN ~ .

'

":.," MC

Jlill

.... / 2 tam /

" ~pF~'L?

]i

,-

OUTCROP L I M I T

;

~

MOVEMENT

I1!,';

ll~

VENOEMOLO~US~AN -."

ZONE .

DIRECflON

SUIUREZONE

FIG. 10. Map of the Variscan of Europe (after Autran et al. 1980; Weber 1981; Zwart & Dornsiepen 1978). B = Brussels. L = London, P = Paris, BM = Bavarian massif, MC = Massif Central, RM = Rhenish massif. A north-dipping suture zone is indicated by high P/low T blueschists associated with eelogites and granulites in Brittany and the Massif Central. The Icartian (--2000Ma) rocks and Cadomian (-600 Ma) zones form a microplate unaffected by Variscan deformation.

99

formation. They trend NW, slightly oblique to the transport direction in the Rhenish massif. They may be coeval with structures in the Rhenish massif as they do not cut through these external zones. Thus in central Europe Variscan tectonics are dominated by NW verging thrusts; strike-slip shears are subordinate and may be analogous to the large strike-slip shear zones which post-date continental collision in the Himalayas (Molnar & Tapponier 1978). In NW France, however, there is less evidence for crustal shortening during the Variscan; the major thrusts occur in south Brittany and the Massif Central. This area of western France contains older gneissic basement material (Autran et al. 1980) and may well have had a different crustal structure before the Variscan. Presumably, therefore, there must have been some zone of strike-slip movement, equivalent to a transform fault, along the Variscan margin SE of London, to maintain compatibility between the German zone of widespread Variscan deformation and the French zone where Variscan affects are less (Fig. 10). The NW to NNW displacement is oblique to the Variscan front along most of its length West of Brussels (Fig. 10). This oblique closure may be responsible for many of the anomalous strains and fold orientations in SW Britain. The Variscan front follows the trend of the pre-Variscan basin margin; it marks pre-Variscan sedimentary fades and thickness changes. As the thrusts developed and moved oblique to the margin, they would need to change their geometry to the ENE, normal to the thrust transport direction, as well to the NNW. Thus they could easily be pinned at their lateral tips leading to rotation around the sticking point, extension along the thrust trace, the development of oblique folds and thrusts and the development of lateral ramps (Fig. 11). The structures in SW Britain are presumably those to be expected of a zone of transpression (Harland 1971). The Variscan orogeny involved early crustal thinning in SW Britain; Dewey (1982) estimates a stretching factor (/Y) of about 2 for north Devon. This thinning may have led to the development of a back-arc basin along the present line of the English Channel (Leeder 1982). Some obduction of this back-arc magmatic material may have occurred in Devonian times, forming the Lizard complex in SW England. Throughout Devonian and Carboniferous times there was thickening of the previously extended crust producing a sequence of foredeep basins. Deformation was

100

M. P. Coward

brittle stretchingon conjugatefaults ~ ~ t a t i o n a l displacement -~ ~ . orogenicfront~thrust front X ~ . \ ~ . -, _~ ~__~.._. followsor~gino/ ductile extension ~

~

l ~ o b l i q u e

y . ~

~

~

_.

~ension

._.. b.~osinmargin

folds

&thrusts

by ductileshears

FIG. 11. Schematic diagraam to show structures associated with an oblique closure as developed in SW Britain. The main conclusion is that major thrusts are pinned at one lateral tip, sometimes suffer rotational displacement and hence develop extensional strains along the thrust trace.

& S. S m a l l w o o d

largely c o n f i n e d to the u p p e r crust, pres u m a b l y by flaking of the elastic lid to the l i t h o s p h e r e ( D e w e y 1982). I n the o r d e r of 150 k m s h o r t e n i n g has b e e n e s t i m a t e d for this u p p e r crustal zone, that is the original British l o w e r crust and l i t h o s p h e r e must e x t e n d for over 150 k m b e n e a t h the English c h a n n e l and the P r e c a m b r i a n b a s e m e n t of n o r t h e r n France. T h e relatively u n a f f e c t e d Icartian a n d C a d o m i a n b a s e m e n t rocks of n o r t h Brittany (Fig. 10) must be a l l o c h t h o n o u s and overlie a c o n t i n u a t i o n of this d e c o u p l i n g zone in the m i d d l e a n d / o r l o w e r crust.

ACKNOWLEDGMENTS: The authors thank colleagues at Leeds for discussion and referees for improvements to the manuscript.

References ANDERSON, E.M. 1951. The Dynamics o f Faulting. Oliver & Boyd, Edinburgh. AUTRAN, A . , BRETON, J.-P., CHANTRAINE, J. & CmRON, J.-C. 1980. Carte tectonique de la France. I:IM. BRGM. BAOHAM, J. P. N. 1982. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 493-504. BARNES, R. P., ANDREWS, J. R. & BADHAM, J. P. N. 1979. Preliminary investigations of south Cornish m61anges. Proc. Ussher Soc. 4, 262-68. BLUR, H., ENGEL, W. & FRANKE, W. 1980. Guide to Excursion Munchberger Gneissmasse and Bayerischer- Wald. Geology Institute and Museum, Gottingen, W. Germany. 100 pp. BROOKS, M., MECHIE, J. & LLEWELLYN,D. J. 1983. Geophysical investigations in the Variscides of Southwest Britain. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 186-97. Hilger, Bristol. BUTLER, R. H. 1982. The terminology of structures in thrust belts. J. struct. Geol. 4, 239-45. COBBOLD, P. R. & QUINQUIS, H. 1980. Development of sheath folds in shear regimes. J. struct. Geol. 2, 119-26. COWARD, M. P. 1976. Strain within ductile shear zones. Tectonophys. 34, 181-97. & KIM, J. H. 1981. Strain within thrust sheets. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 275-92. & MCCLAY, K. R. 1983. Thrust tectonics of South Devon. J. geol. Soc. London, 140, 215-28. - & Pours, G. J. 1983. Complex strain patterns developed at the frontal and lateral tips to shear

zones and thrust zones. J. struct. Geol. 5, 383-99. DEARMAN, W. R. 1963. Wrench faulting in Cornwall and South Devon. Proc. geol. Ass. 74, 265-87. 1969. Tergiversate folds from southwest England. Proc. Ussher Soc. 2, 115-21. -& BUTCHER, N. E. 1959. The geology of the Devonian and Carboniferous rocks of the northwest border of the Dartmoor granite, Devonshire. Proc. geol. Ass. 77, 199-215. DEWEY, J. F. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. London, 139, 371-412. DODSON, M. H. & REX, D. C. 1971. Potassiumargon ages of slates and phyllites from south-west England. Q. Jl geol. Soc. Lond. 126, 465-99. DUNNING, F. W. 1977. Caledonian-Variscan relations in North-West Europe. In: La Chaine Varisque d'Europe Moyenne et Occidentale. Colloques intn. Cent. natn. Rech. scient. 243, 165-80. ELLIOrr, D. 1976. The energy balance and deformation mechanisms of thrust sheets. Phil. Trans. R. Soc. 283, 289-312. & JOHNSON, M. R. W. 1980. Structural evolution in the northern part of the Moine Thrust Belt, NW Scotland. Trans. R. Soc. Edinb. 71, 69-96. FRESHNEY, E. C. 1965. Low-angle faulting in the Boscastle area. Proc. Ussher Soc. 1, 175-80. GIESE, P. 1983. The evolution of the Hercynian crust-some implications to the uplift problem of the Rhenish Massif. In: FUCHS, K., VON GEHLEN, K., MALZER, H., MURAWSK1, H. &

Variscan tectonics of S W Britain SEMMEL, A. (eds) Plateau Uplift, The Rhenish Shield--a case History, 303-14. SpringerVerlag, Berlin. HANCOCK, P. L. 1973. Structural zones in Variscan Pembrokeshire. Proc. Ussher Soc. 2, 509-20. --, DUNNE, W. M. & TRINGHAM, M. E. 1981. Variscan structures in southwest Wales. Geologic Mijnb. 60, 81-8. & 1983. Variscan deformation in southwest Wales. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 47-73. Hilger, Bristol. HARLAND, W. B. 1971. Tectonic transpression in Caledonian Spitzbergen. Geol. Mag. 108, 27-42. HOBSON, D. M. 1971. Deformed agglomerates near Tintagel, North Cornwall. Geol. Mag. 108, 383-91. -& SANDERSON, D. J. 1975. Major folds from the southern margin of the Culm synclinorium. J. geol. Soc. London, 131,337-52. & 1983. Variscan deformation in southwest England. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles. 108-29. Hilger, Bristol. HOLDER, A. P. & BOTT, M. H. P. 1971. Crustal structure in the vicinity of southwest England. Geophys. J.R. astr. Soc. 23, 465-89. HOLWlLL, F. J. W. 1966. Conglomerates, tufts and concretionary beds in the Upper Devonian of Waterside Cove, near Goodrington Sands, Torbay. Proc. Ussher Soc. 1, 238-41. HOUSE, M. R., RICHARDSON, J. B., CHALLONER, W. G., ALLEN, J. R. L., HOLLAND, C. H. & WESTOLL, T. S. 1977. A correlation of the Devonian rocks of the British Isles. Spec. Rep. geol. Soc. London, 8, 110 pp. ISAAC, K. P., TURNER, P. J. & STEWART, I. J. 1982. The evolution of the Hercynides of central SW England. I. geol. Soc. London, 139, 521-31. JENKINS, T. B. H. 1962. The sequence and correlation of the Coal Measures of Pembrokeshire. Q. J! geol. Soc. Lond. 118, 65-101. KELLAWAY, G. A. & HANCOCK, P. L. 1983. Structure of the Bristol District, the Forest of Dean and the Malvern Fault Zone. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 88-107. Hilger, Bristol. KNIPE, R. J. & WHITE, S. H. 1979. Deformation in low grade shear zones in the Old Red Sandstone, SW Wales. J. struct. Geol. 1, 53-66. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles-Caledonide inheritance versus Hercynian plate margin processes. J. geol. Soc. London, 139, 479-91. LEVERIDGE, B. E. 1974. The tectonics of the Roseland coastal section, south Cornwall. Unpublished Ph.D. Thesis, University of Newcastle upon Tyne. MCCLELLAND-BROWN, E. 1983. Palaeomagnetic studies of fold development and propagation in the Pembrokeshire Old Red Sandstone. Tectonophys. 98, 131-49.

101

MEISSNER, R., BARELSEN, H. & MURAWSKI, H. 1981. Thin-skinned tectonics in the northern Rhenish massif, Germany. Nature, 290, 399-401. MOLNAR, P. & TAPPONIER, P. 1978. Active tectonics of Tibet. J. geophys. Res. 83, 5361-75. MOORE, J. MCMAHON 1975. A mechanical interpretation of the vein and dyke systems of the SW England Orefield. Miner. Deposita, 10, 374-88. NAYLOR, D., REILLY, J. A., SEVASTOPULO,G. D. & SLEEMAN, A. G. 1983. Stratigraphy and structure of the Irish Variscides. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt of the British Isles, 20-46. Hilger, Bristol. RAMSAY, J. G. 1980. Shear zone geometry: a review. J. struct. Geol. 2, 83-99. -& WOOD, D. S. 1973. The geometric effects of volume change during deformation processes. Tectonophys. 16, 263-77. RATTEY, P. R. 1980. Deformation in southwest Cornwall. Proc. Ussher Soc. 5, 39-43. & SANDERSON, D. J. 1982. Patterns of folding within nappes and thrust sheets: examples from the Variscan of southwest England. Tectonophys. 88, 247-67. ROBERTS, J. L. & SANDERSON, D. J. 1971. Polyphase development of slaty cleavage and the confrontations of facing directions in the Devonian rocks of North Cornwall. Nature, 230 87-9. SANDERS, L. D. 1955. Structural observations on the southeast Lizard. Geol. Mag. 92, 231-40. SANDERSON, D. J. 1979. The transition from upright to recumbent folding in the Variscan fold belt of southwest England: a model based on the kinematics of simple shear. J. struct. Geol. 1, 171-80. & DEARMAN, W. R. 1973. Structural zones of the Variscan fold belt in SW England: their location and development. J. geol. Soc. London, 129, 527-33. SHACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structure in SW England. J. geol. Soc. London, 139, 533-41. STONLEY, R. 1982. The structural development of the Wessex basin. J. geol. Soc. London, 139, 543-54. STYLES, M. T. & KIRi3u G. A. 1980. New investigations of the Lizard Complex, Cornwall, England and a discussion of an ophiolite model. In: Ophiolites. Proc. int. Ophiolite Symp. Cyprus 1979, 517-26. Cyprus Geol. Surv. Dept. SULLIVAN, R. 1965. The Mid-Dinantian stratigraphy of a portion of central Pembrokeshire. Proc. geol. Ass. 76, 283-300. WALLACE, P. 1983. The subsurface Variscides of southern England and their continuation into Continental Europe. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 198-208. Hilger, Bristol. -

-

-

-

102

M. P. C o w a r d & S. S m a l l w o o d

WATERS, R. A. 1970. The Variscan structures of eastern Dartmoor. Proc. Ussher Soc. 2, 191-7. WEBER, K. 1981. The structural development of the Rheinische Schiefergebirge. Geologie Mijnb. 60, 149-60.

WERNICKE, B. & BURCHFIEL, B. C. 1982. Modes of extensional tectonics. J. struct. Geol. 4, 105-15. ZWART, H. J. • DORNSIEPEN, U. F. 1978. The tectonic framework of central and western Europe. Geologie Mijnb. 57, 627-54.

M. P. COWARD, Department of Earth Sciences, University of Leeds, Leeds LS2 9JT. S. SMALLWOOD, Department of Earth Sciences, Downing Street, Cambridge CB2 3EQ.

Thrust nappe tectonics in the Devonian of south Cornwall and the western English Channel B. E. Leveridge, M. T. Holder & G. A. Day SUMMARY: A review of stratigraphical data from south Cornwall, considered in conjunction with information obtained during a current re-survey by the British Geological Survey of 1:50,000 Sheet 352 (Falmouth), indicates the presence of a major thrust nappe in the Devonian rocks of the district. Seismic records from the western English Channel contain a prominent event that is interpreted as the offshore extension of the thrust. A model of the thrust nappe system of the area is proposed which is consistent with obduction from a small oceanic basin to the south.

The concept that thrusting played a significant role in the geological evolution of south Cornwall, and particularly that the Lizard complex is allochthonous (Flett 1933; Hendriks 1939), has gradually gained support and elaboration (e.g. Green 1964; Strong et al. 1975). Thrusts separating ophiolitic basic/ultrabasic rocks and metamorphic rocks within the complex have been delineated recently (Bromley 1979; Styles & Kirby 1980). Relationships of the lithostratigraphical units north of the Lizard (Fig. 1) are not as clearly defined. There have been various interpretations of the stratigraphy (Hendriks 1937; Sadler 1973; Wilson & Taylor 1976) none of which has prevailed. However, a common view that much of the succession is Lower Devonian (e.g. Floyd 1982; Shackleton et al. 1982) stems largely from a misinterpretation of palaeontological determinations by Sadler (1973). Drawing upon available data and their current work the authors here propose a tectonic model for south Cornwall that is in accord with palaeontological data and relates stratigraphy to major thrust nappe development.

Stratigraphy Gramscatho Group The Gramscatho Group (Fig. 1) is made up of three tormations formerly called the Lower, Middle and Upper Gramscatho Beds by Hendriks (1937). The Lower Gramscatho Formation has the largest outcrop; it consists of sandstone turbidites and dark grey slates interpreted by Hendriks (1937) as a flysch sequence. The Middle Gramscatho Formation, which includes the Veryan Limestone of Sadler (1973), has been recognized only in Roseland. It comprises slates, that are locally manganiferous (Hendriks

1937), interbedded with rhythmic sequences of sandstone and limestone turbidites and radiolarian cherts (Leveridge 1974): an association indicative of deep water deposition. The Upper Gramscatho Formation is formed of slumped and turbiditic sandstones representing a return to the sedimentary regime of the lower Formation. Middle Eifelian conodonts from the turbiditic limestones of the Middle Gramscatho Formation (Sadler 1973, fig. 2A, C) are probably derived but consistent ages suggest that derivation and emplacement of the limestones rapidly followed primary deposition. The age suggests that the very thick Lower Gramscatho Formation is early Middle Devonian and older. The former is consistent with early determinations by Lang (1929) on plant debris from the sandstones.

Roseland and Meneage breccias The belt of rocks here called the Roseland and Meneage breccias (Fig. 1) has been variously interpreted in terms of origin and age and has long been the crux of stratigraphical problems in south Cornwall. It comprises dark grey slates that include dispersed grains and macroscopic clasts of sedimentary, metamorphic, volcanic and magmatic rocks, interbedded with sandstones, framework breccias/conglomerates and volcanic rocks. An established Ordovician age for the quartzites in the belt led Hill & MacAlister (1906) to infer that the whole sequence was Ordovician. Lower Palaeozoic and metamorphic rocks were later recognized as discrete blocks (Hendriks 1931; Flett 1933) and the belt was interpreted as a 'crush zone' adjacent to the Lizard Complex by Flett (1933). Hendriks (1937) proposed that the older rocks occupied thrust 'belts' in an Upper Devonian sequence. Sadler (1973), deducing 103

104

B.E.

Leveridge

from the presence of bedding and sedimentary structures that the rocks formed part of a normal folded sequence, reinterpreted them as a condensed succession of Ordovician and Lower Devonian to Middle Devonian (Roseland Volcanics, Sadler 1973) rocks sliced up by thrust faults. Recent workers, however, have proposed that the Roseland and Meneage breccias are sedimentary (Lambert 1965; Leveridge 1974; Barnes et al. 1979). A primary origin is confirmed by a combination of criteria which include: (i) a sedimentary black slate matrix, (ii) interbedded pyroclastic flows, (iii) clasts lying between sandstone turbidites, (iv) juxtaposition of dissimilar clasts, (v) gradation between polymict framework conglomerates/breccias and the main body of matrix supported breccia, (vi) highly rounded and irregular clast shapes, (vii) continuity of early cleavage from matrix into the less competent clasts. These rocks have been termed wildflysch (Leveridge 1974) and olistostrome (Barnes et al. 1979), being interpreted as the product of turbidity and mass flow, slumping and sliding. The junction with the Gramscatho Group in Roseland (SW 912 379), considered to be a thrust by Sadler (1973), shows these features. Sandstone turbidites of the Upper Gramscatho Formation pass up into typical pebbly mudstone into which large quartzite olistoliths have ploughed. The wildflysch source, identified by olistoliths including schists and shelf facies Ordovician, Silurian and Lower to Middle Devonian sediments (see Hendriks 1937; Sadler 1973) was probably Brioverian basement and Palaeozoic rocks similar to those in Finist6re (Bradshaw et al. 1967). Rocks regarded as contemporaneous with the dark slate matrix of the Roseland and Meneage breccias have yielded conodonts suggesting possible Givetian (Leveridge 1974, fig. 2A, D) and Frasnian (Hendriks et al. 1971, fig. 2A, E) ages.

Mylor Slate Formation The Mylor Slate Formation consists of blue grey slate with local thin interbedded sandstones and siltstones and sporadic beds of pillow lava and sedimentary breccia; the thickest breccias occurring along the south-eastern boundary with the Gramscatho Group. It is considered to be a basinal sequence, the sand-

et al.

stones and siltstones representing distal turbidites (Wilson & Taylor 1976). The Formation crops out at the core of the 'Truro antiform' (Fig. 1) and has been considered to underly the Gramscatho Group (Wilson & Taylor 1976; Rattey & Sanderson 1982). Current work by the authors and B.G.S. colleagues has established that Gramscatho Group rocks on both southern and northern 'limbs' dip and young to the SE (Fig. 2A) and that on the northern flank they underlie the Mylor Slate Formation. There is no evidence that the Mylor Slate Formation passes laterally into the Gramscatho Group. There are two south-coast sections of the junction between the Mylor Slate Formation and the Gramscatho Group, one at Loe Bar and the other at Flushing near Falmouth. At Loe Bar (SW 640 244) the contact is a NE- trending steep normal fault hading and downthrowing Gramscatho Group rocks to the SE. Sedimentary breccias comprising of grey mudstone and sandstone clasts, generally matrix supported but also clast supported, in lenses and continuous beds, form the footwall. Such olistostromes, comprising Gramscatho Group detritus, predominate in the upper 450 m of the Mylor Slate Formation. The junction is exposed at Flushing (SW 816 336) where alternating slates and sandstones of the Gramscatho Group rest on sedimentary breccias composed of slate with sandstone clasts. Such breccias are absent along the mapped boundary between Restronguet Point and Truro. The Mylor Slate Formation has yielded palynomorph assemblages of Upper Devonian age (Turner et al. 1979, fig. 2A, F).

A m a j o r thrust n a p p e Field evidence The pattern of outcrop and stratigraphical evidence in south Cornwall are reconciled if the rocks of the southern 'limb' of the 'Truro antiform' constitute a thrust nappe. The thrust, here called the Carrick thrust, separates the Mylor Slate Formation and the Gramscatho Group (Fig. 2A). The nappe repeats on the south coast a Middle and Upper Devonian succession equivalent to the Gramscatho Group of the north coast and the central Mylor Slate Formation. Differing lithologies of the Mylor Slate Formation and the Roseland and Meneage breccias may be attributable to an initial relative proximity of the latter to the sedimentary source. At Loe Bar, as the faulted junction between the Mylor Slate Formation and Gramscatho

FIG. 1. A geological sketch map of south Cornwall based on B.G.S. 1:50,000 maps.

FIG. 2. (A) A sketch map showing the Carrick nappe and the bounding Carrick thrust in south Cornwall. (B) Metamorphic age zones based on Dodson & Rex (1971).

106

B . E . Leveridge et al.

Group is approached, breccia clasts show a significant increase in flattening in S 1 and early lenticular quartz segregations are elongated in a N N W - S S E direction. At Flushing, sedimentary breccia clasts in the Mylor Slate Formation show strongly increased flattening in a 2 - 3 m zone below the contact with the Gramscatho Group. In Gramscatho lithologies near the contact early folds are sheared with limbs boudinaged and pulled apart. The course of the Carrick thrust from Truro, where north and south Gramscatho Group rocks are brought into juxtaposition, is somewhat speculative due to poor exposure. On geometrical grounds it cannot extend westwards and this is confirmed by a southward younging succession on the north coast between the Meadfoot Beds at Perranporth and the Mylor Slate Formation. The trend of the thrust is thus probably" eastwards in the direction of strike of bedding (and S1); passing into Mevagissey Bay at Pentewan Beach where it appears to correspond with a prominent offshore seismic reflector (Fig. 4B). The northern limit of the Gramscatho Group rocks on I.G.S. 1:50,000 maps has been termed the Perranporth-Pentewan line (Dearman 1971). It also passes into Mevagissey Bay and has been regarded as tectonically significant (Hobson 1976). At its SE end at Pentewan, the Gramscatho Group and Meadfoot Beds are separated by a steep normal fault but at Perranporth Gramscatho Group rocks overlie Meadfoot Beds with a sedimentary contact near Cligga Head (SW 737 537) somewhat to the south of the mapped junction (Fig. 1). A major slide proposed by Sanderson (1971) as a possible expression of the line at Perranporth was refuted by Henley (1973). Work in progress by the authors shows that the main structural feature in the area north of Perranporth is a steep zone of secondary transposition cleavage (S 1 of Sanderson 1971; Henley 1973). The zone lies within the Meadfoot Beds and does not extend across the peninsula to Mevagissey Bay. The shallow dipping seismic reflector in the Bay is thus concluded to be Carrick thrust. Supportive evidence for thrusting is provided by radiometric age determinations and

geochemical studies in south Cornwall. Agezones trending E - W across the peninsula, that are generally younger northwards, were defined by Dodson & Rex (1971) (Fig. 2B) and related to separable tectonic phases, although they could equally well have represented a northward tectonic migration (Dearman et al. 1969). Dodson & Rex (1971) describe an apparently anomalous NE-trending 345-365 Ma zone, encompassing the Lizard, Meneage and Roseland that terminates against a younger E - W zone to the north. Geochemical work by Floyd (1981) has determined the presence of two differing suites of extrusive and intrusive tholeiitic rocks in the district. The north and west of the area (Fig. 1) are characterized by within-plate basalts, whereas adjacent to the Lizard, in a belt running through Roseland, the basic rocks have the chemical signatures of variably enriched mid-ocean ridge basalts. The Carrick thrust nappe (Fig. 2A) could account for the juxtaposition and orientation of the belts of basalt, originating in different tectonic settings, and also the southern metamorphic age zone if the near-source rocks with older metamorphic ages have been transported northwards (see below). Thrusting may also explain the repetition above the Roseland and Meneage breccias of Gramscatho Group lithologies on the Dodman (McKeown 1962), which vary in the direction of dip from slaty to phyllitic rocks, and the 'Variscan' sediments with spilites offshore, up to 2.5 km south of Dodman Point, recorded by Sadler (1974). Offshore evidence In seven seismic reflection profiles obtained offshore south and east of the Lizard, Permo-Triassic rocks with a velocity of 4.25 km s-1 onlap 'basement'. Samples close inshore identify this 'basement' as Devonian/Carboniferous and it is contiguous with the Gramscatho Group and the Roseland and Meneage breccias onshore. On a profile running NNW towards Dodman Point (Fig. 3) a number of events dip in a southerly direction within the 'basement'. The deepest of these events (just above 4 s two-way travel time at the left edge of

FIG. 3. (A) Seismic section trending NNW to approximately 6 km south of Dodman Point (line A-B in Fig. 4B). (B) Interpretation of (A). Individual thrusts are identified by correlation, via adjacent seismic lines, with onshore geology. Numbers at left are seismic two-way travel times in seconds. FIG. 4. (A) Two-way travel-time map of the top of the Devonian basement. Isochrons in seconds (PBF = Plymouth Bay fault). (B) Two-way travel-time map of the persistent prominent reflector in the Devonian basement identified as the Carrick thrust. Isochrons in seconds. The inferred outcrop of the Carrick and Dodman thrusts and the subcrop of the Lizard boundary thrust beneath the Permo-Trias are shown. (PBF = Plymouth Bay fault, A-B = line of seismic section in Fig. 3.)

Thrust nappe tectonics in Cornwall

107

108

B . E . Leveridge et al.

the record) can be correlated with a similar event on adjacent profiles indicating that this is approximately a dip section. A velocity of 5.4 km s -~ is applicable for 'basement' in this area (M. Brooks & J. J. Doody, pers. comm.) and on this basis a true dip of 300-35 ~ is calculated for this event. Extrapolating the event northwards would bring it to the surface where the Carrick thrust is inferred on land. The considerable thickness of 'basement' seen in the seismic section between this reflector and the next parallel reflector above it is thus probably the Carrick nappe. The thickness of the Carrick nappe is about 10.5 km assuming a velocity of 5.4 km s-1. Refraction work however, indicates greater velocities at depth and the thickness may be nearer 12 kin. Further interpretation of seismic records from west of the Lizard has indicated continuity of this reflector into Mounts Bay, where it appears to correlate with the proposed Carrick thrust onshore (Day & Edwards 1983).

Nappe development and deformation history The similarity in age of syn-D 1 metamorphism in SE Cornwall (Dodson & Rex 1971) and the Roseland and Meneage breccias suggests that the latter are syn-tectonic, generated by thrust movements during D 1. Early phase deformation is characterized by slaty cleavage and tight to isoclinal assymmetric folds of variable trend and facing (cf. Turner 1968). In Roseland fold morphology (Dearman 1969), comparable to sheath folds (Cobbold & Quinquis 1980), represents a NNW overriding movement (Leveridge 1974). North of the Lizard folding also records northerly translation, but Rattey & Sanderson (1982) recognize a fold zonation in which belts of 'regional' ENE-trending northward facing folds alternate with N-S zones of oblique folds. They attribute oblique folding to differential advance in the overriding Lizard thrust sheet. However, in the oblique zone between Falmouth and Restronguet Point regional NE-trending secondary structures are similarly oriented N-S (authors' unpublished work), suggesting rotation related to subsequent granite intrusion despite evidence of a lateral ramp to the north (see below). Secondary deformation pre-dating granite related structures (Turner 1968; Rattey 1979) occurs in discrete NE-trending fold belts (Rattey 1980). Some are characterized by gently to moderately inclined northerly verging minor folds plunging between E and SE (cf. Dearman et al. 1980) and others contain steeply inclined,

gently plunging, tight to open folds verging NW (Leveridge 1974). The latter consistently cut the former in Roseland but Rattey & Sanderson (1982) suggest that secondary steep and shallow structures relate along curving shear zones generated by continued northerly movement of the Lizard. Certainly the steeper belts are associated with minor thrusts (e.g. Parsons Beach, SW 788 271) and may have been generated either by shears related to early thrusting or secondary movement over thrust ramps.

A thrust model A model, based on available structural and stratigraphical evidence explaining the geometry and sequence of proposed thrusting in south Cornwall is presented in Figs 5 and 6. Thrusting within the Lizard complex (Bromley 1979) has been omitted for clarity. The onshore information in Fig. 2(A) shows that older stratigraphical units are thrust over younger units in the direction of slip (NNW), but at right angles to slip younger units have been cut out by the thrusts in the manner demonstrated by Elliot & Johnson (1980) in the Moine thrust zone and Harris & Milici (1977) in the Appalachian thrust belt. Offshore a number of SSE-dipping events can be seen in the seismic profiles south of the event identified as the Carrick thrust. Although the Carrick thrust can be mapped from the offshore seismic data (Fig. 4B) the other events do not appear to have the same lateral continuity. In the profile (Fig. 3) the events above the Carrick nappe are interpreted as the Dodman and Lizard boundary thrust planes although this identification is tentative. Permo-Triassic rocks, with low velocities, thicken towards the south (Fig. 4A) and this has the effect of bending the events down to the south on the profile, so that in reality the thrust planes are slightly listric. A number of profiles contain short near-horizontal events at or near the bottom of the record (e.g. the event at 5.2 s two-way travel time in Fig. 3). These may represent reflections from a sole thrust at a depth between 13.6 and 15 km. As no units below the Gramscatho Group are apparently involved in the Carrick nappe, the sole thrust probably lies at or near the base of this Group in south Cornwall. South-west of Falmouth a substantial thickness of the Gramscatho Group in the Carrick nappe has been cut out (Fig. 5) by two frontal footwall ramps (Butler 1982). North and east of Falmouth the greater thickness of the Gramscatho Group

Thrust n a p p e tectonics in C o r n w a l l

109

FIG. 5. Block diagram of the proposed thrust model illustrating the structure of the nappes. (ct = Carrick thrust, dt = Dodman thrust, st = Sole thrust, lbt = Lizard boundary thrust. Arrows indicate slip direction.) Mylor Slate Formation and the Gramscatho Group of the basal unit are cut out to the east by a lateral footwall ramp in the Carrick thrust. The Roseland and Meneage breccias of the Carrick nappe are cut out to the west by a lateral footwall ramp in the Lizard boundary thrust. The D o d m a n nappe is cut out to the west by a lateral hangingwall ramp in the D o d m a n thrust.

FIG. 6. Generalized thrust sequence diagram for south Cornwall. The Mylor Slate Formation is believed to interdigitate with the Roseland and Meneage breccias but its relationship with the Carrick nappe is not known. The Lower Palaeozoic and Brioverian (?) is equivalent to the Icartian microplate of Shackleton et al. (1982). (ct = Carrick thrust, dt = D o d m a n thrust, st = sole thrust, lbt = Lizard boundary thrust, bt = Brioverian thrust (?), MB = Mylor Slate Formation breccias, RMB = Roseland and Meneage breccias, DP = Dodman phyllites.)

110

B.E.

Leveridge

suggests that a simpler single footwall ramp connects the sole thrust with the surface in the Truro area. Linking these footwall ramps between Restronguet Point and Truro a lateral footwall ramp (Butler 1982) is suggested by the cross-cutting of bedding in the Mylor Slate Formation by the Gramscatho Group (Figs 2A & 5). The slump breccias immediately below the Carrick thrust between Restronguet Point and Loe Bar are interpreted as the erosional products of the nappe as it moved across the sediment surface on the Mylor Slate Formation. The Dodman nappe, resting above the Carrick nappe, probably represents a horse created by a footwall ramp failure (Elliot & Johnson 1980) in the Gramscatho Group (Fig. 6). Its western boundary is formed by a lateral hangingwall ramp, converted into a culmination wall (Butler 1982) by overthrusting on to the Carrick nappe, which amalgamates the Dodman thrust and the Lizard boundary thrust (Fig. 5). Overlying the Dodman and Carrick nappes is the Lizard nappe and the rocks of the Lower Palaeozoic and Brioverian (?) thrust sheet. The Lizard complex, and thus the ophiolite (Bromley 1979; Styles & Kirby 1980), is less extensive (Curry et al. 1970) than the probable Lizard nappe and may be a horse on the Lizard boundary thrust.

Sequence of thrusting Thrust movement commenced south of the area, probably in the Lower or Middle Devonian, with erosion of an upthrust mass providing the sediment for the Gramscatho Group. By Upper Devonian times a complex thrust mass, composed of the Lizard nappe and the Lower Palaeozoic and metamorphic basement rocks (Brioverian ?) of a continental foreland may have begun to override the Gramscatho Group flysch from the SE (Fig. 6A). Lower Palaeozoic and metamorphic debris eroded from the front of this thrust mass began to form the Roseland and Meneage breccias on top of the Gramscatho Group. Continued slip on the Lizard boundary thrust would have caused this debris to be overridden, whilst further erosional debris was added to the breccias in front of the nappe (Fig. 6B). Later failure of the frontal footwall ramp beneath the Lizard nappe formed a new glide plane within the Gramscatho Group and detached a thin slice of Gramscatho flysch and its Roseland and Meneage breccia capping. Slip on this glide plane, the Dodman thrust, carried the Dodman nappe and the overlying Lizard nappe in piggyback style (Elliot & Johnson 1980) on to the Roseland and Meneage brec-

et al.

cias and the underlying Gramscatho Group (Fig. 6C). Subsequently the transference of slip from the Dodman thrust to the underlying Carrick thrust transported the entire nappe pile to the NNW, locally over its own erosional debris on top of the Mylor Slate formation (Fig. 6D).

Concluding remarks The seismic profiles of the western Channel compare with those obtained by Meissner et al. (1981) from the Rheno-Hercynian in Germany where shallow dipping reflectors are interpreted as thrusts passing down to a sole thrust extending under the full width of the zone and at least 200 km laterally. Whereas in Germany this sole thrust lies just over 4 km below surface, in the Channel the sub-horizontal seismic reflectors lie at depths greater than 13 kin. Such a difference in depth may be due to: (i) the German profiles lying closer to the deformation front, (ii) greater uplift and erosion of the German nappe pile, (iii) an intra-continental environment in the east of the Variscides (Weber 1981) produced thinner sequences than the deep basin environment in the west. The proposed model of thrusting in south Cornwall, it" combined with the thrust models of Coward & McClay (1983) for south Devon and Isaac et al. (1982) for central Devon, indicates an overall pattern of northerly directed thrusting in the Variscides of SW England. The age of this thrusting, Upper Devonian in south Cornwall and late Vis6an in central Devon (Isaac et al. 1982) indicates northward tectonic migration similar to that in the Variscides of Europe (Shackleton et al. 1982). The Carrick nappe comprises a flysch/wildflysch sequence derived from a continental mass south of the Lizard; the Icartian microlate of Shackleton et al. (1982). The sequence probably accumulated in deep water and includes some large basic bodies that may be ophiolitic (e.g. Nare Head in Roseland). This association has some of the characteristics of a plate collision zone (Mitchell 1974) but Shackleton et al. (1982) compare it with m61anges underlying the obducted ophiolites in Oman and the Himalayas. Emplacement of the Cornish nappes, therefore, probably accompanied closure of either an oceanic back-arc basin (Floyd 1982) or a pullapart basin (Badham 1982). Parts of the basin floor, obducted when the Icartian microplate (Shackleton et al. 1982) was overthrust into the

,Thrust nappe tectonics in Cornwall basin, were incorporated as ophiolitic fragments in the Lizard nappe. ACKNOWLEDGMENTS: We wish to thank our colleagues A. J. J. Goode and R. T. Taylor for information relat-

111

ing to the Gramscatho Group on the northern side of the peninsula and for their contribution t o the paper made in discussion of its contents. This paper is published by permission of the Director, British Geological Survey (N.E.R.C.).

References BADHAM, J. P. N. 1982. Strike-slip o r o g e n s - - a n explanation for the Hercynides. J. geol. Soc. London, 139, 493-504. BARNES, R. P., ANDREWS, J. R. & BADHAM, J. P. N. 1979. Preliminary investigations of south Cornish m61anges. Proc. Ussher Soc. 4, 262-8. BRADSHAW, J. D., RENOUF, J. T. & TAYLOR, R. T. 1967. The development of Brioverian structures and Brioverian/Palaeozoic relationships in west Finist6re (France). Geol. Rdsch. 56, 567-96. BROMLEY, A. V. 1979. Ophiolitic origin of the Lizard Complex. J. Camborne Sch. Mines, 79, 35-8. BUTLER, R. W. H. 1982. The terminology of structures in thrust belts. J. struct. Geol. 4, 239-45. COBBOLD, P. R. & QUINQUIS, H. 1980. Development of sheath folds in shear regimes. J. struct. Geol. 2, 119-26. COWARD, M. P. & MCCLAY, K. R. 1983. Thrust tectonics in Devon. J. geol. Soc. London, 140, 215-28. CURRY, D., HAMILTON, D. & SMITH, A. J. 1970. Geological and shallow subsurface geophysical investigations in the Western Approaches to the English Channel. Inst. geol. Sci Rep. No. 70/3, 12 PP. DAY, G. A. & EDWARDS, J. W. 1984. Variscan thrusting in the basement of the English Channel and SW Approaches. Proc. Ussher Soc. (in press). DEARMAN, W. R. 1969. Tergiversate folds from South-West England. Proc. Ussher Soc. 2, 112-5. 1971. A general view of the structure of Cornubia. Proc. Ussher Soc. 2, 220-36. ~, LEVERIDGE, B. E. & TURNER, R. G. 1969. Structural sequences and the ages of slates and phyllites from south-west England. Proc. geol. Soc. no. 1654, 35-8. , - - , RATTEY, R. P. & SANDERSON, D. J. 1980. Superposed folding at Rosemullion Head, South Cornwall. Proc. Ussher Soc. 5, 33-8. DODSON, M. H. & REX, D. C. 1971 (for 1970). Potassium-argon ages of slates and phyllites from south-west England. Q. Jl geol. Soc. Lond. 126, 465-99. ELLIOTT, D. & JOHNSON, M. R. N. 1980. Structural evolution in the northern part of the Moine thrust belt, NW Scotland. Trans. R. Soc. Edinb. 71, 69-96. FLETT, J. S. 1933. The geology of the Meneage. Mere. geol. Surv. Gt. Br. Suture. Prog. for 1932, pt 2, 1-14. FLOYD, P. A. 1981. Geochemical comparison of basaltic rocks in the Lizard Ophiolite and the Cornubian troughs. Conf the Lizard Complex (abstracts). Geological Society of London. -

-

-

1982. Chemical variation in Hercynian basalts relative to plate tectonics. J. geol. Soc. London, 139, 505-20. GREEN, D. H. 1964. A re-study and re-interpretation of the geology of the Lizard Peninsula, Cornwall. In: Present Views o f some Aspects o f the Geology o f Cornwall and Devon. R. geol. Soc. Cornwall Commem. Vol. 1964, 87-114. HARRIS, L. D. & MILICI, R. C. 1977. Characteristics of thin-skinned style of deformation in the Southern Appalachians, and potential hydrocarbon traps. Prof Pap. U.S. geol. Surv. 1018, 40 pp. HENDR~KS, E. M. L. 1931. The stratigraphy of South Cornwall. Rep. Brit. Ass. 1930, 332. 1937. Rock succession and structure in South Cornwall, a revision. With notes on the Central European facies and Variscan folding there present. Q. Jl geol. Soc. Lond. 93, 322-60. 1939. The Start-Dodman-Lizard Boundary Zone in relation to the Alpine structure of Cornwall. Geol. Mag. 76, 385-402. ~, HOUSE, M. R. & RHODES, F. H. T. 1971. Evidence bearing on the stratigraphical successions in South Cornwall. Proc. Ussher Soc. 2, 270-5. HENLEY, S. 1973. The structure of the Perranporth area, Cornwall. Proc. Ussher Soc. 2, 521-4. HOBSON, D. M. 1976. The structure of the Dartmouth Antiform. Proc. Ussher Soc. 3, 320-32. HILL, J. B. & MACALISTER, D. A. 1906. The geology of Falmouth and Truro. Mere. geol. Surv. G. B. ISAAC, K. P., TURNER, P. J. & STEWART, I. J. 1982. The evolution of the Hercynides of central SW England. J. geol. Soc. London, 139, 521-31. LAMBERT, J. L. M. 1965. A re-interpretation of the breccias in the Meneage crush zone of the Lizard boundary, south-west England. Q. Jl geol. Soc. Lond. 121, 339-57. LAMa, W. H. 1929. On fossil wood (Dadoxylon Hendriksi, n.sp.) and other plant remains from the clay-slates of S. Cornwall.Ann. Bot. 43, 663-83. LEVERIDGE, B. E. 1974. The tectonics of the Roseland coastal section, south Cornwall. Unpublished Ph.D. Thesis. University of Newcastle upon Tyne. MCKEOWN, M. C. 1962. Structural studies at Dodman Point, Cornwall. Proc. Ussher Soc. 1, 22-3. MEISSNER, R., BARTELSEN, H. & MURAWSKI, H. 1981. Thin-skinned tectonics in the northern Rhenish Massif, Germany. Nature, 290, 399-401. MITCHELL, A. G. H. 1974. South-west England granites; magmatism and tin mineralisation in a post-collision tectonic setting. Trans. Inst. Min. metall. B, 83, B95-7. RATTEY, P. R. 1979. The relationship between -

-

-

112

B.E. Leveridge et al.

deformation and intrusion in the Cornubian batholith in south-west Cornwall. J. Camborne Sch. Mines, 79, 60-3. 1980. Deformation in south-west Cornwall. Proc. Ussher Soc. 5', 39-43. & SANDERSON, D. J. 1982. Patterns of folding within nappe and thrust sheets: examples from the Variscan of south-west England. Tectonophys. 88, 247-67. SADLER, P. M. 1973. An interpretation of new stratigraphic evidence from south Cornwall. Proc. Ussher Soc. 3, 535-50. 1974. An appraisal of the 'Lizard-Dodman-Start Thrust' concept. Proc. Ussher Soc. 3, 71-81. SANDERSON, D. J. 1971. Superposed folding at the northern margin of the Gramscatho and Mylor Beds, Perranporth, Cornwall. Proc. Ussher Soc. 2, 266-9. SHACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in SW England. J. geol. Soc. London, 139, 533-41.

-

-

STRONG, D. F., STEVENS, R. K., MALPAS, J. & BADHAM, J. P. N. 1975. A new tale for the Lizard (abstract). Proc. Ussher Soc. 3, 252. STYLES, M. T. & KIRBY, G. A. 1980. New investigations of the Lizard Complex, Cornwall, England and a discussion of an ophiolite model. Extract. Proc. Int. Ophiolite Syrup. Cyprus geol. Surv. Dept 517-26. TURNER, R. E., TAYLOR, R. T., GOODE, A. J. J. & OWENS, B. 1979. Palynological evidence for the age of the Mylor Slates, Mount Wellington, Cornwall. Proc. Ussher Soc. 4, 274-83. TURNER, R. G. 1968. The influence of granite emplacement on structures in south-west England. Unpublished Ph.D. Thesis. University of Newcastle upon Tyne. WEBER, K. 1981. The structural development of the Rhienisches Schiefergebirge. In: ZWART, H. J. & DORNSIEPEN, U. F. (eds) Variscan Orogen in Europe. Geologic Minjb. 60, 149-59. WILSON, A. C. & TAYLOR, R. Y. 1976. Stratigraphy and sedimentation in west Cornwall. Trans. R. geol. Soc. Cornwall, 20, 246-59.

B. E. LEVERIDGE & M. T. HOLDER, British Geological Survey, St Just, 30 Pennsylvania Road, Exeter EX4 6BX, England. G. A. DAY, British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 3LA, Scotland.

A structural cross-section through S W Devon T. J. Chapman, R. L. Fry & P. T. Heavey SUMMARY: A new interpretation for the Variscan structure of SW Devon is presented in the form of a structural cross-section. It is argued that the lowest outcropping part of the Lower Devonian, the Dartmouth Beds, has been transported northwards (NW to NNW) over the upper part of the Lower Devonian and the Middle Devonian. It is also suggested that the Middle Devonian (including the Plymouth Limestone) is thrust northwards over the Upper Devonian. Both these thrusts have previously been recognized in SE Devon. The aim of this paper is to present a new interpretation for the Variscan structure of SW Devon based on a geological cross-section from Plymouth to Bolt Tail (Fig. 1). The crosssection is on a similar line to that of Hobson (1976a). It is based on further detailed mapping of the SW Devon coast but has two further constraints. First is a seismic refraction estimate of depth to a Start schist type of basement below the Devonian sedimentary rocks of south Devon (M. Brooks & J. J. Doody, pers. comm.). The second constraint is the recognition of major thrusts in the Torquay and Dartmouth area by Coward & McClay (1983) which is directly along strike from the P l y m o u t h - B o l t

Tail section. Thrusts have also been recognized on the east side of the Dartmoor granite (Waters 1970; Selwood, Freshney & Durrance 1982; Wilcox 1982), along strike from the northern part of the present cross-section. It is also relevant that thrust-sheets or nappes have been mapped in the area between the Dartmoor and Bodmin Moor granites by Isaac, Turner & Stewart (1982) (see also Shackleton et al. 1982). In producing his structural cross-section Hobson (1976a) has reviewed and revised the stratigraphy of the area between Plymouth and Bolt Tail (see Ussher 1890; Hendriks 1951; Dineley 1961, 1966; Hobson 1976a,b). Hobson's revised stratigraphy is shown in Fig. 2.

FIG. 1. Geological map of south Devon showing previously recognized thrusts and newly proposed thrusts in the south Devon area. Information is from IGS sheets 22,348,349, 350, 355/356 (Waters 1970; Hobson 1976a; Selwood et al. 1982; Willcox 1982; Coward & McClay 1983). 113

114

T. J. Chapman, R. L. Fry & P. T. Heavey

FIG. 2. The geology of the coast between Plymouth and Bolt Tail based on IGS maps 348,349 and 355/356 (Dineley 1966; Hobson 1976a and new work by the authors) (see also Orchard 1978). Major F l fold axial traces are shown. A-B-C-D shows the line of the cross-section (see Fig. 3), and W-X and Y-Z the more detailed sections of Fig. 4.

Overall structure of S W Devon The oldest Devonian beds in Devon are the Dartmouth Beds (or Dartmouth slates) which run in an E - W belt from Dartmouth on the east Devon coast across to the west Devon coast SE of Plymouth Sound (Fig. 1). They extend further westward to Newquay on the north Cornwall coast (Hobson 1976b). Outcropping immediately to the north and south of the Dartmouth Beds are the younger Lower Devonian Meadfoot Group. This fact has led to the interpretation that the Dartmouth Beds occupy the core of the major E - W trending anticlinal structure called the Dartmouth antiform (Dineley 1966; Richter 1968; Hobson 1976a,b). However Hobson (1976a,b) also pointed out that the northern boundary of the Dartmouth Beds to the SE of Plymouth was a major fault. This fault which passes into the sea just north of A n d u r n Point, was estimated by Hobson (1976a) to have a downthrow of 3.7 km to the north. From the surface geology alone this is a reasonable estimate. It is interesting that immediately north of this fault the Meadfoot Group still has a gross younging to the south and lies in the core of an anticline which verges and faces to the north (see Hobson 1978 and Fig. 3a). On the east coast of south Devon the Meadfoot Group north of the same boundary also has a general

younging towards the south (Coward & McClay 1983). The core of the Dartmouth antiform does not apparently lie within the Dartmouth Beds. The northwards verging folds (themselves large scale) are apparently incongrous to the main antiformal structure (see discussion following Dineley 1966). Hobson (1976b) considered this variation in F 1 fold axial surface attitude to be a result of F : folding, that is the Dartmouth antiform is a late major fold. Coward & McClay (1983) have explained downward facing folds at Goodrington, south of Torquay, as a result of F 2 folding of northward verging and facing F1 folds, relating this to sequential thrust formation. In the interpretations presented in this paper we also suggest that thrusting associated with folding can explain such structural relationships.

Thrusts in the Plymouth area Figure 1 shows the postulated position of major thrusts in south Devon. The data, from several sources, include two new thrusts in the Plymouth area. One of these occurs at the base of the Middle Devonian Plymouth Limestone and the other at the base of the Dartmouth Beds although it is not exposed due to later normal faulting. The latter thrust is the westward continuation of Coward & McClay's (1983) Torbay

A cross-section through S W Devon

115

FIG. 3. (a) The geology along the line of the cross-section (see Fig. 2). (b) Structural interpretation based only on surface geology and following Hobson (1976a). The actual position of the basement is somewhat arbitrary but it does indicate the large step required to satisfy the surface geology. The postulated offlap is also arbitrary but is included to reduce the overall sedimentary thickness. (c) Structural interpretation based on surface geology but also constrained by seismic refraction data (M. Brooks and J. J. Doody pers. comm.); the identification a thrust east of Dartmouth, and its tectonic klippe at Torbay (Coward & McClay 1984); and a consideration of the ages of rocks at the surface necessitating ramps at depth. thrust. The evidence for the existence of this thrust is described below in the description of the cross-sections. The northern margin of the Middle Devonian Plymouth Limestone has generally been assumed to be faulted (Ussher 1907; Taylor 1951) and certainly E - W trending normal faults affect the boundary. However, at Billacombe east of Plymouth (SX 523544) the relationship between topography and outcrop pattern indicate that the northern margin of the Plymouth Limestone dips gently southwards. This boundary is exposed as a low angle thrust fault (dipping at less than 20 ~ in a q u a r r y at grid reference SX 519545. This same boundary between Middle Devonian slates and limestones to the south and Upper and MiddleDevonian slates to the north is interpreted as a

thrust (the Forder Green thrust) on the east side of the Dartmoor granite between South Brent and Newton Abbot (Waters 1970; Selwood et al. 1982; Willcox 1982). As Braithwaite (1965) points out, the general dip in the Plymouth Limestone is south with local overturned folds at its southern margin. In view of the low angle contact at the base of the limestones and its identification to the east as a thrust, the entire boundary is interpreted as southerly dipping, low angle thrust. This explains the apparent anomaly of a gross dip to the south but the presence of Upper Devonian rocks to the north of Middle Devonian. Chapman (1984) has recently described thrusts in the Staddon Grits and Jennycliff Slates on the east side of Plymouth Sound just south of the Plymouth Limestone (see Fig. 4a).

116

T. J. Chapman, R. L. Fry & P. T. Heavey

lqG. 4. Detailed cross-sections from parts of the SW Devon coast. Their positions are shown on Fig. 2. The main section (Fig. 3) is based in part on these.

The structural cross-section Figure 3(a) is a cross-section showing the surface geology with a reasonable extrapolation slightly above and below. Figure 4 shows detail from crucial parts of the main section. Figure 3(b,c) shows two alternative interpretative sections. Figure 3(a) shows the general younging to the south in the south of the section and the change in gross attitude from a gentle almost flat zone in the north and middle to a steeper zone in the south. It also shows F 1 folds which are overturned (verging and facing) to the north, becoming upright in the south as shown by Hobson (1976a). Coward & McClay (1983) have explained a similar variation in fold attitude in east Devon as resulting from a later phase of back folding. They take an early phase of folding in the Start complex as F1 assigning the main folds in south Devon to F2 although this is the first phase in the Lower Devonian. They therefore describe the backfolding as F 3. Intense zones of small-scale, post-main cleavage, F 2 folds (F 3 of Coward & McClay) occur in the south of this section as they do on the east coast and these indicate the backfolding event. Richter (1968) had previously considered the variation in orientation of the folds and accompanying axial planar slaty cleavage as a primary variation. Figure 3(a) also shows a gentle antiformal structure north of A n d u r n Point, previously

regarded as the northern part of the Dartmouth antiform. This is discussed below in the second interpretation. Recent geophysical data have been made available to the authors by M. Brooks and J. J. Doody of University College, Cardiff. A N - S trending seismic refraction profile indicates that a basement of similar velocity to the Start complex schists, lies at a depth of 2 - 3 km beneath the Devonian sedimentary rocks. This basement is not affected by any major steps except at the Start complex boundary fault. This information is used as a constraint on the interpretations described below (Fig. 3b,c). The first interpretation (Fig. 3b) is similar to that of Hobson (1976a) and is a logical consequence of the surface geology. There is an overall younging to the south in the southern half of the section; and if these beds are projected downwards in a layer-cake manner they would reach a considerable depth. In view of probably higher velocity basement at 2 - 3 km a postulated arbitrary offlap is included to reduce the thickness of the sedimentary sequence. The most important feature of this interpretation is the 3 km displacement on the fault running just north of A n d u r n Point. This fault has a breccia of probable Permian age along it (Hobson 1978). This would have to affect the basement as shown although its position is rather hypothetical being around a depth of 2 - 3 km. The possibility of a listric normal fault

A cross-section through S W Devon

117

FIG. 5. Possible evolutionary model along the line of the cross-section based on the second interpretation (Fig. 3c). Possible age ranges are from Coward & McClay after Dodson & Rex (1971). cannot be ruled out and offers a third interpretation. It is, however, thought unlikely because of the large displacement involved and the lack of evidence of rotation. The alternative and favoured interpretation is presented in Fig. 3(c). In this model a major thrust is introduced carrying the Darmouth Beds over the rocks lying to the north of the Andurn Point fault. This thrust is not exposed due to the A n d u r n Point fault which, however, still marks the map boundary between the upper and lower thrust sheets. The region of anomalous Start schist type basement is replaced by a Dartmouth Beds to Staddon Grits/Jennycliff Slates sequence. The beds are repeated by a major ramp structure. The attractions of this model are that it more closely satisfies the seismic refraction data described above in that there is a greatly reduced displacement on the A n d u r n Point fault, which becomes reversed relative to the first interpretation (compare Figs 3b and 3c). It also satisfies the constraint of a major thrust, being recognized, directly along strike east of Dartmouth, with Dartmouth Beds thrust over the Meadfoot Group and Staddon Grits and also the presence of a klippe of the same thrust sheet at Torquay (Coward & McClay 1983). The latter is directly along strike from where the postulated thrust would have been north of the A n d u r n Point fault prior to erosion. The Torquay k]tippe carries the same rocks, the Lower Devonian Meadfoot Group and Staddon Grits over Middle Devonian slates (see IGS sheet 350 and Coward & McClay 1983).

A blind imbricate thrust is postulated to account for the antiformal structure north of the A n d u r n Point fault (Figs 4a and 3c), and a ramp may also extend down into Start schist type basement further south. The ramps and imbricate thrusts on the section are necessary to explain the different ages of beds at the surface. The model can also explain the presence of a nearly flat or gently dipping belt north of a steep belt. Such a change in orientation of cleavage from steep in the south to gentle in the north is predicted by Sanderson's (1982) flexual flow model for a ramp structure. A late backfolding event (F 3 of Coward & McClay) is still necessary to give steep and southerly verging folds in the south. This interpretation does not include a simple single Dartmouth antiform. The southerly dipping beds south of the Andurn Point fault are related to one thrust at a ramp, whilst the partial antiform north of the fault is related to a different imbricate thrust. A possible evolutionary model along the line of the cross-section is presented in Fig. 4. The nappe translation direction as indicated by quartz slickensides on minor thrusts and beddings surfaces is towards the N N W and NW. Coward & McClay (1983) recorded a similar direction from extension lineations and slickensides. ACKNOWLEDGMENTS" The authors would like to

express their gratitude to J. J. Doody and M. Brooks for supplying unpublished geophysical information on South Devon. J. J. Doody also provided useful comments on an early draft of this paper.

118

T. J. Chapman, R. L. Fry & P. T. Heavey References

BRAITHWAITE, C. J. R. 1965. A possible reinterpretation of the structure of the Plymouth Limestones. Proc. Ussher Soc. l, 180-82. CHAPMAN,T. J. 1984. A guide to the structure of the Lower to Middle Devonian Staddon Grits and Jennycliff Slates on the east side of Plymouth Sound, Devon. Proc. Ussher Soc. 460-4. COWARD, M. P. & MCCLAY, K. P. 1983. Thrust tectonics of South Devon, J. geol. Soc. London, 140, 215-28. DINELEY, D. L. 1961. The Devonian System in South Devonshire. Fld Stud. 1, 121-40. 1966. The Dartmouth Beds of Bigbury Bay, South Devon. Q. Jl geol. Soc. Lond. 122, 187-217. DODSON, M. H. & REX, D. C. 1971. Potassium-argon ages of saltes and phyllites from south-west England. Q. Jl geol. Soc. Lond. 126, 465-99. HENDRIKS, E. M. L. 1951. Geological Succession and structure in western south Devonshire. Trans. R. geol. Soc. Corn. 18, 255-95. HOBSON, D. M. 1976a. A structural section between Plymouth and Bolt Tail, south Devon. Proc. geol. Ass. 87, 27-43. 1976b. The structure of the Dartmouth Antiform. Proc. Ussher Soc. 3, 320-32. 1978. The Plymouth Area. Geol. Ass. Guide No. 38. ISAAC, K. P., TURNER, P. J. • STEWART, U. J. 1982. The evolution of the Hercynides of central SW England. J. geol. Soc. London, 139, 521-31.

ORCHARD, M. J. 1978. The conodont biostratigraphy of the Devonian Plymouth Limestone, South Devon. Palaeontology, 21, 907-55. RICHTER, D. 1968. Die tectonische Bangeschichte von Sud-Devonshire als Beispiel einer Mehrphrasigen variszischen Pragung. Geol. Rdsch, 57, 425-45. SANDERSON, D. J. 1982. Models of strain variation in nappes and thrust sheets. Tectonophys. 88, 201-33. SHACKLETON, A. C., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in SW England. J. geol. Soc. London, 139, 533-41. SELWOOD, E. B., FRESHNEY, E. C. & DURRANCE, E. M. 1982. The Variscan structures. In: DURRANCE, E. M. & LAMING, D. J. C. (eds) The Geology o f Devon, 66-84, University of Exeter. TAYLOR, P. W. 1951. The Plymouth Limestone. Trans. R. geol. Soc. Corn. 18, 146-214. USSHER, W. A. E. 1890. The Devonian rocks of South Devon. Q. Jl geol. Soc. London. 46, 487-517, 1907. The geology of the country around Plymouth and Liskeard. Mem. geol. Surv. U.K. 156 pp. WATERS, R. A. 1970. The Variscan structures of eastern Dartmoor. Proc. Ussher Soc. 2, 191-7. WILLCOX, A. D. 1982. An introduction to the geology of the area between Buckfastleigh and Ivybridge. Proc. Ussher Soc. 5, 289-95.

T. J. CHAPMAN, R. L. FRY & P. T. HEAVEY, Geology Division, Plymouth Polytechnic, Plymouth PL4 8AA.

Interpretations of seismic and gravity surveys over the eastern part of the Cornubian platform J. W. F. Edwards SUMMARY: Gravity interpretations suggest that the Cornubian granite batholith continues WSW from Land's End as far as 8~ and that the Haig Fras granite is part of a separate, parallel batholith. South-east of the Cornubian batholith there is a zone where seismic sections show dipping reflectors which may be thrusts. The axis of the Cornubian batholith is displaced dextrally between the Scilly Isles and Land's End. There are dextral displacements of the axes of the Cornubian batholith and of the Haig Fras batholith at about 7~ and on the line of these two displacements is an intrusion of high density. A deep horizontal reflector underlies the Cornubian batholith near the Scilly Isles. The Cornubian platform is that part of the continental s]helf which links the Cornubian peninsula of SW Britain with the Goban Spur on the continental margin, and which separates the south Celtic Sea basin from the South Western Approaches basin (Fig. 1). This paper describes interpretations of surveys over the part Of the platform east of 8~ Figure 2(A) shows the geology of the area based on a published sheet (Evans 1982) and pre-publication sheets of the solid geology map which have been prepared by C. D. R. Evans and B. N. Fletcher. Granite crops out in four places on the platform: at Haig Fras, around the Scilly Isles, at Seven Stones, and around Land's End (Smith et al. 1965; Exley 1966). In each case, the granites are

intruded into grey slates which are thought on lithological grounds to be of Devonian or Carboniferous age. On the platform there is a thin cover of a few hundred metres of early Cretaceous and younger sedimentary rocks. In the basins there are thick successions of Permo-Triassic and possible Jurassic sedimentary rocks (Avedik 1975; Kamerling 1979) but these are almost wholly hidden by the early Cretaceous and younger strata, except to the north west of Scilly where there are outcrops o f sandstones ascribed to the Permo-Triassic. These sandstones lie in the Haig Fras sub-basin which crosses the Cornubian platform and joins the two main basins (Evans et al. 1981; Evans & Fletcher 1978). During the Palaeogene, deposition was restricted to the main basins but was more extensive during the Neogene, following the Oligocene transgression.

Gravity survey

FIG. 1. Regional setting. Stippling shows thick accumulations of post-Carboniferous sediments. VF = Variscan front; CBB = Cardigan Bay basin; NCSB, SCSB = north and south Celtic Sea basin; CP = Corrmbian platform; SWAB = South Western Approaches basin; GS = Goban Spur; HF = Haig Fras; HFSB = Haig Fras sub-basin; SI = Scilly Isles; BCB = Bristol Channel basin; FB = Fastnet basin. Late Variscan granites in black. Outline of Fig. 2 dotted. After Gardiner & Sheridan (1981).

The Bouguer anomaly values in the area are shown in Fig. 2(B), which is taken from Edwards (1981), D u n h a m (1975), Armstrong & Howell (1979), Armstrong & Rollin (1979), and two pre-publication sheets. The four outcrops of granite occur in areas with a negative gravity anomaly, and it is suggested here that all the strong negative anomalies seen in Fig. 2 ( B ) are caused by granite bodies having a density lower than that of the surrounding crust; more so because interpretations of the seismic data show that sedimentary basins cannot be responsible for the negative anomalies SW of Land's E n d and N E of Haig Fras. The strong negative anomaly running along the axis of the Cornubian peninsula has been shown to be caused by a granite batholith (Bott et al. 1958) and from Fig. 2 it is suggested that the granite 119

120

J. W. F. Edwards

FIG. 2. (A) Solid and basement geology map and (B) Bouguer anomaly map. Simplified from B.G.S. 1:250,000 series.

Surveys o f the Cornubian platform batholith under the Cornubian peninsula continues WSW from Land's End as far as 8~ I propose to call the whole batholith from 4~ to 8~ the Cornubian batholith, and to interpret the outcrops of granite as cupolas of this batholith. The Haig Fras batholith is separate from the Cornubian batholith and runs parallel to it. Both batholiths are elongate with their axes in a direction different from the conventional direction of the Variscan front, as shown in Fig. 1. The Cornubian batholith may be related to granites postulated to the east under Dorset and Hampshire (Wills 1973) and to the granites postulated to the SW towards the edge of the continental shelf (Day & Williams 1970), and the granite sampled on the edge of the continental shelf (Pautot et al. 1976). Tile axis of the Cornubian batholith is displaced dextrally in the straits between Scilly and Land's End, and there are also displacements of the axes of the Haig Fras and Cornubian batholith at 7~ and 7~ respectively. On the line of these last two displacements, and SSE of the Cornubian batholith, is a positive anomaly suggesting the existence of a dense body in the basement, centred on 49~ 7~ This is named here the Madura body after a wreck in the vicinity. The body was modelled by Ham (Edwards et al. 1983), assuming a density contrast of 0.13 Mg m -3. The model structure he derived has steep sides and comes within 2 km of the surface. Seismic profiles show that the Cretaceous sedimentary rocks thin gradually to a few hundred metres over the area, but it is not clear whether the Madura body lies immediately below this cover. The high density of the body suggests a composition which is basic rather than acidic, and the weak anomaly on the magnetic profiles over the body suggests that basaltic lavas are absent. Beyond this, nothing is known of the petrology or age of the body.

121

found at the end of the record between 5.8 and 6.0 s, corresponding to a depth of 17 km. The lateral extent is shown in Fig. 2, and corresponds to the part of the Cornubian batholith around the Scilly Isles. The reflectors may be restricted to an area under the batholith, or may be more extensive with detection under the batholith being a consequence of velocity pullup. If they represent the base of the batholith, the base is deeper than is suggested from the refraction interpretation. Alternatively, the reflections may be from a cumulate under the batholith, or from a d6collement plane under it (Shackleton et al. 1982). The second set of reflectors is seen on seismic sections as five or more parallel events dipping to the SE at a true angle of 30 ~ and are similar to those illustrated in Leveridge et al. (this volume), which are interpreted as thrusts. These reflectors have not yet been studied in detail in the South Western Approaches and, in particular, individual reflectors have not been correlated from line to line, but the reflector set as whole have been found to fall in a zone elongated N E - S W south of the Cornubian batholith (Fig. 2). From this preliminary study, the zone of reflectors appears to be offset along the lines of displacement of the Cornubian batholith, but the horizontal normal offset of the zone is sinistral.

Profiles

Three profiles have been prepared across the Cornubian platform (Fig. 3) along the lines in Fig. 2(B) based on an integrated interpretation of the gravity, seismic and sampling evidence. The vertical exaggeration of the profiles is 3:1. Further details, including the values of the observations and the calculated anomalies, are ,in Edwards (1983). The density contrast of the granites with the basement was taken as - 0 . 1 3 Mg m -3 as this puts the bottom of the batholith Seismic surveys at the depth found by Holder & Bott (1971), A long refraction survey run SSW from Land's and is also in agreement with density measureEnd (Holder & Bott 1971) detected the Moho ments on samples from outcrops (Bott et al. at 27 km and an increase in velocity starting at a 1958). Models of the onshore part of the Cordepth of 10-12 km, which may be the bottom nubian batholith prepared by A1-Rawi (1980) of the batholith. The whole area shown in Fig. 2 show that slight variations in the density conis covered by the British Geological Survey's trast of the batholith changes the depth of its high resolution seismic surveys, and much of base, but not the general form of its section. By the area is also covered by deep seismic surveys. changing the density contrast of the batholith, it Deep seismic sections obtained commercially would be possible to put the bases of both bathand held in confidence by the B.G.S. show two oliths at the same depth on every profile. The sets of reflectors from within the basement. The density of the crust and the density of the bathfirst set is horizontal, or nearly so, and is oliths have been assumed to be uniform with

122

J. W. F. Edwards

FIG. 3. Profiles along lines in Fig. 2. Values in brackets are density contrasts in Mg m -3. Vertical exaggeration 3:1. depth. The density contrasts of the sediments have been chosen to give agreement between the depth determined from the seismic data, and the gravity anomalies. Assuming a density of 2.7 Mg m -3 for the upper part of the crust, the resulting densities are in agreement with measurements made in boreholes on rocks of

the same age and facies. The regional field has been assumed to be constant along each profile, with values of 40.0, 42.9 and 46.3 reGal for profiles A, B and C respectively. These values are high and, as the Moho is not particularly shallow, they suggest that the lower crust is dense.

S u r v e y s o f the C o r n u b i a n p l a t f o r m The gravity modelling is based on the Talwani algorithm, which assumes that the bodies have the same cross-section throughout their length. The contours of the Bouguer anomaly show that this assumption is reasonable on a large scale, but the profiles show that this is not so in detail. The m e t h o d is insensitive to changes in boundaries which are at depth and not close to vertical, which makes it difficult to test the suggestion of Shackleton et al. (1982) that the granites have been intruded from the SSE. The batholiths have been found to have a bell-shaped cross-section, like the onshore part of the Cornubian batholith (A1-Rawi 1980), although in general the edge of the batholiths are steeper on the NNW side than on the SSE side. Each profile shows the Haig Fras subbasin, the edge of the south Celtic Sea basin, and the edge of the South Western Approaches basin. In Fig. 3 the thin Tertiary sediments have been combined with the Cretaceous sediments for clarity. When the profiles were first modelled, a positive residual anomaly was found in three places, and dense bodies have been introduced into the profiles to explain these residuals. A sharp anomaly seen on a magnetic profile at 50~ 8~ may be caused by a dyke of reversed polarization rising from the body

123

N N W of the Haig Fras batholith on line C, but there is no supporting evidence for the other two bodies. These residuals, and also the projection on the N N W edge of the Cornubian batholith on line B, could also be removed by changing the regional field, or assuming that the cross-section of the bodies varies suddenly along their length. The SSE ends of lines A and B extend into the zone of thrusts seen on the seismic profiles. On line A it was necessary to insert a volume with a positive density contrast above the lowest thrust plane (shown cross-hatched in the profile), but this was not necessary on line B where the lowest thrust plane is shown by a broken line. On both lines the edge of the South Western Approaches basin coincides with the subcrop of the lowest thrust, suggesting that the basin was formed by m o v e m e n t along the thrusts.

ACKNOWLEDGMENTS: I would like to thank Mr J. Bulat of the Hydrocarbons Unit, Institute of Geological Sciences, for discussion of the interpretation of the seismic data and to my colleagues in the Marine Geophysics Unit, who gave assistance and guidance throughout this study. This paper is published with the approval of the Acting Director, institute of Geological Sciences (NERC).

References AL-RAWi, F. R. J. 1980. A geophysical study of the deep structure in south-west Britain. Ph.D. Thesis. University of Wales (University College Cardiff). ARMSTRONG, E. J. & HOWELL, P. M. 1979. Bouguer Gravity Anomaly Map, Scilly, sheet 49~176 1:250,000 series. Institute of Geological Sciences. -& ROLtJN, K. E. 1979. Bouguer Gravity Anomaly Map, Lizard, sheet 49~176 1:250,000 series. Institute of Geological Sciences. AVEDIK, F. 1975. The seismic structure of the Western Approaches and the Armorican Continental Shelf and its geological interpretation. In: WOODLAND,A. W. (ed.) Petroleum and the Continental Shelf of North West Europe, Vol. 1, 29-43. BOTT, M. H. P., DAY, A. A. & MASSON-SMITH,D. 1958. The geological interpretation of gravity and magnetic surveys in Devon and Cornwall. Phil. Trans. R. Soc. A, 251, 161-91. DAY, G. A. & WILLIAMS,C. A. 1970. Gravity compilation in the NE Atlantic and interpretation of gravity in the Celtic Sea. Earth planet. Sci. Lett. 8, 205-13. DUNHAM, K. C. 1975. Bouguer Gravity Anomaly Map, Land's End, sheet 50~176 1:250,000 series. Institute of Geological Sciences. EDWARDS, J. W. F. 1981. Bouguer Gravity Anomaly

Map, Haig Fras, sheet 50~176 1:250,000 series. Institute of Geological Sciences. 1983. Gravity interpretations of the Haig Fras and Scilly granite batholiths. Marine Geophysics Unit Report No. 134. Institute of Geological Sciences, England. --, ARMSTRONG, E. J. & HAM, D. 1983. Geophysical interpretations south-west of the Scilly Isles. Rep. Inst. geol. Sci. 83/10. EVANS, C. D. R. 1982. Solid Geology Map, Scilly, sheet 49~176 1:250,000 series. Institute of Geological Sciences. -& FLETCHER, B. N. 1978. Recent geological exploration of the continental shelf off SW England. J. geol. Soc. London, 135, 478. , Loan', G. K. & WARRINGTON, G. (compilers) 1981. The Zephyr (1977) wells, South Western Approaches and western English Channel. Rep. Inst. geol. Sci. 81/8. EXLEY, C. S. 1966. The granite rocks of Haig Fras. Nature, 210, 365-7. GARDINER, P. R. R. & SHERIDAN, D. J. R. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan Front. J. struct. Geol. 3, 317-31. HOLDER, A. P. & BOTT, M. H. P. 1971. Crustal structure in the vicinity of south-west England. Geophys. J. R. astr. Soc. 23, 465-89.

124

J. W. F. E d w a r d s

KAMERLING, P. 1979. The Geology and Hydrocarbon Habitat of the Bristol Channel Basin. J. Petrol. Geol. Beaconsfield, 2, 75-93. PAUTOT, G., RENARD, V., AUFFRET, G. A., PAS'IOURE~I, L. & DE CHARPAI, O. 1976. A granite cliff deep in the North Atlantic. Nature, 263, 669-72. S HACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in SW England. J. geol. Soc. London, 139, 533-41.

SMITH, A. J., STRIDE, A. H. & WHITTARD, W. F. 1965. The geology of the Western Approaches to the English Channel, IV: a recently discovered Variscan granite WNW of the Scilly Isles. Colston Pap. 17, 287-301. WILLS, L. J. 1973. A palaeogeological map of the Palaeozoic floor below the Upper Permain and Mesozoic formations. Mere. geol. Soc. Lond. 7.

J. W. F. EDWARDS,British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 3LA, Scotland.

Thin-skinned tectonics, basement control and the Variscan front Robert M. Shackleton SUMMARY: The thin-skinned interpretation of Variscan structures in southern England, Wales and Ireland, based on structural and seismic evidence, is used to define the Variscan front as the northern limit of the thin-skinned structures. The original front was where displacement on the d6collement surface diminished to zero. The present front, eroded back from its original line, is near the line previously accepted as the Variscan front. Caledonian (and in places older) structures control the Variscan deformation north of the front. They can also be traced south of it, as post-Variscan structures which have propagated up through the d6collement. A zone of ENE-WSW structures in south Cornwall is thought to reflect the influence of Cadomian basement where the basal thrust has cut down into it.

Thin-skinned tectonics It has been shown elsewhere (Shackleton et al. 1982) that the Variscan structures of SW England require an underlying d6collement zone. This is (a) because the overall recumbency of the slaty cleavage and the high strains must imply the existence of a gently inclined shear zone (Ramsay & Graham 1970) and (b) because the amount of shortening in the Upper Palaeozoic sediments, of the order of 150 km in SW England (Shackleton et al. 1982) cannot apply to the whole crust because after allowing for the previous thinning, which can be calculated from the thickness of sediments (cf. Dewey 1982), the estimated shortening, applied to the whole crust, would have increased its thickness far more than can have occurred. The d6collement surface which is predictable from the structures has been independently identified and mapped by seismic reflection in the Rhenohercynian zone of Germany (Meissner et al. 1981), in Belgium and northern France (C16ment 1963) and in Wiltshire (Kenolty et al. 1981); similar d6collements are found in the Appalachians (Brewer & Oliver 1980). The d6collement so identified is so extensive and the displacement so large that it must continue far beyond the area where it has already been mapped, for example beneath the Celtic Sea. Two features of such a d6collement surface may be emphasized. First, it will be inclined, even if gently, so as to cut up-section in the direction of propagation (in the Variscan, northwards), and shear zone geometry implies that cleavage, thrusts and associated structures will steepen upwards, away from the shear zone, so that their form is listric. Secondly, the displacement on the d6collement surface increases southwards away from its external limit, the tip line (Butler 1982), being at any

point equal to the cumulative shortening external to that point. At the tip line there is no displacement.

The Variscan front It is a cartographic convenience, when making tectonic maps, to be able to draw a line which represents the limit of an orogenic belt, in this case the Variscan. This line is described as the orogenic front. Such a line has no special meaning in plate tectonic terms, in contrast to a suture which separates two plates, and while the limit of the deformation is sharply defined in some places, elsewhere, as for example in both England and Ireland, it is diffuse and strong deformation, with tight folding and associated slaty cleavage, occurs far to the north of the conventionally accepted position of the Variscan front. The concept of a front has consequently come to seem meaningless and misleading (Matthews 1978). However, the d6collement and the thin-skinned tectonics above it do have a definite and definable limit, which is of structural importance and does coincide approximately with the conventional position of the Variscan front. It is proposed that the Variscan front be defined as the present northern limit of the d6collement. The d6collement may end as a blind thrust (Thompson 1981), one on which the displacement diminishes to zero below the present erosion surface, so that the thrust never outcrops. It will then be more difficult to locate the front precisely but it nonetheless has a real existence. If on the other hand erosion has removed the termination of the d6collement, what is exposed will be the thrust, emerging south of the original front and forming what is often described as a thrust front. The original front was external to this, at a distance which may be calculated from 125

R. M. Shackleton

126

the displacement where the d6collement outcrops and the rate of diminution of that displacement northwards. This original front is in the air. The present front, which is the one with most practical value, thus consists of two elements, those representing the outcrop of the d6collement surface, at a thrust front, and those where the thrust is blind, and the tip line is below surface. The original front may be far from the present one; for example in the NW Highlands of Scotland, the original Caledonian front was probably a considerable distance west of Durness, well beyond the present front. In the Variscan, the two are close together and the distinction is less important. To keep as close as possible to current usage it is suggested that front, unqualified, should be used for the present front, as distinct from the original front, which can be so specified where it needs to be distinguished.

Thin-skinned tectonics and basement control The sediments above a d6collement have usually not been tectonically deformed previously. When they are detached, their deformation, uninfluenced by pre-existing structures, will depend essentially on the relative motion of the underlying plate, modified by variations in frictional resistance on the d6collement surface and by differences in the mechanical response of the sediments that are involved. Geometrically therefore, the pattern of thin-skinned structures tends to be rather simple and regular in orientation. On the contrary the structures in the sediments beyond the limit of d6collement (as well as in any that remain attached to the basement below the d6collement) will be influenced and controlled by structures in the basement to which they remain attached. Depending on the difference in orientation between the basement structures and the thin-skin structures, the

change in orientation of the structures in the sedimentary cover may be large or small. In England, Wales and Ireland the angle between the Caledonian and Variscan trends varies, but the change, from Variscan structures which are not controlled by basement to those that are, is generally distinct. Thus the Variscan front is marked not only by a change in intensity of deformation and by the limit of the d6collement but also by a change in orientation of the structures. The latter change is well seen both in south Wales (Hancock et al. 1983) and southern Ireland. The interred line of the Variscan front, as here defined, and located by the criteria indicated above, is shown in Fig. 1. Note that the line so drawn, in southern Ireland, is aligned with the line farther east. It may be argued that the differences between the unconstrained structures above the d6collement and those controlled by earlier structures in rocks external to, or below, the d6collement does not preclude the existence of deeper Variscan thrusts. Such thrusts (not d6collements) would need to cut indiscriminately through older structures, which would nevertheless control the orientation of the Variscan folds. This seems improbable. It seems more likely that the crust beyond the front and its continuation underneath the d6collement, have been heterogeneously deformed by the amplification or reactivation of older structures. The basement continues beneath the d6collement. It is therefore to be expected that later (post-Variscan) structures, initiated in, and propagated upwards from the basement, through the d6collement, will be influenced or controlled by the basement grain. Such control of post-Variscan structures by the Caledonian basement is spectacularly displayed in the Celtic Sea (Gardiner & Sheridan 1981). Faults controlled by the Caledonian Bala and Menai Straits faults can be traced from drilling evidence and gravity signature, as limits of Mesozoic graben, south-westwards across the

FIG. 1. Tectonic map of South England, Wales, Ireland and the Celtic Sea showing the trend of Variscan, Caledonian, Cadomian and other structures, the position of the Variscan front and the alignment of the granite batholiths. AA BB CC DD EE FF

Estimated line of Variscan front. Supposed southern continuation of Malvern Axis (Hawkins 1942). Ophiolite zone (Lefort 1977). Trend of Cornish batholith. Trend of Haig Fras granite. Northern limit of Variscan structures with Cadomian trend.

GG HH JJ KK

Trend of Leinster granite. Bala fault. Church Stretton fault (near Caledonian front?) Wexford boundary lineament.

Based on Dunning (1966); Lefort (1977); Phillips et al. (1979); and Gardiner & Sheridan (1981).

The Variscan front

127

i II Ill irn

ii

I

:".%

11

I

I

, "i. . ".. \, ." ~ i""~ . .i"" . . ' .... -,. . . ~ % ~I~ '

"

"

':"

b'...

%

"-.... 0

ii I i~ jj

"'~

9

,.

.,~ ..

~176176 ~176

._~ ~

9,,.

o

!

"~

"0o

"5

(/)

\

~

0~ -~

i,. =

128

R. M. S h a c k l e t o n

southern Irish Sea far into the Celtic Sea (Gardiner & Sheridan 1981). It is clear that there is no large lateral displacement of these Caledonian structures at the Variscan front. It is the continuation of the posthumous Caledonian structures across the Variscan front which led to the suggestion that the Variscan structures themselves turned south-westwards from SW England and that the Variscan structures in SW Ireland were entirely separate (Gardiner & Sheridan 1981), an interpretation which is rejected here for reasons given above. The SE Caledonian front could be defined in the same way as the Variscan front is defined here, if the d6collement hypothesis of Coward & Siddans (1979) were confirmed; if so it probably lies near the Church Stretton fault. Further east, N - S Malvernian structures may be traced southwards under the Variscan d6collement (Hawkins 1942; Kellaway & Hancock 1983). Southwards, the Variscan d6collement must, unless it is a gravitational structure for which there is no evidence, cut downwards into the basement. The Variscan structures above it will then again be influenced by the basement grain instead of being unconstrained by pre-Variscan structure. This change is unlikely to be sudden because the thickness of basement above the thrust will only increase gradually. It is suggested that this above-thrust basement influence can be recognized in south Cornwall as the line where the trend of the Variscan structures changes from E - W to E N E - W S W (Sanderson & Dearman 1973). This E N E - W S W trend has sometimes been referred to as Caledonian but as it is far beyond the Caledonian front this is misleading. It appears to be parallel to the Cadomian structures in Normandy, as well as to a zone of probable ophiolites (Lefort 1977) under the English Channel, and it is suggested that this zone of E N E - W S W Variscan structures in south Cornwall reflects a Cadomian

influence. The Icartian ( - 2 0 0 0 Ma) gneisses in the Channel Islands (Roach 1977) appear to show a N N W - S S E trend. This might be regarded as a detached continuation of the Birrimian ( - 2 0 0 0 Ma) structures of W and NW Africa, which trend consistently N - S or N N W - S S E over a vast area. A further problem concerns the emplacem e n t of the SW England granites. These are well known to lie in a zone which trends E N E - W S W , oblique to the E - W Variscan structures into which they are intruded. The Haig Fras granite is shown by geophysical work to have a similar trend (Edwards, this volume). These orientations suggest basement control, as previously suggested by Dunning (Dunning & Max 1976). It has been argued (Shackleton et al. 1982) that the crust into which the granites were intruded was not then thick enough to melt, although isotopic evidence indicates a significant crustal contribution, and so it was suggested that the granitic magma was injected laterally. If so, it would seem that the same pre-Variscan structures which controlled the Variscan structures in south Cornwall also controlled the intrusion of the granites.

Conclusions The thin-skinned interpretation of Variscan structures in the Rhenohercynian zone of western Europe explains the features which have in the past been implicitly used to locate the Variscan front. This front is defined as the northern limit of thin-skinned tectonics above a d6collement surface. The Variscan structures north of the front are strongly influenced by the Caledonian basement grain. Pre-Variscan structures can be traced south of the Variscan front from posthumous structures which have propagated up through the thin-skinned Variscan structures from the underlying basement.

References BREWER, J. A. & OLIVER,J. E. 1980. Seismic reflection studies of deep crustal structure. Ann. Rev. Earth planet. Sci. 8, 205-30. BUTLER, R. W. H. 1982. The terminology of structures in thrust belts. J. struct. Geol. 4, 239-45. CLI~MENT, J. 1963. R6sultats pr61iminaires des campagnes gEophysiques de reconnaissance dans le permis de recherche 'Arras et Avesnes' de l'association Shell Franqaise--P.R.C.B.-S.A.F.R.E.P.--objectif du forage profond Jeumont--Marpent No. 1. Soc. gdol. Nord. 83, 237-41. COWARD, M. P. & SIDDANS,A. W. B. 1979. The tec-

tonic evolution of the Welsh Caledonides. In: HARRIS, A. L., HOLLAND,C. H. & LEAKE,B. E. (eds) The Caledonides o f the British Isles--Reviewed. Spec. Publ. geol. Soc. Lond. 8, 187-98. Scottish Academic Press, Edinburgh. DEWEY,J. F. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. 139, 371-412. DUNNING,F. W. 1966. Tectonic Map o f Great Britain, 1: 189 Institute of Geological Sciences. & MAX, M. D. 1976. Explanatory notes to the geological map of the exposed and concealed Precambrian basement of the British Isles. In: HARRIS, A. L., SHACKLETON,R. M., WATSON,J.,

The Variscan front DOWNIE, C., HARLAND, W. B. & MOORBATH, S. (eds) Precambrian: a correlation o f the Precambrian rocks o f the British Isles. Spec. Rep. geol. Soc. 6, 11-14. GARDINER, P. R. R. & SHERIDAN, D. J. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan front. J. struct. Geol. 3, 317-37. HANCOCK, P. L., DUNNE, W. M. & TRINGHAM, M. E. 1983. Variscan deformation in Southwest Wales. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 47-73. Hilger, Bristol. HAWKINS, H. L. 1942. Some episodes in the geological history of the South of England. Q. Jl geol. Soc. Lond. 98. i-lxx. KELLAWAY, G. A. & HANCOCK,P. L. 1983. Structure of the Bristol District, the Forest of Dean and the Malvern Fault Zone. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 88-106. Hilger, Bristol. KENOLTY, N., CHADWICK, R. A., BLUNDELL, D. J. & BACON, M. 1981. Deep seismic reflection survey across the Variscan Front of southern England. Nature, 293, 451-3. LEFORT, P. J. 1977. Possible 'Caledonian' subduction under the Domnonean domain, North Armorican area. Geology, 5, 523-26. MATTHEWS, S. C. 1978. Caledonian connexions of Variscan tectonism Z. dt. geol. Ges. 129, 423-8. MEISSNER, R., BARTELSEN,H. & MURAWSKI,H. 1981.

129

Thin-skinned tectonics in t h e northern Rhenish Massif, Germany. Nature, 290, 399-401." PHILLIPS,W. E. A., FLEGG, A. M. & ANDERSON,T. B. 1979. Strain adjacent to the Iapetus suture in Ireland. In: HARRIS, A. L., HOLLAND, C. H. & LEAKE, B. E. (eds) The Caledonides o f the British isles--Reviewed. Spec. Publ. geol. Soc. Lond. 8, 257-62. Scottish Academic Press, Edinburgh. RAMSAY, J. G. & GRAHAM, R. H. 1970. Strain variation in shear belts. Can. J. Earth Sci. 7, 786-813. ROACH, R. A. 1977. A review of the Precambrian rocks of the British Variscides and their relationships with the Precambrian of N.W. France. In: La Chaine Varisque Europe Moyenne et Occidentale. Coll. int. CNRS, Rennes, No. 243, 61-79. SANDERSON, D. J. & DEARMAN, W. R. 1973. Structural zones of the Variscan fold belt in S.W. England, their location and development. J. geol. Soc. London, 129, 527-36. SHACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in S.W. England. J. geol. Soc. London, 139, 533-41. THOMPSON, R. I. 1981. The nature and significance of large 'blind' thrusts within the northern Rocky Mountains of Canada. In: MCCLAY, K. R. L. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 449-62. Blackwell Scientific Publications, Oxford.

ROBERT M. SHACKLETON, Department of Earth Sciences, The Open University, Milton Keynes MK7 6AA, England.

The Ribblesdale fold belt, NW England--a Dinantian-early Namurian dextral shear zone Russell S. Arthurton SUMMARY: Folds and faults in various en dchelon patterns affect the Tournaisian-early Namurian sedimentary rocks of the Craven basin, a rift basin which includes Dinantian and possibly late Devonian strata some 5 km thick. The structures are interpreted as a response to regional dextral shear, which accompanied sedimentation from mid-Dinantian to early Namurian times at Least. The principal Ribblesdale folds are viewed as products, in the Carboniferous cover, of primary wrenching on inherited fractures in the Lower Palaeozoic basement. Some en dchelon sets of minor folds are regarded as products of secondary wrenching in the axial zones of 'primary' folds. The block-basin transition zone south of the North Craven fault appears to have been affected by dextral transtension during the late Brigantian (late Dinantian).

The Ribblesdale fold belt (Phillips 1836) is a product of the Hercynian deformation of the Craven basin, a ?late D e v o n i a n - C a r b o n i f e r o u s intracratonic sedimentary basin some 300 km north of the Variscan front (Fig. 1). It comprises an e n O c h e l o n set of N E - S W - t r e n d i n g anticlines with broad intervening synclines, and is cut by faults or fault zones trending N W - S E , including the South Craven fault system (Fig.

FIG. 1. The Dinantian rift basins of northern Britain in relation to the Variscan front. (After Leeder 1982, with minor amendments; position of Variscan front by N. J. P. Smith.)

2). The fold belt extends E N E - W S W over at least 80 km, and is up to 25 km wide. It lies between the Askrigg block to the north and the Pennine and Rossendale blocks (Miller & Grayson 1982) to the south. The Craven basin was interpreted by Leeder (1982) as being one of several rift basins within a late D e v o n i a n - D i n a n t i a n rift province which extended northwards from St George's Land (Fig. 1); the basins having formed in areas of less buoyant crust as a result of crustal stretching. The province lay within a zone of shear which was an expression of right-lateral transform faulting in the Canadian Maritimes extending north-eastwards into the British Isles during Dinantian times (Dewey 1982). According to Leeder (1982) the rift province became largely defunct during the Namurian and Westphalian, with fault-bounded, 'rift' subsidence being replaced by general 'sag' subsidence. The Craven basin, like the Northumberland basin (Fig. 1), was uplifted during Stephanian times; it became deeply eroded before the deposition of the Permo-Triassic sandstone of which remnants are preserved resting on midDinantian rocks near Clitheroe (Earp et al. 1961). Major faults both within and peripheral to the basin were active d u r i n g t h e Dinantian and early Namurian (e.g. H u d s o n 1930). Some of these faults may have had a Variscan component of lateral displacement (Wager 1931; Hudson & Mitchell 1937; Hudson & Dunnington 1944; Earp et al. 1961; Moseley & A h m e d 1967; Westoll in discussion to Moseley & A h m e d 1967). The study of one of the Ribblesdale folds (Skipton anticline) by Hudson & Mitchell (1937) led them to the conclusion that in addition to a post-Millstone Grit (Namurian) phase 131

132

R. S. A r t h u r t o n

of folding there were several earlier phases during the mid-Dinantian-early Namurian. This paper examines the relationship between faults and folds in the central part of the Ribblesdale fold belt and its immediate surroundings, and in particular assesses the significance of the en dchelon patterns of folds and minor faults.

~ ' ~ ~ 0

Stratigraphy The Craven basin (Fig. 2) is bounded to t h e north, by a fault zone including the North and Middle Craven faults. Westwards from Settle the northern margin of the basin during the Dinantian is speculative. Gravity data suggest the existence of high density ?Lower Palaeozoic

/

~St~

Askrigg Block

Farrn~f'----~ CR~'Ab ~ - ~ - ~ F A U

Settle ~L~RA~'~NR~EEF9 BEL--EET-~~

,++..,~ ,,,,.+~ ................~ % . ~ !

_

+o++O+.+++.......~">

~ ~

~0,.,,~. . . ~ ~

"

~c

__

- ~

~-~

~ f ~ , < o "

+"+,,~/;+"

,~ , O + ~ /,,~ pv I"

-

/.

M,:;lh+,.,.,

'

\ \~V o~\ \~.h

Swinden ~_ [ Borehole@]

+

_x/ " , 0~

"~l ....

,~bklpton

~n

r +'' .,++.,W+""

\ +: + + , + o + ; + ~ ' 2 ~ : ~ + + +

) ~,,,o,o_~

+o_

....+ ............... ..........-" k..."("

=

..... ~-'gOL

--

,,~ ...'""'Y .oO."

(~ Borehole

~ I

10 Km

J

3o

Holme Chapel Borehole @

1

<=< u_

~+oul+wo+th

~

%........ ~'~ Burnleyll

+

1_

rr

<03

_

g ~

L

Craven

L

Basin

+z

1__i

~ A > + + + + . + § I ~ ~

~skrigg Block

z z

"

WESTPHAL/AN

•+

" " " - " ; " : .....-........... . ..... :.:.:.:.:i;.""

FIG. 2. (A) The central part of the Ribblesdale fold belt, showing the principal anticlines and faults. (Based on published and unpublished maps of the British Geological Survey.) (B) Cross-section on line shown in (A). The principal chrono-stratigraphic divisions of the Carboniferous cover rocks are indicated.

The Ribblesdale fold belt rocks within about 1 km of the surface on a line immediately north of the Sykes and Catlow anticlines, with an abrupt thickening of ?late D e v o n i a n - D i n a n t i a n strata to the south (J. D. Cornwell, pers. comm.). In early Brigantian strata in the Catlow anticline the author has identified olistoliths of Asbian reef limestone, These are analogous to those identified by D. J. C. Mundy (pers. comm.) immediately south of the Craven reef belt, east of Settle (Fig. 2), and indicate proximal derivation from a Dinantian shelf margin to the north. The south-eastern limit of the basin is taken at the Pendle monocline (see Miller & Grayson 1982) (Fig. 2), a flexure interpreted here as being the result of the late Variscan reversal of movement on a basin margin hinge line. South of this monocline, Lower Palaeozoic rocks were proved under a thin (490 m) Dinantian sequence in the Holme Chapel borehole. The general stratigraphy of the ?late Devonian-Carboniferous strata of the central part of the Craven basin is illustrated in the crosssection of Fig. 2.

Tournaisian Tournaisian strata are absent on the Askrigg Block north of the North Craven fault, but crop out in the more deeply eroded anticlines of the fold belt. Their base, however, has not been penetrated within the basin, although 704 m were drilled in the Swinden borehole (Fig. 2) where they consist of mudstone and limestone (Ann. Rep. I.G.S. for 1978). Gravity data in the vicinity of the Swinden bore suggest that strata with a density similar to that of the Swinden sequence may be at least 3 km thick, resting on Lower Palaeozoic rocks (J. D. Cornwell, pers. comm.). South of the Pendle monocline a much thinner sequence of Tournaisian rocks was proved in the Boulsworth borehole (Fig. 2) (Ramsbottom 1974).

Vis~an Vis6an strata are estimated to attain a thickness of about 2 km within the central part of the basin, on the northern limb of the Gisburn anticline. These are mainly mudstones although limestone dominates in the lowest 500 m over much of the area. Limestone also dominates in the much thinner Vis6an sequences flanking the basin, north of the Craven reef belt and in the Holme Chapel and Boulsworth boreholes in the south (Fig. 2).

133

Namurian Namurian strata are about 2 km thick in the Pendle monocline, and about 1400 m in the Holme Chapel borehole. There is no evidence of the original thicknesses of either these or the Westphalian strata over the central part of the basin. However, in the Lancaster Fells and Keasden areas to the west of Settle, Namurian thicknesses are given by Moseley (1954, 1956) as 1200 and 500 m plus respectively.

The Ribblesdale fold belt The distribution of the Ribblesdale folds in the central part of the basin is shown in Fig. 2. The anticlines are mostly between 5 and 10 km long. They are generally single-hinged and concentric at the level of outcrop. Many are asymmetrical with their northern or north-western limbs being the steeper. Some of the anticlines terminate against faults; others are cut by, and are laterally displaced across faults (Earp et al. 1961). The horizontal shortening of Carboniferous strata across the fold belt, estimated on the base of the Vis6an (Fig. 2), is of the order of 10%. Turner (1936) suggested that the folds might represent a posthumous Caledonide range within a Caledonide arc. Westoll (in discussion with Moseley & A h m e d 1967), however, favoured an origin connected with dextral wrenching on the Craven faults. Folding of the Ribblesdale anticlines may have occurred over much of Dinantian-early Namurian time, though spasmodically. A Chadian phase is indicated by the author's mapping in the Slaidburn anticline, where late Chadian Worston Shales rest with angular unconformity on the mid-Chadian Thornton Limestone; on the southern limb of this fold the Thornton Limestone is progressively cut out towards the fold culmination. An Arundian phase is suggested in the Eshton-Hetton anticline, where Arundian Worston Shales with a limestone debris bed at their base rest with irregular unconformity on the Thornton Limestone; the oldest part of the Worston Shale sequence at nearby Swinden (Hudson & Dunnington 1944) is locally unrepresented. Folding of this age was proposed also by Hudson & Mitchell (1937) in the Skipton anticline, where they described the (Arundian) Embsay L i m e s t o n e - - a l s o with limestone debris at its base--resting uncomformably on (?Chadian) Halton Shales-with-Limestone; according the George (1978) perhaps as much as 600 m of Chadian strata were removed as a

134

R. S. Arthurton

result of earth movements pre-dating the Embsay Limestone. Evidence for a late Pendleian phase of folding is widespread. This phase and its related erosive episode are represented on the Askrigg block by an unconformity under the Grassington Grit (Rowell & Scanlon 1957) and in parts of the Craven basin also by an unconformity under the Grassington Grit (Mundy & Arthurton 1980), or its equivalent the Warley Wise Grit (Earp et al. 1961). In the Catlow anticline, the Pendle Grit ( s e n s u Earp et al. 1961) thickens markedly down dip both to north and south, and is strongly involved in the folding; while the overlying Grassington Grit (late P e n d l e i a n Arnsbergian) is only weakly involved and its base is unconformable. The importance of the existence of en d c h e l o n folds in plastic cover rocks as indicators of oblique slip in more rigid basement rocks has been emphasized by Harland ( 1971) and Reading (1980). Thus the possibility of wrenching in the Lower Palaeozoic basement is considered here as an explanation of the various en # c h e l o n patterns of folds and minor faults in the Carboniferous strata of the Craven basin described below.

A 9

..........

........

6o-

,'"

i./

9

I

1

B ~

90

"~ ~

malham

"-,..7%+\ --60

~ #

~-~

15

En bchelon folds related to b a s e m e n t faults

i

--

,

,1 ,, ANTICLINE

I

c The Barnoldswick fault (Fig. 2) is one of several Variscan dextral wrench faults identified in the south central part of the Craven basin (Earp et al. 1961). It cuts three en # c h e l o n Ribblesdale anticlines (Gisburn, Middop-Thornton and Lothersdale) and displaces the axial trace of one of these (Middop-Thornton) by about 800 m. The fault is here interpreted as the manifestation in Carboniferous cover of wrenching on a cross-basin basement fracture. Dextral wrench displacement in this fracture would have led first to the formation of en d c h e l o n folds in the cover (see for comparison Wilcox et al. 1973), then to fracturing of the cover, leading to the formation of a main wrench f a u l t - - t h e Barnoldswick fault--parallel with the basement fracture. The South Craven fault system (Fig. 2) is a complex strike-slip zone which crosses the Craven basin. Wrench displacement on this system was invoked by Hudson & Mitchell (1937) and Hudson & Dunnington (1944). The Eshton, Skipton and Lothersdale anticlines all terminate against various parts of this system. South of Malham there is a group of chevron folds affecting Arundian strata in the vicinity of the fault system; their axial traces lie en ~chel o n and sub-parallel to that of the adjacent Esh-

o.,-,oo~

F

015

020

530

l-

"~0 "

/

/

Subsidiary Faults

Anticlines [

200m

I

I FIG. 3. En gchelon patterns in minor folds and faults within the Ribblesdale fold belt. (Based on unpublished maps of the British Geological Survey.) (A) Subsidiary folds in the axial zone of the Catlow anticline, and a chevron fold zone on the north-western limb of the Slaidburn anticline. (B) Chevron fold group associated with the South Craven fault system near Malham. (C) Subsidiary folds and faults related to a main wrench fault (Skipton Rock fault) in the axial zone of the Skipton anticline.

ton anticline (Fig. 3B). The folds typically have wavelengths of the order of 5 - 2 0 m and amplitudes of similar dimension. The South Craven fault system was probably active during the late Dinantian-early Namurian times. Abrupt thickness variation

T h e R i b b l e s d a l e .fold belt

135

in Arundian sediments at the western end of the i axial zone of this anticline, cutting Tournaisian Eshton anticline (mapped by the author) is and Chadian strata (Fig. 3C). Associated en interpreted as a consequence of contempor- Ochelon chevron folds and high-angle, reverse aneous movement on this fault system, and splay faults have been interpreted as the proPendleian movement is indicated some 2 km ducts of dextral transpression within the axial south of Settle, where gently folded Grassing- zone of the main told, the displacement culton Grit (late Pendleian) rests unconformably minating in the formation of the Skipton Rock on tightly flexured Pendle Grit (sensu Earp et at. fault as a main wrench fault (Arthurton 1983). 1961). The formation of the Skipton Rock fault and Like the main folds associated with the Bar- its associated structures post-dates the main noldswick fault, the chevron folds in the Arunfolding of early Dinantian sediments which dian strata south of Malham are interpreted as a form the core of the anticline. This axial zone manifestation of dextral wrenching on an wrenching is envisaged, as at Catlow, to have underlying basement fracture. However, in been secondary to the basement wrenching view of the evidence of synsedimentary move- which produced the main fold. As suggested by ment on the fault system, it is suggested that the Hudson & Mitchell (1937), the latter is prechevron folds may have formed somewhat sup- sumed to have occurred on the South Craven erficially during late Dinantian or early fault system, against which the main fold termiNamurian times in response to wrench dis- nates. A n o t h e r zone of chevron folds, affecting placement on wrench faults already formed in underlying Dinantian strata. Arundian-Brigantian mudstones and limestones, is present on the north-western limb of the Slaidburn anticline (Fig. 3A). These folds En dchelon subsidiary structures related to m a i n trend N E - S W , parallel to the subsidiary folds folds in the Catlow anticline, and are associated with The Sykes and Catlow anticlines (Fig. 2) N N W - S S E faults. They may be interpreted as form an en Ochelon pair of complex periclines at indicators of a E N E - W S W dextral wrench or near the postulated Dinantian margin of the zone; the faults being antithetic fractures in this basin (see above). These folds are considered to zone. have formed as a result of dextral displacement on a basement wrench zone underlying the Dinantian hinge. The anticlines carry en Oche- En dchelon minor faults in a block-basin transition zone ton sets of subsidiary folds (periclines) on their crests (Moseley 1962), and at Catlow these subThe North and Middle Craven faults define sidiary folds subtend angles of about 25 ~ with an elongate fault block within a zone transithe overall axial trace of the main fold (Fig. tional between the Askrigg block and the Cra3A). In the Sykes anticline the crestal periclines ven basin (Fig. 2). These faults were active durtrend NNE on a NE-trending main fold. ing late Dinantian and early Namurian times at Moseley explained this discordance in terms of least (Hudson 1930, 1944), and are here inter'an E - W regional stress field in conjunction preted as members of a suite of rift faults with some factor producing NE structures, with bounding the basin. The Dinantian and early a resulting sinistral rotational shear pattern'. Namurian strata preserved in this fault block The subsidiary folds are here interpreted as are affected, not by en dchelon folds, but by two the products of secondary dextral wrenching en ~chelon sets of minor faults, mapped in detail along the axial zone of the main fold, post- by D. J. C. Mundy (Fig. 4). dating the formation of the latter. The main and The dominant en Ochelon set consists of faults trending variously N W - S E to N N W - S S E . the subsidiary folds of the Catlow anticline, at the level of outcrop, were formed largely in late These faults have normal vertical separations Dinantian-early Namurian times. Folding there exceptionally up to 90 m, either to the NE or was particularly intense during the Pendleian, the SW, the fault slices forming stepped zones, the Grassington Grit being only weakly affected horsts and grabens. Their traces form a pattern and resting with marked unconformity on older which is interpreted as a product of divergent Pendleian strata (see above). dextral wrenching (sensu Wilcox et al. 1973) on In the Skipton anticline (Fig. 2) a zone of en the North and Middle Craven faults, the fault Ochelon chevron folds is present in the axial zone itself having been one of transtension zone and southern limb (Hudson & Mitchell (sensu Harland 1971). Dextral movement along 1937; Arthurton 1983). A high-angle reverse the Craven faults was claimed by Westoll (in fault (Skipton Rock fault) extends along the discussion with Moseley & Ahmed 1967), the

136

R . S. A r t h u r t o n

/

Askrigg

~

Settle

~

,

~

.....

~ ~VEN

I

[

"

FAULT

Craven

...J////~J /

....

1Malharn

I

.

Basm

I

FIG. 4. En Ochelon minor faults in the North-Middle Craven fault block, including a weak NE-SW set and dominant NW-SE to NNW-SSE set. (Based on unpublished maps of the British Geological Survey, mostly by D. J. C. Mundy.) South Craven fault forming a splay fault in relation to such movement. This en ~ c h e l o n fault set formed during late Brigantian times, accompanying substantial movements on the Middle Craven fault (D. J. C. Mundy, pets. comm.). Late Brigantian-early Pendleian mudstones (Bowland Shales) rest with angular unconformity on Asbian-mid-Brigantian limestones. By far the weaker of these two sets consists of faults throwing up to only a few metres and trending N E - S W . Minor faults of this trend are prominent farther north on the Askrigg block (Fig. 4), and indicate a phase of sinistral movement on the North Craven fault, as advocated by Wager (1931) and Moseley & A h m e d (1967). The age of any such phase is unknown.

Discussion A qualitative synthesis of the wrench displacements on the major cross-basin and peripheral faults (both observed and postulated) is shown in Fig. 5. The displacements are those deduced from the observations of the en # c h e l o n patterns of folds and faults described above. With the exception of the North Craven fault, none of the major faults provides evidence of any reversal in the sense of displacement through its wrench history. Stratigraphic evidence indicates that the principal wrench features of the Ribblesdale told belt were imparted, o r at least initiated, during the interval late Chadian-Pendleian. It is argued that, just as the pattern of rifting which produced the Craven basin in ?late D e v o n i a n - D i n a n t i a n times (Leeder 1982) was predetermined by existing fractures in the lower Palaeozoic basement, so the pattern of Dinantian-early Namurian wrenching within the, regional 'distributed zone of shear' (Dewey 1982)

[

"/ ,O<m

y

,

7

q

FIG. 5. A qualitative synthesis of primary wrench displacements on major cross-basin and peripheral faults (large half-arrows), and displacements on secondary wrench zones (small half-arrows). The relative displacement of the Askrigg and Rossendale/Pennine blocks is indicated by bold arrows. was also determined by those fractures in the basement. Thus the observed en Ochelon patterns of folds and faults in Dinantian and Namurian cover may be viewed as the products of displacement on local wrench zones within a general dextral shear regime. Individual en # c h e l o n trends are specific to the wrench zone which produced them. Where wrench zones are of a secondary nature, for example those formed in the axial zones of main folds, the en d c h e l o n trend contrasts with the trend of the main fold. The en 6 c h e l o n fold trends produced in both primary and secondary transpressive zones depend on the degree of convergence on the

The Ribblesdale fold belt w r e n c h zone; H a r l a n d (1971) states that in a transpressive situation the folds form p e r p e n d icular to the m a x i m u m principal stress on the translation direction. T h e fold t r e n d m a y be modified by rotation during w r e n c h i n g .

137

(3) S e c o n d a r y w r e n c h zones f o r m e d in the axial zones of some main folds, and disp l a c e m e n t on these p r o d u c e d s e c o n d a r y en dchelon fold a n d fault sets.

ACKNOWLEDGMENTS: This paper is based largely on the results of the recent re-survey of the Settle 1:50,000 Geological Sheet carried out by the British Geological Survey. The author acknowledges the advice and generous support of D. J. C. Mundy, J. D. Cornwell, E. W. Johnson and L. C. Jones in the development of the ideas expressed in the paper. He alone is responsible, however, for any misconception or misrepresentation. He is particularly indebted to D. J. C. Mundy, F. Moseley and D. Sanderson for improvement of the manuscript. The paper is published with the approval of the Director, British Geological Survey (N.E.R.C.).

Conclusions (1) T h e R i b b l e s d a l e fold belt was established in late C h a d i a n - P e n d l e i a n times in r e s p o n s e to regional dextral shear. (2) T h e m a i n en dchelon fold set is r e g a r d e d as having f o r m e d in C a r b o n i f e r o u s cover as a result of lateral d i s p l a c e m e n t , with varying d e g r e e s of c o n v e r g e n c e , on lower P a l a e o z o i c b a s e m e n t fractures which acted as p r i m a r y w r e n c h zones.

References ARTHURTON, R. S. 1983. The Skipton Rock Fault--an Hercynian wrench fault associated with the Skipton Anticline, northwest England. Geol. J. 18, 105-14. DEWEY, J. F. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. London, 139, 371-412. EARP, J. R., MAGRAW, D., PooLE, E. G., LAND,D. H. & WHITEMAN, A. J. 1961. Geology of the country around Clitheroe and Nelson. Mere. geol. Surv. HMSO, London. ix + 346 pp. GEORGE, T. N. 1978. Eustacy and tectonics: sedimentary rhythms and stratigraphical units in British Dinantian correlation. Proc. Yorks. geol. Soc. 42, 229-62. HARLAND, W. B. 1971. Tectonic transpression in Caledonian Spitzbergen. Geol. Mag. 108, 27-41. HUDSON, R. G. S. 1930. The Carboniferous of the Craven reef belt; the Namurian unconformity at Scaleber, near Settle. Proc. geol. Ass. 41, 29O-322. 1944. A pre-Namurian fault scarp at Malham. Proc. Leeds phil. lit. Soc. 4, 226-32. --& DUNNINGTON,H. V. 1945. The Carboniferous rocks of the Swinden Anticline, Yorkshire. Proc. geol. Ass. 55, 195-215. & M~TCHELL, G. H. 1937. The Carboniferous geology of the Skipton Anticline. Summ. Prog. geol. Surv. Gt Dr. (for 1935), Pt. II, 1-45. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles--Caledonide inheritance versus Hercynian plate margin processes. J. geol. Soc. London, 139, 479-91. MILLER, J. & GRAYSON, R. F. 1982. The regional context of Waulsortian facies in northern England. In: BOLTON, K., LANE, R. H. & LE MONE, D. V. (eds) Symposium on the Paleoenvironmen-

-

tal Setting and D&tribution o f the Waulsortian Facies, 17-33. E1 Paso Geological Society and The University of Texas at E1 Paso. MOSELEY, F. 1954. The Namurian of the Lancaster Fells. Q. Jl geol. Soc. Lond. 109, 423-54. 1956. The geology of the Keasden area, west of Settle, Yorkshire. Proc. Yorks. geol. Soc. 30, 331-52. 1962. The structure of the south-western part of the Sykes Anticline, Bowland, West Yorkshire. Proc. Yorks. geol. Soc. 33, 287-314. & AHMED, S. M. 1967. Carboniferous joints in the north of England and their relation to earlier and later structures. Proc. Yorks. geol. Soc. 36, 61-90. MUNDY, D. J. C. & ARTHURTON, R. S. 1980. Report of field meeting to Settle and Flasby. Proc. Yorks. geol. Soc. 43, 32-6. PHmLIPS, J. 1836. Illustrations o f the Geology o f Yorkshire, Part II. The Mountain Limestone District. Murray, London. xx + 253 pp. RAMSBOTTOM, W. H. C. 1974. Dinantian. In: RAYNER, D. H. & HEM1NGWAY,J. E. (eds) The Geology and Mineral Resources o f Yorkshire. Yorkshire Geological Society. READING, H. G. 1980. Characteristics and recognition of strike-slip fault systems. In: BALLANCE, P. F. & READING, H. G. (eds) Sedimentation in Oblique-Slip Mobile Zones. Spec. Publs int. Ass. Sediment. 4, 7-26. Blackwell Scientific Publications, Oxford. ROWELL, A. J. 8~; SCANLON, J. E. 1957. The relation between the Yoredale Series and the Millstone Grit on the Askrigg Block. Proc. Yorks. geol. Soc. 31, 79-90. TURNER, J. S. 1936. The structural significance of the Rossendale Anticline. Trans. Leeds' geol. Ass. 5, 157-60. -

-

-

-

138

R. S. A r t h u r t o n

WAGER,L. R. 1931. Jointing in the Great Scar Limestone of Craven and its relation to the tectonics of the area. Q. Jl geol. Soc. Lond. 87, 392-424.

Wmcox, R. E., HARDING,T. P. & SEELY, D. R. 1973. Basic wrench tectonics. Bull. Am. Ass. Petrol. Geol. 57, 74-96.

RUSSELL SCOTT ARTHURTON,British Geological Survey, Keyworth, Nottingham NG12 5GG, England.

Variscan tectonics of the Alston block, northern England M. F. Critchley SUMMARY: Several 'blocks' can be recognized in the Carboniferous of northern England. These acted as rigid units and were only affected by the last phases of Variscan deformation. The Alston block provides an example of the deformational history of block regions during the Variscan orogeny. A pre-Carboniferous granite provided resistance to deformation of the block until late Carboniferous times, and fracturing of the Carboniferous sediments above cupolas of the granite is less intense than elsewhere. Away from the cupolas, where a thick succession of Lower Palaeozoic rocks underlie Carboniferous strata, there is a more random and denser distribution of faults and vein fractures. There were three tectonic phases to the structural evolution of the Alston block during the Variscan orogeny. The first phase, occurring prior to the intrusion of the Stephanian Whin sill, resulted in local folding of the block sediments and the formation of early shear joints under E-W compression. The second tectonic phase was a period of uplift of the Alston block, and resulted in a reticular pattern of ENE and NNW normal faults and extension joints. The third phase was characterized by renewed E - W compression. This resulted in the formation of ESE sinistral strike-slip faults, and horizontal reactivations along existing fractures. During Lower Carboniferous sedimentation, northern England consisted of a series of 'blocks' and 'basins' (Fig. 1, insert). The block areas generally have a core of late Caledonian granite, whose low density resulted in the less rapid subsidence of the blocks as compared with the basins. Carboniferous sediments across the whole of northern England show a cyclic nature. On the blocks a relative shallow water sequence of limestones, sandstones and shales were deposited. Rapid subsidence in the basins accompanied the deposition of thick impure and muddy limestones during the Dinantian. By late Namurian and early Westphalian times, the distinction between basin and block sedimentation became less pronounced. During this time a general 'sag' developed over northern England from the Southern Uplands to the Derbyshire massif (Leeder 1982). In the late Stephanian period, the commencement of uplift was accompanied by the intrusion of the basaltic Whin sill and associated dykes. Vein and replacement mineralization occurs in the Carboniferous cover rocks of the Alston block. The results of recent structural studies of the vein fractures and their spatial distributions are described below.

structures, now downthrowing away from the block (Dunham 1933). The Stublick-Ninety Fathom fault line to the north, and the Lunedale fault to the south, acted as contemporary hinge lines during Carboniferous sedimentation. D u n h a m (1948) gives stratigraphic and structural evidence that the movements on the northern and southern boundary faults were reversed during the first phase of Variscan tectonism. Most of the Pennine fault system, along the western edge of the Alston block, was not active during Carboniferous sedimentation. It was initiated by overthrusting and folding from the west during the early stages of Variscan deformation (Burgess & Holliday 1979). The present downthrow to the west is ascribed to the Tertiary period by Bott (1974).

NW faults on the Alston block Within the Carboniferous sediments of the Alston block there are several major N W striking faults; the principal of which is the Teesdale fault. The majority of N W faults downthrow to the NE, and carry little or no mineralization.

Folding Major structures of the Alston block Boundary faults The northern, western and southern margins of the Alston block are bounded by major fault

Intense folding of the Carboniferous rocks in northern England is restricted to the basinal sediments. On the Alston block the Carboniferous rocks are only locally folded; monodines associated with veins are the most typical fold structures. 139

M. F. Critchley

140

FIG. 1. Simplified vein map of the Northern Pennine orefield, showing major structures and locations referred to in text (based on Dunham 1948). Dotted line represents outer margin of fluorite zone. Insert shows location of Alston block and the major Carboniferous blocks and basins of NE England, together with major faults. BFD: Burtreeford disturbance, LF: Lunedale fault; PF: Pennine fault, SF: Stublick fault, T-G-H F: Thieves-Gyle-Harthorpe fault, 90 Fm F: Ninety Fathom fault.

On a larger scale, the sediments have been warped into a broad and gentle dome: the Teesdale dome ( D u n h a m 1931). Today, the highests points of the Teesdale dome occur around Cross Fell (NY687345). In the Durham coalfield, the Carboniferous sediments tilt gently eastwards at about 2 ~ This tilt to the east is probably related to Tertiary uplift along the Pennine faults. After removing the eastward tilt of Carboniferous sediments on the Alston block, two structural 'highs' are noticeable; both of which overlie the cupolas of the Weardale granite. This association could have formed by uplift of the granite, and flexing of the strata above.

Burtreeford disturbance The most prominent fold structure on the Alston block is the N - S trending Burtreeford disturbance. The disturbance is an east facing monocline, with a vertical displacement of up to 150 m. Locally, the monocline may be overturned and thrusted. The formation of the Burtreeford disturbance pre-dates, or was contemporaneous with, the intrusion of the Whin sill. D u n h a m (1948) interprets the disturbance as a compressional feature, related to the same stresses which produced thrusting along the Pennine fault system and the N W faulted monoclines. There is no evidence to dispute this

interpretation, and folding could have taken place on the Alston block at an early stage of compression. However, the Burtreeford disturbance is clearly a major structural feature, probably extending into the underlying granite. If the granite is a homogeneous body then it is difficult to envisage the large strains associated with the formation of a structure of this size, with the relatively small-scale of folding elsewhere. If the Burtreeford disturbance is purely a compressional feature, then why are there no other 'disturbances' on the Alston block? A n alternative hypothesis is that the Burtreeford disturbance marks the junction between two separate intrusions of the granite; the Tynehead granite to the east and the Weardale granite to the west. There was no relative movement between the two granites during Carboniferous sedimentation. Failure along the junction between the two halves of the granite may have occurred during a period of uplift, and the structure tightened by later E - W compression.

Vein fractures The Northern Pennine orefield occupies the central portion of the Alston block ( D u n h a m 1934, 1948). Vein fractures can generally be classified into one of three groups: E N E veins, N N W 'cross' veins, or E - W to ESE veins.

Tectonics o f the Alston block

ENE and NNW veins The ENE veins may be traced for tens of kilometres along strike, although they are not continuously mineralized. Displacements on the ENE veins are rarely more than a few metres. In the competent beds, such as the limestones and sandstones, the veins are nearly vertical. In the less competent beds, such as shale, the dip is shallower (Forster 1809). This feature, taken together with the concentration of vein fill in the steep parts of the veins, suggests that the ENE vein fractures were produced by normal dip-slip movement. The NN W veins are poorly mineralized, with vertical throws generally greater than the ENE veins. Displacements, however, seem to be confined to a few large fractures. Subsidiary, parallel striking fractures have small displacements and dip towards the main NNW vein fractures. The geometry is consistent with normal faulting. The location of the NNW vein fractures may have partly been controlled by the preexisting NNW to N W monoclines and folds of this orientation. The offset relationships between the ENE and NNW veins are highly variable, and it would appear that the two fracture sets developed contemporaneously (Wallace 1861). Following the formation of the ENE and NNW vein fractures by normal faulting, both vein sets were subjected to a second phase of movement. This second phase of movement was strike-slip in nature and occurred before the mineralization; as evidenced by horizontal slickensides covered by undisturbed vein minerals (Critchley 1981). The NNW vein fractures generally moved sinistrally, and the E N E dextrally; although there are exceptions to this rule. Nentsberry mine (NY779466) and the Nenthead mines (NY780430) contain good examples of vein-fracture reactivation. The geometry of the veins showing strike-slip movement suggests reactivation under E - W compression.

ESE veins The ESE veins are the most complex of the three vein sets. Lateral branching is particularly common in these veins, but the refractive effect of normal faults is less evident. The oreshoots show vertical continuity, but restricted strike extension. Dunham et al. (1965) note that oreshoots in the Red Vein of Rookhope (NY930430) were confined to E - W sections of the vein. This strike control was inferred to be due to sinistral strike-slip movement along the vein. The oreshoots on the ESE veins are often

141

thickened by structural overlap, due to the juxtaposition of ore bodies by strike-slip movement. In other cases, the two halves of an oreshoot have been offset by strike-slip movement. Greenwood & Smith (1977) cite the sinistral displacement by 27 m of a transgression in the Whin sill at Blackdene mine (NY868390) as an example of pre-mineralization strikeslip movement. In the Whiteheaps mine (NY955464), the Sike Head oreshoot is developed in the overlap region between two en-echelon sinistral shears (Critchley 1981).

Jointing Jointing has been studied in the Carboniferous rocks at many locations across the Alston block (Critchley 1981). Space prevents a detailed discussion of the joint patterns, but a few relevant points can be made. In Teesdale early dextral and sinistral shear joints, orientated 050 ~ and 120 ~ respectively suggest formation under E - W compression (Critchley in prep.). The dominant joints are orientated parallel to the ENE and NNW veins. These are extension joints, and may have formed during uplift and stress release following the formation of the ENE and NNW vein fractures. In several localities the extension joints were rotated during the strike-slip reactivation of the E N E and NNW veins. Both the early shear joints and the extension joints formed before mineralization; both joint types contain vein minerals adjacent to the vein fractures.

Vein structures at Blackdene m i n e , Weardale In many of the recently operating fluorspar mines, the ESE veins show marked changes in structure as they are followed upwards. In the Blackdene mine, Weardale, Co. Durham, the Slitt Vein in the Whin sill is a major continuous, E - W structure. About 80 m higher, in the Six Fathom Hazle (a sandstone), the vein has split up into several en-echelon sections. Each section is about 100 m long, and strikes at approximately 80 ~ (Fig. 2b). The arrangement of the vein in the Six Fathom Hazle and the Whin sill is comparable to the structures generated in clay model experiments of strike-slip faults. Wilcox, Harding & Seely (1973) describe the formation of Riedel shears above a sinistral strike-slip fault; Fig. 2(a) shows the Riedel shears developed during their clay model experiment. There is a close

M. F. Critchley

142

tions using trend surface analysis. Their results show linear zones across the Alston block within which the strikes of the veins have been systematically rotated. These linear zones were suggested by Carter & Moore to result from strike-slip m o v e m e n t along basement wrench faults. Plotting the position of their zones show good correlations with known major structures of the Alston block (Fig. 3).

10cm

o~

iiiit

~ "

k.._

100 m

Veininlhe

I Whir~Sill

_

- -

I , i

M a i n vein~,on the 4 O f t level

FIG. 2. Comparision of Riedel shears and the Slitt Vein, Blackdene mine, Weardale. (a) Synthetic Riedel shears (R) and antithetic Riedel shears (R') generated during clay model experiment (after Wilcox et al. 1973). (b) Veins in Whin sill and on the 40 ft level at Blackdene mine (mine grid shown). For discussion see text. similarity between the synthetic Riedel shears in the clay model and the en-echelon veins in the Six Fathom Hazle at Blackdene mine. This comparision would suggest that sinistral movement on the Slitt Vein in, and below, the Whin sill resulted in the formation of Riedel shears in the higher strata. Reactivation of pre-existing joints could explain en-echelon patterns of this type. However, the joints in the Six Fathom Hazle strike at about 70~ that is 10 ~ north of the strike of the veins. The geometry of the fracture pattern along the Slitt Vein suggests failure under the conditions of E - W compression. Veins at a high angle to the main Slitt Vein, which could be interpreted as antithetic Riedel shears, are absent at Blackdene mine. However, the E N E veins in the mine area show evidence of a dextral c o m p o n e n t of movement. This dextral m o v e m e n t could have taken place during the period of compression associated with the formation of the Slitt Vein by sinistral strike-slip movement.

Fracture

trace

,

,

, '

,

,

i

analysis

The fracture trace analysis described here is primarily concerned with the spatial distribution and variations of veins over the Alston block and surrounding area.

Trend surface analysis Carter & Moore (1978) have investigated the variations in vein strike from their m e a n direc-

60

SF

,6"

.,.~_~--~-

-

9. .

l~o ~,~''~-

.

-

~J

o~

| ." ~ "',. -"-

,

/t

~

""'iiii"

3o

\k..""/~...M. "'- ~%.2.a / ~.. - - . ~. \ _ ~.-.Y~ ~

"20

~o

70~

~ '~'"*

,.:',,~'Bp'-

lOKm _^ 9,0

NY~_NZ

oo

1. o t - . . . . , . ~ : ~ _2_p

/Outer margin / Whin dykes ....-Major faults fluorite zone . . . Lineaments (after Carter & Moore 1978)

FIG. 3. Correlation of major structures and lineaments of the Alston block. The outline of the fluorite zone approximately coincides with the underlying Weardale granite. Major faults and Whin dykes, along the northern and southern margins of the block are E-W. To the east of the centre of the block, both dykes and major faults appear to be controlled by the ENE lineaments of Carter & Moore (1978). BD: Brandon dyke, BFD: Burtreeford disturbance, HD: Hett dyke, IPF: Inner Pennine fault, LF: Lunedale fault, SF: Stublick fault, 90 Fm F: Ninety Fathom fault.

Vein density and entropy During the present study the statistical distributions of the veins in the central part of the Alston block was studied. Data for this analysis were in the form of digitized vein and fault traces taken from the published 1:50,000 geological sheet. These data were automatically divided into 4 x 4 km grid cells, and various statistics calculated. The length density of fractures for a 1 km area, and the relative entropy proved to be the most useful statistics. Relative entropy is a measure of randomness; a value of 100% indicates perfect randomness, 0% indicates no randomness. Contours of length density and relative entropy are shown in Fig.

Tectonics o f the Alston block

143

FIG. 4. Vein-fracture trace statistics for central part of the Alston block. (a) Contours of vein-fracture density (units are km of veins per km). (b) Contours of relative entropy (in per cent). Shaded areas show positions of cupolas of underlying granite; TG: Tynehead granite, WG: Weardale granite.

4(a,b). The length densities, expressed in kilometres per km square, show two low regions, centred on grid coordinates NY780400 and NY960400. Relative entropy around these two localities is at a minimum; less than 60%. Both areas coincide with the approximate cupolas of the Weardale and Tynehead granites. The analyses suggest that veins over the granite cupolas are present in fewer numbers than in the surrounding sediments and that these veins show a moderately ordered distribution. This latter result is at first sight unexpected; typically the fractures above a granite cupola show radial fracture patterns due to up-doming of the sediments (Norman 1976). The granite in the Northern Pennines pre-dates the Carboniferous sediments, and its intrusion could not therefore have up-domed the Carboniferous sediments. Assuming that the Carboniferous rocks were deposited horizontally, then doming of the sediments could only occur by uplift of the granite basement. The association of low fracture length density and relative entropy may be explained by the rigidity provided by the underlying granite. Where the Carboniferous sediments are underlain by Lower Palaeozoic rocks, fracturing would have taken place more easily due to the lower competence of the pre-Carboniferous strata and the

presence of existing Caledonian structures. The more random distribution in areas away from the granite cupolas may partly be explained by failure along pre-existing structures in the Lower Palaeozoic rocks. In the Teesdale inlier (NY850290), Ordovician mudstones of the Skiddaw Group crop out beneath the Carboniferous cover-rocks. The main Caledonian cleavage in the Teesdale inlier is E - W , with a second cleavage N - S (Burgess & Holliday 1979). The areas having a high entropy value for Variscan fractures show an unusual number of E - W and N ' S fractures, which, together with the typical vein directions, gives an overall random pattern. These unusual E - W and N - S fractures could have been controlled by the existing Caledonian cleavage of the basement.

Structural history of northern England during the Upper Palaeozoic Carboniferous sedimentation The interpretation of the results of the trend surface analysis of Carter & Moore (1978) suggests that major basement faults were present

144

M. F. Critchley

FIG. 5. Evolution and effect of basement faults on the Alston block. (a) Annealing of basement faults by the intrusion of the Weardale granite. (b) Carboniferous sedimentation controlled by basement faults and granite. (c) Whin dykes controlled by basement faults and margin of the Alston block. (d) Rejuvenation of basement faults.

before the Carboniferous. The Alston block is centred around a late Caledonian granite, and bound in part by pre-Carboniferous ENE and N W structures (Fig. 5a). The block, by virtue of its low density granitic core, underwent less subsidence than the surrounding basins, and the basement faults acted as contemporay hinge lines during Carboniferous sedimentation (Fig. 5b). Late Carboniferous structures

The earliest structures in the Northumberland basin were generated by E - W compression before the intrusion of the Stephanian Whin dykes (Robson 1976). Typical early structures of the Northumberland basin are N-S anticlines, and ENE dextral wrench faults (Jones et al. 1980). The orientation of these structures suggests that the first deformational phase in the basins resulted from E - W compression. On the Alston block, E - W compression also resulted in the first phase of Variscan deformation in late Carboniferous times, causing N-S to NNW trending monoclines across the block, and thrusts along the western margin of the Alston block (Fig. 6a). Early shear joints locally formed in the Carboniferous block sediments. In late Stephanian times the intrusion of the Whin dykes were controlled by the pre-existing ENE set of basement faults (Fig. 5c), or enechelon Riedel shears between fault pairs (Randall 1980). The associated Whin sill was

probably intruded at the Commencement of uplift the Alston block (Jones et al. 1980), after the cessation of Carboniferous sedimentation. The Burtreeford disturbance formed by failure of the underlying granite at the onset of uplift. Lateral expansion probably accompanied uplift, reducing the horizontal stresses. Thus, the vertical stress became the maximum principal stress, and conjugate ENE and NNW normal faults developed on the Alston block. With continued uplift the tensile horizontal stresses became greater than the tensile strength of the rocks and extension joints developed. This was the second phase of the Variscan deformation on the Alston block (Fig. 6b). Early Permian structures

The third tectonic phase of the Variscan was characterized by a period of renewed E - W compression on the Alston block after the formation of the ENE and NNW normal faults. This second period of E - W compression may have occurred during the very late Carboniferous or early Permian. On the Alston block the compression resulted in the formation of the ESE sinistral vein fractures, probably contemporaneously with mineralization (Fig. 6c). Horizontal movements also occurred along the existing ENE and NN W normal vein fractures. In the Cheviot block major dextral NE and sinistral? NW wrench faults were formed in the granitic basement (Robson 1976): the orientations suggesting E - W compression. Similar

Tectonics of the Alston block TECTONIC

I

PHASE

East -west compression

STRESS

~

~

o-~ o"3- Horizontal a-2 - Vertical

1T Uplift

o-~ - Vertical o-3-Alternating from ENE to NNW

STRUCTURES

CONDITIONS

__ ~-

o',

145

Gonjugate early shear joints and NNW folds

~

~" o'3

a2 .~-J

~/~2o~1

/- / / ~ o~3 ~.~' ~ ;;/

Conjugate normal faults and extension joints

\

~.,,cr a

o-3 Trr

9 .~ . ~ S ' g ~

Renewed e a s t west camp . . . . ion

m,o'~-Horizontal

-~o--~:

"'J"

\xxxX-~

~1

Reactivation of normal faults by strike- slip movement. Formation of ESE sinistral wrench veins

cr2 - Vertical 0-3

FIG. 6. Summary of the tectonic evolution of the Alston block during the Variscan orogeny. Directions of principal stress axes are shown, together with the main structures.

structures in the granitic basement of the Alston block would give the arrangement of linear zones of fracture rotation as shown by the trend surface analysis (Fig. 5d).

Conclusions (1) Many of the Carboniferous structures in northern England have been influenced by the basement. A postulated series of E N E and NW basement faults controlled Carboniferous sedimentation and the location of the late Stephanian Whin dykes. On the Alston block the basement faults were annealed by the intrusion of the Weardale granite. They were rejuvinated after the formation of the vein fractures on the block, and moved horizontally, rotating the veins above them. (2) The Weardale granite provided a rigidity to the Alston block until late Carboniferous times. The structural highs of the Teesdale dome on the Alston block coincide with the cupolas of the Weardale granite. The Burtreeford disturbance was formed by differential movement of the two halves of the Weardale granite. Fracturing of the Carboniferous sediments above cupolas of the granite is less intense than elsewhere: this is due to the rigidity of the granite. Away from the cupolas, where a thick succession of Lower Palaeozoic rocks underlie the Carboniferous strata, there is

a more random distribution in the orientations of faults and vein fractures. (3) The vein fractures of the Northern Pennine orefield, on the Alston block fall into three groups. The E N E and NNW vein fractures were initiated as normal faults, prior to mineralization, by shear failure. The E N E and NNW vein fractures were probably formed during the continued uplift of the Alston block, after the intrusion of the Whin sill. The ESE vein fractures were formed as sinistral wrench faults during a period of E - W compression. This compression also produced horizontal movements along the existing E N E and NNW fractures. The ESE vein fractures appear to have originated in the lower parts of the succession, and were propagated upwards by Riedel shearing. (4) There were three tectonic phases to the structural evolution of the Alston block during the Variscan orogeny. The first phase occurred during the late Carboniferous, but prior to the intrusion of the Whin sill. E - W compression produced overfolding along the western margin of the block and the formation of the Burtreeford disturbance. The second tectonic phase was a period of uplift of the Alston block. Initially, the uplift warped the Carboniferous cover rocks into the Teesdale dome. The Whin sill was intruded into the partly domed sediments. Continued uplift caused stratal extension and normal faulting, together with the formation of extension joints. The third phase

146

M. F. Critchley

was characterized by renewed E - W compression. It resulted in the formation of the ESE vein fractures by sinistral strike-slip movement, and horizontal movements along existing faults. Mineralization occurred during this third phase of tectonic activity. ACKNOWLEDGMENTS" I would like to thank Mr J. McM. Moore and Mr D. A. Greenwood for con-

tinued advice and supervision during the period of my work on the Alston block, and Adrian Phillips for discussion and constructive criticism. The British Steel Corporation generously allowed access to the mines under their control. The receipt of a N E R C CASE Studentship with the British Steel Corporation at Imperial College is gratefully acknowledged. Finally, thanks to Elaine Cullen for drafting.

References Boaq', M. H. P. 1974. The geological interpretation of a gravity survey of the English Lake District and the Vale of Eden. J. geol. Soc. London, 130, 309-31. BURGESS, I. C. & HOLLIDAY, D. W. 1979. Geology of the country around Brough-under-Stainmore. Mere. geol. Surv. G.B. 131 pp. CARTER, J. S. 8r MOORE, J. McM. 1978. Some major lineaments in the Northern Pennine orefield. 7rans. lnstn rain. Metall. 87, B90-3. CRITCHLEY, M. F. 1981. Structure and formation mechanisms of vein-fracture systems in the north Pennine orefield. Unpublished Ph.D. Thesis, University of London. DUNHAM, K. C. 1931. Mineral deposits of the North Pennines. Proc. geol. Ass. 36, 107. 1933. Structural features of the Alston Block. Geol. Mag. 69, 241-54. 1934. The genesis of the North Pennine ore deposits. Q. Jl geol. Soc. Lond. 90, 684-717. 1948. Geology of the Northern Pennine Orefield--Vol. 1 Tyne to Stainmore. Mere. geol. Surv. G.B. 357 pp. , DUNHAM, A. C. HODGE, B. L. & JOHNSON, G. A. L. 1965. Granite beneath Visean sediments with mineralisation at Rookhope, northern Pennines. Q. Jl geol. Soc. Lond. 121, 384-417. FORSTER, W. 1809. A treatise upon a section o f the strata from Newcastle on Tyne to Cross Fell, with remarks on mineral veins. Preston & Heaton, Newcastle. 158 pp.

-

-

-

-

GREENWOOD, D. A. & SMITH, F. W. 1977. Fluorspar mining in the Northern Pennines. Trans. Instn min. Metall. 86, B181-90. JONES, J. M., MAGRAW, D., ROBSON, D. A. & SMITH, F. W. 1980. Movements at the end of the Carboniferous period. In: ROBSON, D. A. (ed.) 7he Geology o f north east England. Spec. Publ. nat. Hist. Soc. Northumb. 113 pp. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles--Calendonide inheritance versus Variscan plate margin processes. J. geol. Soc. London, 139, 481-94. NORMAN, J. W. 1976. Photogeological fracture trace analysis as a subsurface exploration technique. Trans. Instn rain. Metall. 85, B52-62. RANDALL, B. A. O. 1980. The Great Whin Sill and associated dykes. In: ROBSON, D. A. (ed.) The Geology o f north east England. Spec. Publ. nat. Hist. Soc. Northumb. 113 pp. ROBSON, D. A. 1976. A guide to the geology of the Cheviot Hills. Trans. nat. Hist. Soc. Northumb. 43, 1. WALLACE, W. 1861. The laws which regulate the deposition o f lead-ore in veins, illustrated by examination o f the geological structure o f the mining district of Alston Moor, Stanford, London. 258 pp. WILCOX, R. E., HARDING, T. P. 8r SEELY, D. R. 1973. Basic wrench tectonics. Bull. Am. Ass. Petrol. Geol. 57, 74-96.

M. F. CRITCHLEY, Department of Geology, Trinity College, Dublin 2, Ireland.

Structural variation across the northern margin of the Variscides in NW Europe David J. Sanderson SUMMARY: Four areas which cross the northern margin of the Variscan fold belt are examined and models of their structural evolution, from basin development to folding etc., are compared. Foreland Britain is characterized by heterogeneous deformation localized in basins, whose development accompanied E-W dextral shear. The Munster basin in Ireland contains steep folds and cleavage. A variety of minor structures indicates an important dextral shear component to the deformation, and basin inversion by transpression is proposed. SW England is dominated by thrust tectonics, including the obduction of the Lizard ophiolite. Deformation was preceded by south to north migration of basin development. Brittany is characterized by major dextral shear zones delimiting areas of variable ductile strain. The structural geometry and history of these areas is interpreted in terms of regional dextral transpression. The common transport direction is to the NNW or NW and oblique to the belt. Variable amounts of thrusting and dextral shear produced the different structural patterns. A case is made for an essentially 'thick-skinned' style of heterogeneous deformation produced in a dextral, oblique collision orogen, which involved microplates and a small back-arc basin. A traverse across the northern margin of the Variscan belt from Ireland, through SW England, to Brittany presents a very varied selection of structural styles and sequences. The pattern and sequence of structures must reflect the boundary conditions and internal rheological variations of the orogen. If we knew the boundary conditions, say from plate tectonics, palimspastic restoration etc., and the rheological properties of rocks under geological conditions we would be in a position to m a k e numerical and scale models of a developing orogen. Clearly geology has not reached that stage! What I wish to examine in this paper is how far we can take the observed pattern of structures and integrate it into simple models of regional deformation. It is my belief that these models might supply the necessary link between observation and interpretation of some of the complex boundary conditions involved in orogeny. Much published data and many regional syntheses exist for the Variscan structure in SW England, south Wales and Brittany. In this paper I will discuss the structure of SW Ireland, and then attempt to integrate the models developed there with the rest of the area shown in Fig. 1.

S W Ireland U p p e r Palaeozoic rocks dominate the surface geology of much of south and central Ireland. In the south, they are highly folded and cleaved, whereas throughout the Midlands they are only

FORELAND

FORELAND

Va.riscan

j'/

~=

~;

I J

Mesozoic cover

lOOkm,

FIG. 1 Map showing the extent of Variscan and pre-Variscan rocks in study area. MB--Munster basin; MV--Midland Valley of Scotland; SWE--SW England; S-NF--Silvermines-Navan fault system; SASZ--South Armorican shear zones. 149

D. J. S a n d e r s o n

150

'

"~

D/N6LE

I

Li

"

i/

'

Head

B2

""

D ~

~

/ 7

I

......

/

I

FIG. 2. Map of Munster basin showing Devonian isopachs (in metres), Variscan volcanism and the Dingle-Dungarvan and Cork-Kenmare lines. AB--line of gravity profile (Fig. 8) and CD--line of geological cross-section (Fig. 4). gently folded, faulted and locally cleaved. Traditionally a line from Dingle Bay to Dungarvan (Fig. 2) is taken as the Variscan front, and considered to separate these two tectonic styles. The precise location and definition of such a 'front' in Ireland has proved problematical (see discussion in Naylor & Sevastopulo 1979 and Naylor et al. 1983); it appears to represent different things, in different places, to different people. I will consider Ireland in two regions. One is the Munster basin (Fig. 2), which mainly lies to the south of the 'front' and in which folding and cleavage are fairly intensly developed. The rest of Ireland represents a foreland region in which deformation is heterogeneous and generally weak, but may locally be as intense as in areas within the Munster basin. Munster basin

The Munster basin represents a thick accumulation of Devonian and Carboniferous rocks. Its northern margin is usually defined by a zone of rapid southward thickening of Upper Devonian sediments along an E - W trending line from Dingle through the Galty and Comeragh Mountains. The earliest sediments are a late Middle Devonian (Clayton & Graham 1974; Russell 1978) and Upper Devonian fluviatile sequence which attains a maximum thickness of 6 - 7 km (Naylor & Jones 1967; Graham 1983). The isopachs and Bouguer anomaly map (Murphy 1960) clearly demonstrate an asymmetrical basin, probably produced as a half-graben with basement faulting along the northern margin,

which corresponds in places to the postulated position of the Variscan front. Graham (1983) describes the fining of the fluvial fades towards the interior of the basin with a widespread fluvial coastal plain facies at the top of the sequence. Thus the Devonian sediments indicate a rapid fluvial infill of a basin initiated in the Middle Devonian. Near the Devonian-Carboniferous boundary shallow marine conditions transgressed northward. The main facies boundary in the Lower Carboniferous occurs along a line from Kenmare to Cork City (Fig. 2), separating shelf carbonates to the north from carbonate-poor facies to the south. Naylor & Sevastopulo (1979) suggest that this line may also be controlled by basement faulting. The maximum thickness of Lower Carboniferous rocks in the basin is about 2 km, representing a less rapid subsidence than in the Upper Devonian. Thus from mid-Devonian to mid-Carboniferous times basin development was the dominant structural style in south Ireland. This requires major subsidence and implies considerable crustal (and lithospheric) thinning (McKenzie 1978). The rate of sediment infill allows some estimate of the amount of crustal stretching (fl) and thinning ( l / t ) in the McKenzie model. If we assume fairly typical lithospheric parameters (lithosphere thickness -- 125 km, crustal thickness 30 km, density of mantle, crust and sediment infill = 3.3, 2.8, 2.5 g cm -3, temperature at top of asthenosphere = 1333 ~ and thermal expansion coefficient = 3.28 x 10-5~ -1 then an initial sediment accumulation of - 3 . 9 km is

The northern margin o f the Variscides

half its original value and an - 8 km rise in the Moho (Fig. 3b). /7 = 2 is considered by Dewey (1982) to produce the onset of volcanicity. Thus the localized volcanicity in Beara (Coe 1966; Coe & Selwood 1963) and along the northern margin of the basin (Penny 1978) support the model and appears to be located along major faults (Matthews, Naylor & Sevastopulo 1983).

E I':__~- T T.- - [ 7 ]-G=H

a

b

151

c

FIG. 3. Diagrammatical representation of crustal thickness evolution of the Munster basin. (a)Initial crust (Early Devonian). (b) After lithospheric thinning and basin development (mid-Carboniferous). fl = 2 in McKenzie model. (c) After deformation, thickening and uplift (end-Carboniferous). = 1.7 in transpression model.

Structures within the Munster basin

predicted for instantaneous stretching of/7 = 2. This would be followed by a further timedependant subsidence producing - 1 . 6 km of sediment in 20 Ma. Thus fl = 2 is consistent with a 5.5 km sequence of l a t e - M i d d l e Devonian to U p p e r Devonian basin sediments. By mid-Carboniferous a further 1.5 km of sediment would be predicted. These sedimentation rates accord well with those observed in the basin and imply a thinning of the crust to about

In this section I will describe the various structural elements, based mainly on mapping on the Beara peninsula in west Cork. I have tried to group structures which I consider representative of the entire west Cork region, but how far east these can be traced I do not yet know.

Folding The rocks of the basin outcrop in a series of major E N E - W S W trending folds, which become more E - W to the east of Cork City.

N

\

,/"

\

I

-I

\~\

I

\

\

1-

\

\

-I

/

\

I BANTRY

S

BINGLE DUNGARVAN

CORK KENPIERE -

BAY

CARBONIFEROUS

GABRIEL

,, , /~l'/~lxl,\//"

/1\/ \ I\I)/\

I z

,, ',l

" .;-'t/,

i,,

,/\1 \'' \

/ "I " '

-

~

\

I

\

t

I"

\

~_

/

FIG. 4. Cross-sections through Munster basin (line CD in Fig. 2), based on surface structure of Naylor (1978). (a) Thin-skinned model, with d6collement at basement-cover interface. (b) Thick-skinned model.

152

D. J. Sanderson

The folds are mainly upright or steeply inclined and are developed in a broad fan across the basin (Fig. 4). Near Killarney, the folds are more overturned, with axial planes locally dipping gently south. Similar broad fans of axial planes are seen in the sections to the east (Cooper et al. 1984). The fold axes plunge both to the WSW and ENE indicating a series of culminations and depressions along individual major fold hinges. On the Beara peninsula, the culminations and depressions of minor folds occur in zones which trend obliquely across the major fold hinges, breaking up the continuity of individual major folds. Fold style varies with lithology, but the typical minor folds are rounded, asymmetrical, and congruous to the major folds. Locally 'structural terraces' or 'step folds' are common (Coe & Selwood 1963), with flat-lying, short limbs sub-normal to cleavage, and steep long limbs at a low angle to cleavage (Fig. 5a). The low competence contrast of the layers within the Devonian strata probably gave rise to slowly amplifying folds with much layer-parallel shortening. Thus cleavage generally penetrates all layers with little refraction and fanning. In such cases

'flattening' of early buckles gives rise to folds with relatively thickened hinges (or flat limbs of step folds). Cleavage A single, pervasive, steep cleavage is developed, which in the argillaceous horizons produces a slaty cleavage and in the siltstones and fine sandstones is a spaced cleavage. Even where slaty, the cleavage in thin section consists of closely spaced solution seams and there is widespread evidence, on all scales, for the operation of pressure solution controlled diffusion creep as the dominant deformation mechanism. Detrital chlorite and quartz grains have sutured boundaries, with the development of fibrous quartz/chlorite/mica pressure fringes. Veining accompanies cleavage formation and evidence for the 'crack-seal' mechanism (Ramsay 1980) is common. The close association of veins, cleavage and folds is ubiquitous. Although cleavage is often more or less parallel to fold axial planes, many examples of nonaxial planar cleavage may be found locally. There seems to be little systematic obliquity of cleavage and axial planes, as is reported from the para-tectonic Caledonides of Ireland (Sand-

Fly. 5. Minor structures in the rocks of the Beara peninsula. (a)Typical 'step' fold style. (b,c) Boudinage. (d,e) Shear zones. (f) Combination of boudinage and shear zones. (g) E-W veins (and faults). (h,j) Extension veins. (k) Kink-bands.

The northern margin o f the Variscides erson et al. 1980). There is some indication on the Beara peninsula of cleavage striking slightly anticlockwise of the major folds, to produce a change in the pitch of the bedding/cleavage intersections across major fold hinges. Early reports of two cleavages and polyphase deformation (Gill 1962) have not been substantiated on a regional scale. Local zones of crenulation cleavage, usually adjacent to faults, or in the complexly strained contact regions around early or syntectonic veins, cannot be correlated on a regional scale. A very interesting feature of the timing of cleavage development was recognized by Coe (1966) and Coe & Selwood (1963) on the Beara peninsula, where minor intrusions of diorite and tuffisite cross-cut folds, but are themselves deformed and cleaved. They attributed this to a two-phase development of cleavage, but, in the absence of any clear petrographic or field evidence for superposition of cleavages, I would prefer to think of this as resulting from syntectonic igneous intrusion during a single, prolonged phase of regional deformation. Very similar relationships are seen between quartz veins and cleavage throughout the region, with early formed veins being folded and later veins cross-cutting the folds and cleavage (Fig. 5h). High pore fluid pressure, and high magma pressure, during deformation would produce these brittleductile transitions. Boudinage Coe (1959) has described boudinage from west Cork. Although layers may be extended in many directions, boudins are mainly developed with long axes pitching steeply in bedding and indicating strike-parallel extension (Fig. 5b). This type of boudinage becomes less frequent to the north and south of Beara, where subvertical extension and flat-lying veins and boudin long axes are more common. Early shear zones Another common manifestation of strikeparallel extension is the development of ductile shear zones in sandstone beds. These usually occur in conjugate sets, with displacement of bedding and deflection of cleavage indicating dextral and sinistral shear (Fig. 5d). En echelon quartz veins (Fig. 5e) indicate similar shear senses. These zones have a close kinematic and genetic relationship to boudinage, especially where quartz extension veins develop (Fig. 5f). Strain indicators Oblate (K --~ 0) shaped reduction spots occur

153

widely, but no systematic measurement is yet available. Work on deformed dessiccation cracks supports the widespread development of flattening strain and preliminary results suggest an increase in strain towards the centre of the basin, the highest strains being recorded on the Beara Peninsula. Using cleavage as an indicator of the XYplane and grain alignment, fossil distortion, pressure fringes etc. to locate the stretching direction (X), the orientation of the strain ellipsoid can be obtained from simple field observation. The XY-plane is generally steep and stretching lineations poorly developed, supporting the view of widespread flattening strain. Where observed, X is generally steeply/moderately inclined, but, locally on the Beara peninsula it is subhorizontal. Vein systems A complex system of veins is developed; most can be grouped as follows: (1) Bedding-parallel slickensides These are relatively rare in the Devonian rocks, where flexural-slip folding is uncommon, but are more common in the Carboniferous strata. Where seen such slickenside veins have fibres at a high angle to fold axes and are often cut and folded by the main cleavage. They represent early slip along bedding accompanying the onset of folding. (2) Early extension veins These often occur as steep, roughly N-S trending veins which are strongly folded within cleavage (Fig. 5j). They are very common on the Beara peninsula, where they are often mineralized (chalcopyrite, siderite etc.). They indicate early, subhorizontal, strike-parallel extension pre-dating much of the cleavage forming strain. These veins form a continually developing set which pass into shear zone extension veins (en echelon arrays) and the 'late extension veins' (see below). (3) Late steep extension veins These form a prominent set of N-S trending veins along acjoints (Fig. 5h), which commonly develop in the sandstones. They may accommodate up to 15% extension along strike and are frequently filled with quartz fibres and pegs. Generally these veins cut across cleavage, but in boudin necks and en echelon arrays some deflection of the cleavage occurs. Thus they are regarded as being 'late tectonic' in age and together with the early extension veins suggest strike-parallel extension throughout cleavage development. (4) Flat extension veins

These are developed

154

D. J. S a n d e r s o n

in steeply dipping sandstone layers and indicate subvertical lengthening of the layer. Whilst they occur on Beara, they become more dominant both to the north and south and largely replace the steep veins in Dingle and Mizen Head. (5) Inclined extension veins These represent structures intermediate between the steep and flat extension veins. They generally dip to the east, a feature which can be attributed to the combination of subvertical extension, vertically directed shear and E - W dextral shear. The last two shear elements effectively produce an oblique shear with the shear direction pitching west and hence the veins have an easterly dip component. (6) E - W veins These commonly occur along faults and longitudinal joints. They are frequently mineralized and at Allihies form the main lodes (Sheridan 1964). They often have pinnate veins and extensional offsets (Fig. 5g) which indicate dextral shear. They typically occur as a stockwork of veins along faults, and early formed pinnate veins are folded in cleavage. They are thus thought to have formed in association with folding and cleavage and, by implication, the associated strike faults must also be early. Faulting Faults are common throughout the basin and form two broad categories; strike faults, which are mostly early in the deformation history, and 'cross' faults (including 'diagonal' or 'oblique' faults) which are generally later and cut folds (Naylor et al. 1983). Strike faults have been mapped by many workers and have been interpreted in a variety of ways. I consider that there are three distinct types of strike fault. Early, pre-tectonic, normal faults control the basin development and involve basement. Faults of this type probably formed the northern margin of the basin and influenced the siting of the Variscan front, either as a structure or a cartographic convenience! These faults may also have been utilized for igneous intrusion as in the Beara-Bantry area (Matthews et al. 1983). Minor, low-angle, listric? thusts, with movement to the north and south, are locally common and have been used to infer major thrusting in the Cork area (Cooper et al., this volume). In the Glandore area, Reilly & Graham (1972) recognize important E - W dextral strike-slip faults controlling mineralization and I have mapped many minor, high-angle faults with dextral movement on the Beara Peninsula. These faults, and those mapped by Capewell

(1957) on the Iveragh peninsula seem to be a common feature of west Cork and Kerry and control many of the E - W vein systems, as at Allihies and Glandore. The 'cross' and 'diagonal' faults are oblique to the strike and fold axes, and generally trend within 45 ~ of north. They are usually obliqueslip faults, with both vertical and horizontal components of displacement. Those trending to the east of north generally have a dextral component, whilst those to the west have a sinistral one (for review see Naylor et al. 1983). Most of these faults cross-cut the folds and strike faults and appear to be late in the structural history. Philcox (1964), however, has suggested that the fold pattern changes across some of these, with the development of folds within 'compartments' defined by contemporaneous cross faults. This type of deformation has also been suggested from the Variscan of south Wales (Owen & Weaver 1983) and SW England (Dubey 1980). Kink bands In many of the finer grained, cleaved lithologies kink bands are common. These clearly post-date cleavage and folding, and indicate a component of strike-parallel shortening (Fig. 5k). Although conjugate kink bands with steep kink planes are occasionally seen, typically only one set of steep, dextral kink bands develops with a N - S trend and at a high angle to cleavage. As Dewey (1966) has pointed out, this set is common throughout the Variscides of Ireland and implies a late sinistral shear along the belt, but one which represents little strain. Models of Variscan deformation within the Munster basin

Dextral transpression model Within the Munster basin folding, cleavage, boudinage, veining and strike faulting are broadly contemporaneous. Any model of deformation must consider all of these structures and also the overall history of the basin. The N - S shortening, folding, local thrusting and the inversion from subsidence (and crustal stretching) to uplift (and crustal shortening) suggest a 'typical' N - S Variscan compression. The existence of strike-parallel extension and the E - W dextral wrench faulting, however, suggest the operation of a horizontal dextral shear component. This would also account for the obliquity and swing in trend of many of the structures (see also Max & Lefort, this volume). These shortening and shear components may be combined in a transpression model (Harland 1971).

The northern margin o f the Variscides Z

.,~x

"-~r FIG. 6. Transpression model; y is angular shear parallel to the boundary (7 = tan qJ) the crustal thickenning, and a-1 the shor-' tening across the zone. Sanderson & Marchini (1984) have examined a transpression model involving a zone subject to a combination of transcurrent shear parallel to its margins and compression across it (Fig. 6). The shortening across the zone (~-1) is compensated by a vertical thickening (~). In effect the model encompasses a spectrum of cases between pure shear (~ -- 0) and simple shear (~ = 1). Sanderson & Marchini demonstrate that such deformation, with ~ > 1 (i.e. involving shortening and thickening of the crust), leads to the following strain features: (a) The finite strain will be oblate (K < 1). (b) Depending on the values of ~ and ~ the X-axis can be either vertical of horizontal. (c) The XY-plane is vertical and makes an angle < 45 ~ to the zone boundary. The precise angle depends on the values of and :~, but will always be less than the corresponding angle for simple shear alone. These features are typical of the Munster basin; but to make a semi-quantitative comparison we need to estimate the parameters ~ and 7. Assuming that transpression operated on a crust (plus sediment) thinned to - 2 0 km during basin development (Fig. 3b) to produce a 'normal' crust of - 3 0 km, then a vertical stretch, = 1.5, is required. Allowing for subsequent uplift and erosion, this may be increased to 1.7. To achieve K = 0 strain and an interchange of X from vertical to horizontal (as on Beara) requires ~ ~- 2. This combination of parameters would produce a horizontal strain ratio of - 4 - 5 , with the XY-plane at - 1 6 ~ to the zone boundary. These values are almost precisely

155

those observed on the Beara peninsula. Thus the model is consistent with the limited data available. Decreasing amounts of shear (r) to the north and south of Beara would give lower strains and a vertical X-axis, as observed. Such regions would have smaller, but still significant, components of dextral shear. The model needs to be tested more rigorously and elsewhere in the basin. Since the model implies crustal thickening within the basin there should be strain discontinuities or zones of complex strain at its margains. These may have reactivated the bounding faults as high-angle reverse faults, as is observed in places adjacent to the DingleDungarvan line (Philcox 1964; Morton 1965; Brennand 1965; Walsh 1968). Uplift might also promote gravity driven, low-angle thrusts (Gill 1962; Cooper et al., ~his volume). A very good analogy of this 'flower structure' pattern (Fig. 4b) is described by Sylvester & Smith (1976) from the Salton Trough of southern California. Figure 7 shows how transpression might result from the oblique closure of a basin, and is based on the simple transpression model of Harland (1971) and Sanderson & Marchini (1984).

Thick or thin skinned tectonics? From the field evidence in west Cork, I have argued for a transpressive closure of the Munster basin in late Carboniferous times. Within any fold belt field evidence is restricted to those rocks at present exposed at the earth's surface. The deeper crustal structure must be interpreted in such a way as to be consistent with this

FIG. 7. Simple transpression resulting from oblique closure of a basin (stippled area). The unshaded square (upper diagram) is transformed to the parallelogram (lower diagram).

D. J. Sanderson

156

surface evidence and any geophysical data available. Two general models of deformation style have been widely applied to fold belts and can be termed 'thin-skinned' and 'thickskinned'. The essential difference lies in the depth to which the near surface structural elements are considered to extend. Figure 4 considers, in somewhat diagrammatic form, extreme cases of the two models applied to a cross-section through west Cork and Kerry (after Naylor 1978). A detailed application of thin-skinned models is given by Cooper et al.

(this volume). In the absence of deep seismic reflection data the two models are difficult to test directly. Various types of data place constraints on both models and I shall examine some of these. One might expect field observation of the faults to provide definitive evidence. Most of the exposed strike faults are steep, but the critical factor is their geometry at depth. Are they listric or do they penetrate deep into the crust? The 'master d6collement' or 'sole thrust' in a thin-skinned model will only reach the surface along the thrust front, and here the displacement will be small and could consist of a series of ramps with steep dips (see Cooper et al., this volume). These 'blind thrusts' can only be proved by seismic reflection data or deep boreholes, neither of which are available. The Munster basin is coincident with a well developed gravity low (Murphy 1960). A - 3 5 mgal, anomaly coincides with the basin, but appears to be superposed on a larger wavelength, regional anomaly of some - 1 0 mgal (Fig. 8). These anomalies can be modelled, using the two-dimensional methods of Talwani, Worzel & Landisrrian (1959), by an asymmetrical basin developed above a region of crust with the Moho depressed by 1-2 km. Thus the Munster basin is interpreted as exist-

30 *

";?,~

rngats 10' ZO i 30 1

Moho

FIo. 8. Bouguer anomaly profile (AB in Fig. 2) across the Munster basin and its interpretation. The anomaly consists of a regional low produced by the depressed Moho and a local low due to the basin. No attempt has been made to model the gravity low at the northern end (left) of the profile.

ing above slightly thickened continental crust. If this is compared with the thinned crust necessary during basin development in the late D e v o n i a n - e a r l y Carboniferous, considerable crustal thickening during Variscan deformation must have occurred. Thus I consider the gravity data, combined with McKenzie's model of basin development, to support a thickskinned model. The amount of crustal thickening, ~ in the transpression model, can be estimated (Fig. 3c). The calculation assumes an initial crustal thickness of 30 km, a 2 km depression of the Moho at present under the basin, a/~-value of 2, and erosion of some 5 km of strata. Thus during deformation a - 2 2 km crust was thickened to - 3 7 km giving a ~ 1.7, and a horizontal shortening of - 4 0 % . Cooper et al. (this volume) argue that horizontal shortening of some 30% in east Cork, as measured from balanced sections, precludes thick-skinned tectonics. Clearly such a value is also consistent with the thick-skinned transpression model proposed here. I conclude that there is much evidence for crustal thickening and a thick-skinned model of deformation. Taken together with the evidence for dextral shear, a transpression model for the closure of the Munster basin is proposed. Some of the crustal thickening may have been achieved by thrusting both to the north and south.

Foreland deformation

Devonian and Carboniferous strata crop out extensively in the British Isles to the north of the Variscan front. I shall discuss a few important features from three areas within this foreland region. In central Ireland (Fig. 9) a thin Upper Palaeozoic shelf sequence, with local basins (e.g. Shannon, Dublin, etc.), was developed on Caledonian basement. Variscan deformation began in the Lower Carboniferous. Critical evidence for this is found at Navan, where much of the faulting and folding pre-dates a Visean unconformity (Brown 1979). Elsewhere E - W to E N E - W S W faults have been demonstrated to be active during the Dinantian, as at Tynagh (Moore 1975) and Silvermines (Taylor & Andrew 1978; Coller, this volume). This implies that deformation was taking place during the development of the Visean and Namurian basins. The deformation in Central Ireland occurs mainly in zones with two main trends (Coller, this volume). E - W to E N E - W S W trending

The northern margin of the Variscides

I

1ookm

~/~.~\ ~

'

~

FiG. 9. Location map of the structures and areas discussed in the foreland region of the British Isles. AI--Alston block; As--Askrigg block; D--Dublin; F--Fergus shear zone; M--Mendips; MV--Midland Valley of Scotland; N--Navan fault; RFB--Ribblesdale fold belt; S--Silvermines fault; T--Tynagh; VF--Variscan front. The trends of folds and dykes are shown diagrammatically. dextral zones are usually associated with major faults, such as the Navan and Silvermines faults (Fig. 9). N - S to N N E - S S W trending sinistral zones are usually more diffuse and represent smaller displacements. These zones, especially the ENE trend, reflect underlying basement control, e.g. the Navan-Silvermines fault system follows the trace of the Iapetus suture (Phillips, Stillman & Murphy 1976). A somewhat similar picture is seen in north England, where deformation is mainly confined to zones or basins surrounding more stable blocks (e.g. Alston and Askrigg blocks). The Ribblesdale fold belt (Fig. 9) represents an E - W trending basin which Arthurton (this volume) has demonstrated was undergoing dextral shear throughout much of the Carboniferous, at least from mid-Dinantian to early Namurian. Late Caledonian granites underly the blocks and probably helped resist stretching and prevented isostatic subsidence. They also sealed basement faults, reducing subsequent deformation (Critchley, this volume). One of the most clearly defined zones of Variscan deformation in the foreland is in the Midland Valley of Scotland (Fig. 9). Leeder (1982) envisaged the early evolution of this area to involve subsidence of a remnant Caledonian

157

fore-arc basin, but Bluck (1980) has emphasized its strike-slip origin as a pull-apart basin. The total Devonian-Carboniferous thickness in the basin exceeds 12 kin, and Dewey (1982) suggests this implies /~ = 6 in the McKenzie model. This is an extremely large value, which is consistent with the widespread volcanicity in the area. Such a large stretch cannot be continued west into Ireland, and its localization in a strike-slip pull-apart basin, as envisaged by Bluck, seems most likely. Closure of the basin was mainly in the Upper Carboniferous, but with minor folding in the Middle Devonian. This need not have produced much crustal thickening (~ < 1.5) and may have occurred mainly by dextral simple shear. The obliquity of folds, normal faults and Permo-Carboniferous dykes to the bounding faults (Fig. 9) corresponds closely to that seen in simple wrench tectonic models (e.g. Wilcox, Harding & Seely 1973). These few examples of deformation in the Variscan foreland serve to illustrate two important general features: (a) The deformation is often localized in approximately E - W trending zones of dextral transpression. These zones reflect basement faults. (b) The timing of both basin development and deformation varies. Deformation occurred, at least locally, throughout much of the Carboniferous and was contemporaneous with basin development elsewhere. There may have been earlier 'pulses' of deformation, such as the Middle Devonian folding in the Midland Valley.

SW England The structure of SW England has been reviewed by a number of authors (Sanderson & Dearman 1973; Matthews 1977; Shackleton, Ries & Coward 1982; Hobson & Sanderson 1983). The structural evolution involves a fairly complex sequence of basin evolution and tectonism. Basin development took place throughout much of the Devonian, with a northward overall progression of deposition (Matthews 1977). Gramscatho basin and Lizard obduction

This basin (Fig. 11) was developed in late Silurian-early Devonian times and was accumulating flysch sediment throughout the Devonian. The Lizard complex lies on its

158

D . J. S a n d e r s o n

southern margin, and there is a growing acceptance of this as an early Variscan ophiolite. The precise age of the Lizard complex is still controversial (see discussion in Floyd 1983), but an early Variscan age (say 375-400 Ma) seems most likely (Halliday & Mitchell 1977; Styles & Rubdle 1981; Davies 1984). A Lower Devonian age for the Lizard rocks would overlap with that of the sediments in the Gramscatho basin (Sadler 1973). The relative position of the two in Lower Devonian times is not known. However, the presence of ophiolite m61ange (Barnes, Andrews & Badham 1979) and the widespread tholeiitic volcanism (Floyd 1983) in the Gramscatho basin would be consistent with their proximity. I follow Badham (1982) in considering the Lizard ophiolite to represent a small, isolated piece of oceanic crust developed within, or very close to the Gramscatho basin (Fig. 10a). Floyd (1982) notes that the tholeiitic basalts of SW Cornwall have geochemical characteristics of enriched midocean ridge basalts or those from back-arc basins. The latter environment is preferred on general geological grounds. The crustal thinning necessary to produce the basin and oceanic crust may have arisen as a local pull-apart associated with dextral strikeN Devon basin

TB ~

Tr e v o n e

-

~

slip faulting (Badham 1982). In addition to providing a convenient explanation for the localized nature of the Lizard rocks, the N W - S E trend of the sheeted basalt dykes implies N E - S W extension or spreading, which is consistent with an E - W dextral shear. The obduction of the Lizard complex involved NNW to NW directed thrusting over the Gramscatho basin, and was accompanied by deformation of the Devonian sediments (Rattey & Sanderson 1982, 1984). Two phases of deformation (Da and D2) were developed during obduction, with D1 being accompanied by low temperature/high pressure regional metamorphism of pumpellyite-actinolite facies (Barnes & Andrews 1981). The tectonic style and metamorphism of this flysch filled basin is similar to that of a small accretionary prism produced during limited subduction of a marginal sea basin. The timing of this deformation is more problematical. It is tempting to use the Upper Devonian m61ange as an indicator of the start of obduction (Barnes et al. 1979). This is implicit in the model of Shackleton et al. (1982), but they attribute a Lower Devonian age to the m61ange. The c. 365 Ma K-Ar ages from slates in SW Cornwall (Dodson & Rex 1971) are taken as

Culm basin

" ~

f A '~ j~.o

--- ~

,,..

6ramsca fho basin

S

Lizard oMucfion

..................... ""

"-"

"

"

1 ....

C

I

,.,.,.'.,4

FIG. 10. Sketch sections showing the evolution of SW England. (a)Early Devonian. (b)Late Devonian. (c) 'Mid'-Carboniferous. (d) End-Carboniferous. AB--alkali basalt, TB--tholeiite and vertical striping indicates oceanic crust.

159

The northern margin o f the Variscides

Fold axes -~

Sfrefching

/~

Sense of Thrusting v

Alkali

9

Tholeiific

dtrecfions

B ~

CULM

BASIN

basal~s basatfs

Oranife "~v

TREVONE

BASIN v

GRAMSCATHO 9

v

v

v

vI

[ vv

' ~-I ""-- ~

back folding fhzu.._sfing

BASIN Start

50

L i z a r d \^1

\

km

~

FiG. 11. Map of places, volcanicity and structures discussed in SW England. B--Bude; E--Exeter; M--Mevigissey; P--Plymouth; PP--Perranporth; Pn--Pentire Point; Pz--Polzeath; T--Tintagel; TB--Torbay. The diagonal shading indicates the zone of moderately or steeply inclined folds at southern margin of the Trevone basin. support for an Upper Devonian age of deformation. There is little stratigraphic constraint as the youngest rocks in the basin are also Upper Devonian. Trevone basin

In this basin Lower Devonian shallow-water sediments (Dartmouth slates) are overlain by a thick sequence of Middle Devonian to Upper Carboniferous basinal sediments. To the east, around Plymouth and Torbay, thick linestones are present. Rapid facies changes characterize the basin (House 1971), with localized condensed, nodular limestones being laterally equivalent to deeper, basinal deposits. This is the equivalent of the 'Schwellen und Becken' development in the Rheinish Schiefergebirge of Germany (Goldring 1962). A belt of Upper Devonian alkali basalts, mainly submarine, extends along the northern margins of the basin from Pentire Point to Torbay (Fig. 10b). Near-shore facies in the Lower Devonian suggest development of the basin on continental crust. The alkali basalts have an intra-plate geochemistry (Floyd 1982), typical of rifted continental crust. Stratigraphic thickness is dif-

ficult to estimate, due to later deformation, but is probably c. 5 kin, and represents a crustal stretching of fl = 1.5-2. The basin was deformed sometime after the Lower Carboniferous, the exact timing is uncertain (Hobson & Sanderson 1983). In Fig. 10(c) a mid-Carboniferous tectonism is shown, but this may have taken place somewhat later. The deformation produced north facing folds which correspond to the early primary deformation recognized by Roberts & Sanderson (1971) in the Polzeath area. These folds were accompanied by thrusting in south Devon (Coward & McClay 1983; Shackleton et al. 1982). Culm basin

This basin was initiated by crustal stretching in the Upper Devonian and received restricted sediment up to the Lower Carboniferous, producing a thin, moderately deep-water sequence accompanied by alkali basalt volcanism along the southern margin from Tintagel to Exeter (Fig. 10c). The main basin infilling was in the Upper Carboniferous with something in excess of 2 km of sediment being deposited. The basin was deformed in late Carbonifer-

160

D. J. S a n d e r s o n

ous times to produce the Culm synclinorium, a complex structure involving a broad fan of facing directions (Fig. 10d). South facing folds are developed along the southern margin and may be attributed to southerly directed thrusting (Sanderson 1979; Rattey & Sanderson 1982). These folds extend southward to Polzeath and, hence, effect the Trevone basin. The central part of the Culm synclinorium is characterized by upward facing folds which farther north become overturned to the north. North Devon basin

In north Devon a thick (0.6 km) sequence of Devonian and Carboniferous shallow marine and continental sediments, with no volcanics, is exposed. This basin was developed on continental crust, detected at a depth of >6 km by seismic refraction (Brooks, Mechie & Llawellyn 1983). Dewey (1982) interprets sediment thickness to indicate a crustal stretch of fl = 2, although I favour a slightly lower value. The basin was deformed, producing upright and north facing folds, at the end of the Carboniferous, some 100 Ma after initiation. This allowed fairly complete thermal subsidence (McKenzie 1978) and a Moho rise of - 6 km is envisaged. Discussion of deformation in SW England

Figure 10 brings out the broad pattern of a northward (towards continent) progression of basin development, volcanism and deformation, although detailed relationships are more complex and poorly understood. The facing of the early folds is to the north or NNW, except at the southern margin of the Culm basin. The direction of thrusting can be deduced from the stretching lineation in the rocks and from differential movement within the thrust sheets (Rattey & Sanderson 1982) and is to the NNW throughout much of the region. At the southern margin of the Culm synclinorium this movement is reversed (Sanderson 1979; Rattey & Sanderson 1982) in what Shackleton et al. (1982) regard as a large 'pop-up' structure. This direction of transport is oblique to the trend of the Variscan front in Britain (Fig. 1), and is consistent with a component of dextral strike-slip along the belt. There are very few other indications of dextral transpression. In north Cornwall and south Devon there are some late-tectonic vein arrays which show dextral shear parallel to the strike of the fold axial planes. The strain involved i s very small and their regional significance debatable. The N E - S W trend of fold axes in SW Corn-

wall (Sanderson & Dearman 1973; Rattey 1979) is oblique to the Variscan belt as a whole and to the northern margin of the Gramscatho basin and the Lizard thrust (Fig. 11). Folds developed anticlockwise of bounding shears are characteristic of dextral shear. This N E - S W trend of minor folds persists northwards into the Trevone basin and eastwards to the Plymouth area (Hobson 1976a,b). Accompanying this obliquity of minor folds is a steepening of their axial planes (Fig. 11). Sanderson (1971) and Hobson (1976a) attribute this to refolding by later phases, which Shackleton et al. (1982) and Coward & McClay (1983) term backfolding and attribute to SSE directed backthrusting. An alternative explanation is that strike-parallel transcurrent movements produce the steepening of cleavage. Oblate strain characterizes this zone and locally the stretching direction is subhorizontal and parallel to strike (Coward & McClay 1983). Both these features are predicted by including a shear component in a transpression-type model. I do not wish to overemphasize this dextral shear component, since the dominant deformation style in SW England is clearly one of thrusting to the NNW and SSE. Shackleton et al. (1982) interpret the thrust tectonics in terms of a thin-skinned model, with a basal d6collement which dips gently south and underlies the entire area, including much of south Wales. This d6collement is thought to form the base of the granites at a depth of some 12 km under Cornwall. The backthrusting and 'pop-up' structure is seen as developing from this master d6collement. Granite related structures in SW Cornwall post-date Lizard obduction and the D a and D 2 fold phases (Rattey & Sanderson 1984). Elsewhere the granites are post-D~ as at Dartmoor (Dearman 1959). Thus much of the thrust related deformation is pre-granite in age. Clearly granites could not have originated within the thin-skinned nappe (even if 12 km thick). Shackleton et al. (1982) argue that the granites cannot have originated from the crust directly below the thrust pile, due to it not being sufficiently thick, and that they were emplaced 'side-ways'. The present crust is --27 km thick (Bott et al. 1970) and may have been about 33 km thick allowing for later crustal extension (only 5% according to Shackleton et al.) and erosion (probably c. 5 kin). The Cornubian grarotes are typical 'S'-types of Chapple & White (1974) and are ascribed to a lower crustal origin by Exley, Stone & Floyd (1983) and Floyd, Exley & Stone (1983). A geothermal gradient of 30~ km -1 is required to produce an adamel-

161

The northern margin o f the Varbscides litic biotite-granite melt at a depth of - 2 8 km (see arguments in Exley et al. 1983). It is my contention that such melting is possible. Basin development gives rise to crustal thinning and an increased heat flux (McKenzie 1978). Subsequent tectonism producing crustal thickening by thrusting would depress the Moho into a region with a still elevated heat flow and produce melting. Thus granite formation is triggered by deformation and crustal thickening. Hence granite emplacement occurs after early deformation. Shackleton et al. (1982) estimate that the total shortening across SW England is - 5 0 % . In a thick-skinned model this would require 100% crustal thickening by a pure shear mechanism. A considerable amount of the shortening in the Gramscatho basin took place by underthrusting and underplating of oceanic, and thinned continental, crust. The remainder of the area had an initial crustal stretching of # = 2, which would have thinned a 3 0 k m crust to 15 km and developed a - 5 km sedimentary basin, thus giving a - 2 0 km crust prior to deformation. To thicken this to - 3 3 km at the end of the Carboniferous requires ~ = 1.7 and a crustal shortening of - 4 0 % . I would suggest that crustal thickening by thrusting could account for most of the shortening seen in the surface structure, with dextral shear making up the difference. This shear component would not give rise to crustal thickening. Thus I see little need for a master d6collement and envisage shortening of the basement at depth as well as in the cover.

Brittany The rocks of the Armorican massif preserve structures related to the Cadomian and Variscan orogenies. Three major domains can be delimited (Fig. 12), separated by major dextral shear zones--the North Armorican shear zone (Chauris 1969; Watts & Williams 1979) and the South Armorican shear zone (Cogne 1960; Jegouzo 1980). Whilst major dextral movements occur on the north and south Armorican shear zones there is significant Variscan deformation elsewhere in the region. Northern Brittany (Fig. 12) is characterized by Cadomian deformation, but with at least local Variscan reworking (see particularly Cabanis, in Roach 1980). The Central Domain was widely deformed during the Variscan, which produced steep folds and cleavage accompanied by a subhorizontal stretching lineation. These can be attributed to a dextral shear component (Gapais & Le Corre

~

A

I

N

~+~_.~-~<+5S

,__/] ~ DOMAIN

i R

FIG. 12. Map of Brittany showing the main structural zones and shear zones; NASZ--North Armorican shear zone, SASZ(N,S)--South Armorican shear zones (north and south branches). Granites (crosses), basins (stipple) and cleavage traces (lines) are only shown for the central domain. B--Brest; CB--Chateaulin basin; G--Ile de GrOix; N--Nantes; Q--Quintin granite; R--Rennes.

1980; Hanmer, Le Corre & Berthe 1982). This deformation intensifies towards the South Armorican shear zone (Berthe, Choukroune & Jegouzo 1979; Jegouzo 1980). South Brittany underwent complex deformation and metamorphism. The low temperature/high pressure assemblages of the Ile de Gr6ix and the high temperature/low pressure metamorphism of the mainland have naturally led to speculation of a paired metamorphic belt and related subduction (Lefort 1979). The age equivalence of the belts is still uncertain, but any subduction would seem to be an early, or even pre-, Variscan event. Ouinquis & Choukroune (1981) attribute this deformation to NW directed overthrusting, antithetic to northward underthrusting of subducted oceanic crust beneath south Brittany at around 4 0 0 M a . Lefort (1979) favours a later age for northward dipping subduction during the Middle and Upper Devonian. The timing of deformation in Brittany appears to extend from the late Devonian until late Carboniferous, Darboux et al. (in Roach 1980) describe Visean-Namurian molasse in the Chateaulin basin (Fig. 12) which lies unconformably on previously deformed Devonian. The latter would indicate a 'Bretonic' event, with the molasse basin being deformed in the late Carboniferous. The location of the Variscan leucogranites (dated at 320-300 Ma by

D. J. Sanderson

162

Vidal 1973; Mifdal 1979) is controlled by the South Armorican shear zone, but they are cut and mylonitized by later movements along it. This again indicates prolonged deformation which was not complete until the late Carboniferous, as small Stephanian basins, containing clasts of the mylonitized granites, are also mylonitized along the shear zones.

Wrench/thrust models One interesting feature of the strain field in Central Brittany is that away from the South Armorican shear zone the fabric still trends E - W , subparallel to the zone, for example S W of Rennes (Fig. 12). This is not the pattern predicted by simple shear (cf. Ramsey & Graham 1970). It is, however, consistent with a transpression model (Sanderson & Marchini 1984). Away from the shear zone, the compression normal to the zone produces a weak E - W fabric (e.g. west of Rennes). As the shear strain increases towards the northern branch of the South Armorican shear zone, the cleavage forms a gentle sigmoidal pattern as it swings firstly anticlockwise then clockwise (Fig. 12). This, coupled with the contact strains around the granites, forms the main cleavage pattern in central Brittany. Dextral shear essentially overlaps granite emplacement and basin development. Wrench tectonics, especially if controlled by basement faults, would be expected to give rise to local areas of dilation and compression at bends or terminations of the faults. Thus granite emplacement may have been controlled by local, transient zones of dilation along the main shear zones. The small Stephanian basins along the South Armorican shear zone and the larger basins, such as the Chateaulin basin (Fig. 12), may represent pull-apart basins. Whatever the age for the subduction in the Ile de Gr6ix, it clearly pre-dates the collision and dextral shearing responsible for the deformation in Central Brittany. It is, however, of some interest that the movement direction of this thrusting is between N - S and N W - S E . Since the zone must have trended roughly N W SE, the subduction would have been markedly oblique to it, thus generating a dextral shear component along the zone (Fig. 12).

Conclusions This analysis of the structural history of various areas in the northern m a r g i n of the Variscan belt in the British Isles and France shows

the following general features: (1) Basin development commenced at different times with little overall polarity, although a local sequence of south to north migration of basins in SW England is evident. (2) Basin development is often contemporaneous with regional deformation, and only locally is there a sequence of crustal stretching and later compression. This suggests no regional pattern of fluctuating extensional and compressional tectonics and is more consistent with a strike-slip model. Basins are considered to arise as local pull-aparts during an overall dextral shearing. (3) Only locally, as in the Lizard, is there evidence of basins being floored by oceanic crust. Most of the volcanicity is associated with continental rifting. (4) Deformation began in the Devonian and continued until the end of the Carboniferous. The end-Carboniferous deformation seems more important in Britain, with significantly earlier events in Brittany. Although the onset of deformation may have swept northwards much local complexity in timing exists and deformation continued to the end of the Carboniferous in all regions. (5) Most areas show evidence for a transpressive regime, with variable amounts of dextral shear and shortening normal to the belt. The most widespread manifestation of this is the obliquity of structures to the main 'zones' within the belt. The local transport direction is usually to the NW or NNW, in contrast to the general ESE trend of this part of the belt (Fig. 1). (6) The Variscan 'front' is poorly defined and the northern margin of the fold belt is best viewed as a transition from a more pervasive style of deformation, within the belt, to one of markedly heterogeneous deformation in the foreland. (7) Thrust tectonics dominate S W England and the Ile de Gr6ix blueschists belt. Even in these regions the transport direction is oblique to the belt implying a component of dextral shear along it. These features can be interpreted in terms of a dextral transpression model. Within a framework of dextral movement, local extensional and compressional 'events' occur. The small ocean basin, represented by the Lizard Complex, was obducted by NW thrusting which also affected the flysch trough (the Gramscatho

T h e n o r t h e r n m a r g i n o f the V a r i s c i d e s

basin) to the north. This deformation was accompanied by low t e m p e r a t u r e / m o d e r a t e pressure metamorphism and may be likened to a small accretionary subduction complex. The limited extent of oceanic crust and the lack of a blueschist belt suggest obduction within a marginal sea basin as suggested by Reading (1973) and A n d e r t o n et al. (1979) on largely stratigraphic grounds. Apart from this area the rest of the margin is en-sialic. The transpressive nature of the deformation (including 'transtension' in the sense of Harland 1971) suggests some sort of oblique or strikeslip boundary conditions. The distinction between oblique subduction and/or collision and strike-slip tectonics is one of degree. Using Fig. 7 as a model for the entire region, the angle of approach (~) determines the relative importance of strike-slip and compressional tectonics. Subduction would increase the plate convergence and hence p r o m o t e thrust tectonics (as in S W England?). H e t e r o g e n e o u s deformation within such a model would tend to concentrate deformation in the basins where the crust is thinnest and heat flow highest. Thus

163

the structure appears as a series of inverted basins developing at different times and showing variable amounts of dextral shear. Indeed some areas may remain undeformed. If these are large enough they may be considered as microplates, such as the Cadomian-Icartian microplate of Shackleton et al. (1982), which includes the north Brittany domain. Thus the structural style and sequence at the northern margin of the Variscan belt in west Europe support the dextral strike-slip plate tectonic model of Badham & Halls (1975) and Badham (1982). This dextral shear would be produced by oblique collision of the European and African plates, as is implicit in the palaeomagnetic reconstructions of Smith, Briden & Drewry ( 1 9 7 3 ) a n d Scotese et al. (1979). Finally if the N W tectonic transport, deduced for much of the belt in west Europe, operated on a more N E - S W trending boundary then no dextral c o m p o n e n t would occur. Instead a (more usual?) thrust-type tectonic regime would result as in the Appalachians and central Europe.

References

ANDERTON, R., BRIDGES, P. H., LEEDER, M. R. & SELWOOD, B. W. 1979. A Dynamic Stratigraphy of the British Isles. Allen & Unwin, London. BADHAM, J. P. N. 1982. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 493-504. - - & HALLS, C. 1975. Microplate tectonics, oblique collisions and the evolution of the Hercynian orogenic system. Geology, 3, 373-6. BARNES, R. P. & ANDREWS,J. R. 1981. Pumpellyiteactinolite grade regional metamorphism is south Cornwall. Proc. Ussher Soc. 5, 139-46. & BADHAM, J. P. N. 1979. Preliminary Investigations of south Cornish melanges. Proc. Ussher Soc. 5, 139-46. BERTHE, D., CHOUKROUNE,P. & JEGOUZO,P. 1979. Orthogneiss, mylonite and non-coaxial deformation of granites: the example of the South Armorican Shear Zone. J. strucr Geol. 1, 31-42. BLUCK, B. J. 1980. Evolution of a strike-slip faultcontrolled basin, Upper O.R.S. Scotland. In: BALLANCE, P. F. & READING, H. E. (eds) Sedimentation in Oblique-slip Mobile Zones. Spec. Publs int. Ass. Sediment. 4, 63-78. Blackwell Scientific Publications, Oxford. BOTT, M. H. P., HOLDER, A. P., LONG, R. E. & LUCAS,A. L. 1970. Crustal structure beneath the granites of southwest England. In: NEWALL,C. & RAST, N. (eds) Mechanism of Igneous Intrusion. Geol. J. Spec. Issue, 2, 83-102. BRENNAND, T. P. 1965. The Upper Carboniferous (Namurian) stratigraphy northeast of Castle-

island, Co. Kerry, Ireland. Proc. R. Ir. Acad. 41-63. BROOKS, M., MECHIE,J. & LLEWELLYN,D. J. 1983. Geophysical investigations in the Variscides of southwest Britain. In: HANCOCK,P. L. (ed.) Variscan Fold Belt in the British Isles, 186-97. Hilger, Bristol. BROWN,A. C. 1979. The Navan (Tara) deposit. In: BROWN, A. C. (ed.) Prospecting in Areas o f Glaciated Terrain, 25-33. Irish Association for Economic Geology. CAPEWELL,J. G. 1957. The stratigraphy and structure of the country around Sneem, Co. Kerry. Proc. R. Ir. Acad. 58B, 167-83. CHAPPELL,B. W. & WmTE,A. J. R. 1974). Two contrasting granite types. Pacif Geol. 8, 173-4. CHAURIS, L. 1969. Sur un important accident structural dans le nore-Ouest de l'Armorique. C. r. hebd. Sdanc. Acad. Sci., Paris, 268, 2859-61. CLAYTON, C. & GRAHAM, J. R. 1974. Miospore assemblages from the Devonian Sherkin Formation of southwest County Cork, Republic of Ireland. Pollen Spores, 16, 565-88. CoE, K. 1959. Boudinage structure in west Cork, Ireland. Geol. Mag. 96, 191-200. - 1966. Intrusive tufts of west Cork, Ireland. Q. J. geol. Soc. Lond. 122, 1-28. - - & SELWOOD,E. B. 1963). The stratigraphy and structure of the Beara peninsula, Co. Cork. Proc. R. Ir. Acad. 63B, 33-59. COGNE,J. 1960 Schistes cristallins et granites en Bregagne meridionale. Le domaine de l'Anticlinal 64B,

164

D . J. S a n d e r s o n

de Cornouaille. Mdm. Serv. Carte gOol. det. Fr. i-382. COWARD, M. P. & MCCLAY, K. R. 1983. Thrust tectonics of South Devon. J. geol. Soc. London, 140, 215-28. DAVIES, G. R. 1984. Isotopic evolution of the Lizard Complex. J. geol. Soc. London, 141, 3-14. DEARMAN, W. R. 1959. The structure of the Culm Measures at Meldon, near Okehampton, north Devon. Q. Jl geol. Soc. Lond. 115, 65-106. DEWEY, J. F. 1966. Kink bands in Lower Carboniferous slates of Rush, Co. Dublin. Geol. Mag. 103, 138-412. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. London, 139, 371-412. DODSON, M. H. &REX, D. C. 1971. Potassium-argon ages of slates and phyllites from southwest England. Q. Jl geol. Soc. Lond. 126, 465-99. DUBEY, A. K. 1980. Model experiments showing simultaneous development of folds and transcurrent faults. Tectonophys. 65, 69-84. EXLEY, C. S., STONE, M. & FLOYD, P. A. 1983. Composition and petrogenesis of the Cornubian granite batholith and post-orogenic volcanic rocks in southwest England. In: HANCOCK,P. L. (ed.) The Variscan Fold Belt in the British Isles, 153-77. Hilger, Bristol. FLOYD, P. A. 1982. Chemical variation of Hercynian basalts relative to plate tectonics. J. geol. Soc. London, 139, 505-20. 1983. Composition and petrogenesis of the Lizard Complex and pre-orogenic basaltic rocks in southwest England. In: HANCOCK,P. L. (ed.) The Variscan Fold Belt in the British Isles, 130-52. Hilger, Bristol. , EXLEY, C. S. & STONE,M. 1983. Variscan magmatism in southwest England--discussion and synthesis. In: HANCOCK,P. L. (ed.) The Variscan Fold Belt in the British Isles, 178-85. Hilger, Bristol. CAPAIS, D. & LE CORRE, C. 1980. Is the Hercynian belt of Brittany a major shear zone? Nature, 288, 574-5. GILL, W. D. 1962. The Variscan fold belt in Ireland. In: COE, K. (ed.) Some Aspects o f the Variscan bold Belt, 49-64. Manchester University Press. GOLDRING, R. 1962. The bathyal lull: Upper Devonian and Lower Carboniferous stratigraphy of north Cornwall and west Devon. In: COE, K. (ed.) Some Aspects o f the Variscan Fold Belt, 75-91. Manchester University Press. GRAHAM, J. R. 1983. Analysis of the Upper Devonian Munster basin, an example of a fluvial distributary system. In: COLLINSON, J. O. & LEWIN, J. (eds) Modern and Ancient Fluvial Systems. Spec. Publs int. Ass. Sediment. 6, 473-83. Blackwell Scientific Publications, Oxford. HALLIDAY, A. N. & M~TCHELL,J. G. 1977. K-Ar ages of basaltic dykes from the Lizard Complex. Proc. Ussher Soc. 4, 24. HANMER, S. K., LE CORRE, C. & BERTHE, D. 1982. The role of Hercynian granites in the deformation and metamorphism of Brioverian and Palaozoic rocks of Central Brittany. J. geol. Soc. London, 139, 85-93.

-

-

HARLAND, W. B. 1971. Tectonic transpession in Caledonian Spitsbergen. Geol. Mag. 108, 27-42. HOBSON, D. M. 1976a. A structural section between Plymouth and Bolt Tail, South Devon. Proc. geol. Ass. 84, 283-313. 1976b. The structure or the Dartmouth antiform. Proc. Ussher Soc. 3, 320-32. -& SANDERSON, D. J. 1983. Variscan deformation in southwest England. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 108-29. Hilger, Bristol. HOUSE, M. R. 1971. Devonian faunal distributions. In: MIDDLEMISS, F. A., RAWSON, F. A. & NEWALL, G. (eds) Faunal Provinces in space and Time. Geol. J. Spec. Issue, 4, 77-94. JEGOUZO, P. 1980. The South Armorican Shear Zone. J. struct. Geol. 2, 39-47. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles--Caledonian inheritance verses Hercynian plate margin processes. J. geol. Soc. London, 139, 479-91. LEFORT, J.-P. 1979. Iberian-Armorican arc and Hercynian orogeny in western Europe. Geology, 7, 384-8. MATrHEWS, S. C. 1977. The Variscan fold belt in southwest England. Neues Jb. Miner. Geol. Paliiont. Abh. 154, 94-127. --, NAYLOR, D. & SEVASTOPULO, G. D. 1983. Palaeozoic sedimentary sequence as a reflection of deep structure in southwest Ireland. Sediment. Geol. 34, 83-9. MCKENZIE, D. P. 1978. Some remarks on the development of sedimentary basins. Earth planet. Sci. Lett. 40, 25-32. MIFDAL, A. 1979. La methode Rubidium-Strontium. Example d'apptication: le granite de Pontivy. Unpublished D.E.A. University de Rennes. MOORE, J. McM. 1975. Fault tectonics at Tynagh mine, Ireland. Trans Insm min. metall. 84B, 141-5. MOORE, J. McM. 1975. Fault tectonics at Tynagh mine, Ireland. Trans. Inst. Min metall.. 84B, 141-5. MORTON, W. H. 1965. The Carboniferous stratigraphy of the area northwest of Newmarket, Co. Cork, Ireland. Sci. Proc. R. Dubl. Soc. 2A, 47-65. MURPHY, T. 1960. Gravity anomaly map in Ireland, sheet 5--south west. Geophys. Bull., Dubl. 18. NAYLOR, D. 1978. A structural section across the Variscan fold belt, southwest Ireland. J. Earth Sci. R. Dubl. Soc. 1, 63-70. & JONES, 1967. Sedimentation and tectonic setting of the Old Red Sandstone of southwest Ireland. ln: OSWALD, D. H. (ed.) Proc. int. Syrup. Devonian System, Calgary, 1089-99. --, REILLY, T. A., SEVASTOPULO, G. D. & SLEEMAN, A. C. 1983. Stratigraphy and Structure in the Irish Variscides. In: HANCOCK, P. L. (ed.) Variscan Fold Belt in the British Isles, 20-46. Hilger, Bristol. -• SEVASTOPULO, G. D. 1979. The Hercynian 'front' in Ireland. Krystalinikum, 14, 77-90. OWEN, T. R. & WEAVER, J. D. 1983. The structure of the Main South Wales coalfield and its margins. -

-

T h e n o r t h e r n m a r g i n o f the V a r i s c i d e s

-

-

In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 74-87. Hilger. Bristol. PENNY, S. R. 1978. Devonian lavas from the Comeragh Mountains, Co. Waterford. J. Earth Sci. R. Dubl. Soc. 1, 71-6. PnlLCOX, M. E. 1964. Compartment deformation near Buttevant, County Cork, Ireland and its relation to the Variscan thrust front. Sci. Proc. R. Dubl. Soc. 2A, 1-11. PHILLIPS, W. E. A., STILLMAN, C. J. & MURPHY, T. 1976. A Caledonian plate tectonic model. J. geol. Soc. London, 132, 579-609. QUINQUIS, H. & CHOUKROUNE, P. 1981. Les schistes bleus de L'ile de Groix dans la chaine hercynienne: implications cinematiques. Bull. Soc. gdoL Fr. 23, 409-18. RAMSAY, J. G. & GRAHAM, R. H. 1970. Strain variation in shear belts. Can. J. Earth Sci. 7, 786-813. RAMSAY, J. R. 1980. The crack seal mechanism of rock deformation. Nature, 284, 135-9. RATTEY, P. R. 1979. Deformation in south-west Cornwall. Proc. Ussher Soc. 5, 39-43. & SANDERSON, D. J. 1982. Patterns of folding within nappes and thrust sheets: examples from the Variscan of SW England. Tectonophys. 88, 247-67. & -1984. The structure of SW Cornwall and its bearing on the emplacement of the Lizard Complex. J. geol. Soc. London, 141, 87-95. READING, H. G. 1973. The tectonic environment of southwest England: discussion of Floyd (1972). Proc. geol. Ass. 84, 237-9. REILLY, T. A. & GRAHAM, J. R. 1972. The historical and geological setting of the Glandore mines, southwest County Cork. Bull. geol. Surv. lrel. 2, 1-13. ROACH, R. A. 1980. Structure and tectonic evolution of the Armorican Massif. J. geol. Soc. London, 137, 211-6. ROBERTS, J. L. & SANDERSON, D. J. 1971. Polyphase development of slaty cleavage and the confrontation of facing directions in the Devonian rocks of north Cornwall. Nature, 230, 87-9. RUSSELL, K. J. 1978. Vertebrate fossils from the Iveragh Peninsula and the age of the Old Red Sandstone. J. Earth Sci. R. Dubl. Soc. 1, 151-62. SADLER, P. 1973. An interpretation of new stratlgraphic evidence from south Cornwall. Proc. Ussher Soc. 2, 535-50. SANDERSON, D. J. 1971. Superposed folding at the northern margin of the Gramscatho and Mylor beds, Perranporth, Cornwall. Proc. Ussher Soc. 2, 266-9. 1979. The transition from upright to recumbent folding in the Variscan fold belt of southwest England: a model based on the kinematics of simple shear. J. struct. Geol. 1, 171-80. -

-

-

-

--.,

165

ANDREWS, J. R., PHILLIPS, W. E. A. & HUTFON, D. W. H. 1980. Deformation studies in the Irish Caledonides. J. geol. Soc. London, 137,289-302. • DEARMAN, W. R. 1973. Structural zones of the Variscan fold belt in SW England: their location and development. J. geol. Soc. London, 129, 527.-33. & MARCHINI 1984. Transpression. J. struct. Geol. in press. SCOTESE, C., BAMBACH, R. K., BARTON, C. VAN DER VOO, R. & ZIEGLER, A. 1979. Palaeozoic base maps. J. Geol. 87, 217-78. SHACKLEJON, R. M., RIES, A, C. & COWARD, M. P. 1982. An interpretion of the Variscan stuctures in SW England. J. geol. Soc. London, 139, 533-41. SHERIDAN, D. J. 1964. The structure and mineralization of the Mountain Mine area, Allihies, west Co. Cork, Ireland. Sci. Proc. R. Dubl. Soc. 2A, 21-7. SMITH, A. G., BRIDEN, J. C. & DREWRY, G. E. 1973. Phanerozoic world maps. In: HUGHES, N. F. (ed.) Organisms and Continents Through Time. Spec. Pap. palaeont. Lond. 12, 1-42. STYLES, M. T. & RUNDLE, C. 1981. A whole rock Rb/Sr isochron for the Kennack gneisses and a discussion of the age of the Lizard Complex. Proc. Ussher Soc. 3, 255-62. SYLVESTER, A. G. & SMITH, R. R. 1976. Tectonic transpression and basement controlled deformation in San Andreas fault zone, Salton Trough, California. Bull. Am. Ass. Petrol. Geol. 60, 2081-102. TALWANI, M., WORZEL, J. L. & LANDISMAN, M. 1959. Rapid gravity computations for twodimensional bodies with application to the Mendocino submarine fracture. J. geophys. Res. 64, 49-59. TAYLOR, S. & ANDREW, C. J. 1978. Silvermines orebodies, Country Tipperary, Ireland. Trans. Instn Min. metall. 87B, 111-24. VIDAL, PH. 1973. Premieres donnees geochronologiques sur le granites hercyniens du Sud du Massif armoricain. Bull. Soc. g~ol. Fr. 15, 239-45. WALSH, P. T. 1968. The Old Red Sandstone west of Killarney, Co. Kerry, Ireland. Proc. R. Ir. Acad. 16B, 9-26. WATTS, M. J. & WILLIAMS,G. D. 1979. Fault rocks as indicators of progressive shear deformation in the Guingamp region, Brittany. J. struct. Geol. 1, 323-32. WILCOX,R. E., HARDING,T. P. & SEELY,D. R. 1973. Basic wrench tectonics. Bull. Am. Ass. Petrol. Geol. 57, 74-96. -

-

DAVID J. SANDERSON, Department of Geology, Queen's University, Belfast BT7 INN, Ireland.

Structural style, shortening estimates and the thrust front of the Irish Variscides M. A. Cooper, D. Collins, M. Ford, F. X. Murphy & P. M. Trayner SUMMARY: Previous work in the Irish Variscan orogen has emphasized the significance of the major folds in the overall structure. A series of three balanced sections through the orogen reveals that thrusts are an equally important element of the orogenic structure. The 'thrust front' is not a fundamental structural line delimiting different deformational zones as previously proposed. Major thrusts occur along the 'thrust front' between Killarney and Mallow, but are absent east of Mallow. This change in character of the 'thrust front' is a function of a change in the stratigraphic level of the sole thrust to the orogen and the current erosion level. The deeper level of the sole thrust produces the numerous Silurian cored inliers to the NE of Mallow. The balanced sections show that the overall bulk shortening in the orogen increases from 33% in the east, to 42% in the central areas. This increase may be related to the arcuate nature of the orogen.

In the southern part of Ireland the Variscan orogeny strongly deformed t h e ' U p p e r Palaeozoic sedimentary sequence together with the underlying Lower Palaeozoic and Precambrian basement. The location map (Fig. 1 inset) shows the major tectonic features of the orogen. To the south of the Dingle Bay-Dungarvan line, traditionally regarded as the 'thrust front' (Gill 1962), the structure is dominated by large upright folds, faults and thrusts. In this area the structures have E N E - W S W to E - W trends, elsewhere the underlying Caledonian trend may exert an influence. On the south coast the folds face southwards, becoming upward facing around Cork City and northward facing as the 'thrust front' is approached. Cleavage is usually developed throughout this zone and is apparently axial planar to the folds. Small-scale thrusts are commonly seen in outcrop, and larger-scale thrusts may be inferred from surface structures. Steeply dipping dextral wrench faults are developed parallel to orogenic strike in west Cork. To the north of the thrust front the deformation has been considered less intense (Gill 1962). Many of the folds face north and have a morphology indicative of underlying d6collements. Naylor (1978) presents a cross-section through the western part of the orogen which emphasizes the major fold structures. The aim of this paper is to present three balanced sections through the orogen and to discuss the nature of the Variscan 'thrust front' in Ireland.

The evidence for thrusting in the Irish Variscan The Irish Variscan has many structural features typical of other recognized foreland fold and thrust belts, e.g. the Canadian Rockies (Price 1981), the Appalachians (Roeder et al. 1978; Gwinn 1970) and the Norwegian Caledonides (Hossack et al. 1984). Published data suggest that many of the faults in the orogen lie parallel to orogenic strike (Naylor et al. 1981, fig. 13E, F). These have previously been termed strike faults (Reilly & Graham 1972; Naylor 1969). In most cases the sense of displacement on these steeply dipping faults is unknown, although some have been shown to be reverse faults (Naylor 1978; Naylor & Sevastopulo 1979; Naylor et al. 1981). These faults were not termed thrusts as they differ from the old definition of a thrust as a low angle reverse fault. This definition is now redundant in view of recent developments in thrust tectonics (Dahlstrom 1970; Price & Mountjoy 1970; Harris & Milici 1977). We prefer the definition proposed by McClay (1981) that 'a thrust is a map scale contraction fault'. Under this definition many of the strike faults in the Irish Variscan could be considered to be thrusts. Minor thrusts are often exposed in the areas we have studied, e.g. at White Bay on the east side of Cork Harbour where the Kinsale Formation is imbricated b y a number of thrusts dipping 30-70~ Major thrusts are rarely

167

168

M. A . C o o p e r et al.

exposed, an important exception however occurs at D r o m a n e e n House near Mallow. At this locality sandstones of the Kiltorcan Formation are thrust over supra-reef Visean limestones (T. Mulcahy, pets. comm.). The movement on this thrust is at least 2 km and it transfers significant amounts of sole thrust slip into the Namurian. Where exposure is poor major thrusts can be inferred from space problems encountered during detailed mapping. An example occurs on the northern limb of the Ardmore syncline (Fig. 1). Here there is insufficient room to accommodate the Kiltorcan and Kinsale Formations between outcrops of the Old Red Sandstone and the Courtmacsherry Formation. The missing stratigraphy is due to a foreland dipping thrust; the hinge of the anticline complimentary to the syncline is truncated in the hanging wall. Movement surfaces are ubiquitous, often mineralized by quartz and chlorite displaying slickenside fibres. Some of these surfaces are due to flexural slip during folding, some are thrusts and others have been active during both processes. Measurement of the fibres reveals that they are oriented perpendicular to fold trend and orogenic strike (Figs 1C, D, E & 3C). This is consistent with their development on thrust and flexural slip surfaces.

Several workers have previously regarded thrusts as an important element of the orogen's structure. Philcox (1964) in a perceptive paper on the structure of the Buttevant area immediately north of the 'thrust front', describes a number of south-dipping thrusts. He found that structurally distinct areas were compartmentalized by N - S faults, each fault block having an individual structure. Such faults, oriented perpendicular to orogenic strike may reflect the position of subsurface lateral ramps (Dahlstrom 1970; Boyer & Elliott 1982; Hossack et al. 1984). We believe that this may be a valid model for the Buttevant area. Slightly further north in Co. Tipperary, Shelford (1963) has described minor thrusting in the Namurian and a major thrust carrying the Old Red Sandstone over the Namurian succession. Nevill (1966) describes thrusts from the Kanturk Coalfield, situated at the northern end of Section 3. The coal seams are pulverized and locally tectonically thickened, suggesting they have acted as d6collement planes. Several N - S wrench faults apparently connect thrust planes in the coalfield (W. Nevill, pets. comm.) which we again interpret as reflecting subsurface lateral ramps. Thrust tectonics exist elsewhere along the Variscan 'front'. They have been described in South Wales (Hancock et al. 1981), in north-

F1G. 1. Section 1. (A) is the deformed section. (B) is the restored section. (C), (D) and (E) are stereograms of slickensides +, minor fold axes and cleavage/bedding intersection lineations 0; (C) Dungarvan Syncline; (D) east of Youghal; (E) the area between Youghal and Cork. The inset map shows the locations of the three sections and the major structural features of the orogen; Co, Comeragh Mountains; Kn, Knockmealdown Mountains; CS, Cork syncline; DS, Dungarvan syncline; TS, Tallow syncline.

169

T h r u s t f r o n t o f the Irish Variscides western France where the structure is largely obscured by Mesozoic cover (Wallace 1968; Cooper et al. 1983) and in Germany (Meissner et al., this volume). It is thus consistent that thrust tectonics should exist in Ireland.

Method of section construction The sections presented in this paper are balanced sections (Hossack 1979). They were constructed following the recommended procedure of Elliott & Johnson (1980). An inherent assumption of the sections is that the folds that dominate the outcrop pattern of the Irish Variscan are underlain by a major d6collement horizon. The relationship between folds and underlying d6collements has been discussed by Carey (1962), Dahlstrom (1969) and Gwinn (1970). The deformation is in this case developed above a d~collement and does not affect the strata beneath the d~collement. The alternative is that folding continues downwards to depths where ductile processes can accommodate the shortening. However, where pre-Devonian rocks are exposed in the Irish Variscan no significant Variscan ductile deformation is seen, thus precluding the latter model. We therefore prefer the model of deformation above a d6collement with the sole thrust of the orogen lying largely within pre-Devonian strata (but see Sanderson, this volume). The sections were balanced by both line length and area, and are thus internally consistent. Area balancing accounts for any layer parallel shortening process, e.g. by cleavage formation, which may well be important in the finer grained clastic sediments and locally in the limestones. The pin lines chosen at the northern, foreland, end of the sections are either axial planes or perpendicular to horizontal strata. The lack of data to the north precludes extending the sections northwards beyond the limit of obvious Variscan deformation. The southern limit of the sections is the coast. The surface data on the outcrop of contacts and bedding dip have been taken from Geological Survey of Ireland 1" to 1 mile Geological Map sheets, published work and our own observations along the section lines. Detailed mapping along the section lines is still in progress, and when complete will improve the validity of the sections. The poor control on deep structure may be helped by the proposed BIRPS deep seismic line. In constructing a balanced deformed section a restored section must be constructed simultaneously. In Section 1 (Fig. 1B) stratigraphic

thicknesses remain fairly constant and a layercake template is easily constructed on to which the restored section can be plotted. Sections 2 and 3 (Figs 2B & 3B) present a more complicated problem as stratigraphic thicknesses vary due to facies changes. No simple template can be constructed on which to plot the restored section; instead an initial approximation of the template based on measured sections and estimates of bulk shortening is iteratively modified during the balancing procedure. The errors introduced during the construction of Sections 2 and 3 were small since the restored sections approximate to gradually thinning wedges over large distances. The accuracy of the deformed sections is clearly partially dependent on the quality of data used in constructing the templates. Much stratigraphic and sedimentological research has been conducted in the Irish Variscan. Consequently there is no shortage of published data on stratigraphic thicknesses derived from detailed logging. We have used stratigraphic nomenclature based on that of Naylor et al. (1981) with some modifications to facilitate section construction. The Toe Head Sandstone Formation in the coastal areas and the Kiltorcan Beds in the north have both been included within the Old Red Sandstone. The Kinsale Formation in the south has been correlated with the Lower Limestone Shales in the north. Although these units are not exactly time or facies equivalent the errors are minimal on the scale at which the sections are presented. Their base, the Devonian-Carboniferous boundary, was chosen as the horizontal datum for the sections. The basement shown on the sections includes all pre-Devonian rocks. It should be stressed that the three sections presented are not unique solutions; they are possible interpretations of the available data. The minor differences in the style of the sections reflects individual preferences in section construction. We opted for structures which yielded the minimum shortening estimates, the involvement of less basement in the thrusts would increase shortening.

Section 1 This section runs from the coast at Youghal along the River Blackwater through Cappoquin to Ardfinnan in the north (Fig. 1). The stratigraphic template was constructed using the thickness data of Shelford (1963); Sleeman et al. (1978); MacCarthy et al. (1978); Naylor et

170

M.A.

C o o p e r et al.

al. (1981); and our own observations, which

extend as far north as the Dungarvan syncline. We are grateful to S. Baldy (pers. comm.) for information on the structure of the Devonian rocks eastwards along strike from the northern part of the section. Considerable thickening of Devonian rocks in the Comeragh Mountains has been noted (Capewell 1957; Penney 1980). The geometry of the Comeragh sedimentary basin is unknown, and basal units are predominantly conglomeratic (Penney 1980). An important d6collement remains at a virtually constant depth beneath the top of the Old Red Sandstone, coinciding with the change from conglomerates to finer grained sediments. This local thickening of the conglomeratic facies appears to behave in structural unity with the basement and for that reason is ornamented as basement on the section. One of the main features of the section (Fig. 1A) is the large footwall ramp in the sole thrust beneath the north limb of the Dungarvan syncline. The northward movement of the hangingwall block caused stratigraphic duplication and produced the box-fold morphology of the Knockmealdown anticlinorium (S. Baldy, pers. comm.). Shortening to the south of this large footwall ramp is accomplished by a series of orogen and foreland dipping thrusts, which have climbed from the major ddcollement at the base of the finer grained Devonian sedi-

ments. The major foreland dipping thrusts have been mapped in the field by the geometry of stratigraphic cut-out along strike. The Cork and Dungarvan synclines are the result of intersecting orogen and foreland dipping thrusts producing large triangle zones (Boyer & Elliott 1982). At the northern end of the section the sole thrust is still beneath the current erosion level and carries 10 km of slip northwards beyond the pin-line into the foreland. The overall shortening represented in this section is 33% ( - 0 . 4 0 natural strain) measured on the Devonian-Carboniferous boundary.

Section 2 This section (Fig. 2A) runs from Frower Pt near Kinsale on the south coast north to Buttevant via Ballincollig and Mallow (Fig. 1 inset). The stratigraphic thickness data, on which the template for this restored section (Fig. 2B) is based, comes from the following sources: Hudson & Philcox (1965); Naylor (1966), Naylor (1975), Sleeman et al. (1978), MacCarthy et al. (1978), MacCarthy & Gardiner (1980), Gardiner & MacCarthy (1981), Naylor et al. (1981) and Sevastopulo (1982). in addition we received much valuable information from Dr A. Sleeman (pers. comm.), on revised stratigraphic thicknesses and outcrop distribution south of

FIc. 2 Section 2. (A) is the deformed section. (B) is the restored section which is split at D'.

171

Thrust front o f the Irish Variscides the Cork syncline. These data together with the map of Philcox (1964) for the Buttevant area and our own observations in the Mallow and Cork Harbour area have added considerably to the surface outcrop and structural data available on the 1" to 1 mile Geological Survey sheets. The sole thrust lies within the basement and climbs to the base of the Devonian in two footwall ramps, the larger of which occurs beneath the northern limb of the Cork syncline (Fig. 2A). The matching hangingwall ramp has been moved c. 13 km northwards and the stratigraphic duplication produces an anticlinorium south of Mallow (Fig. 2A). In the Mallow area much of the slip on the sole thrust is transferred via a series of thrust ramps to higher stratigraphic levels, notably within the Westphalian and Namurian (Fig. 2A, B). The overall bulk shortening represented by Section 2 is 37% ( - 0 . 4 6 natural strain) measured on the D e v o n i a n - - C a r b o n i f e r o u s boundary. The section could have been balanced without involving the basement in the thrusting, but would have required large-scale overthrusting of the Carboniferous by Devonian rocks resulting in a bulk shortening of 70% ( - 1 . 2 natural strain). The figure of 37% is thus a minimum estimate.

Section 3 Section 3 runs from Galley Head to Kanturk via Macroom (Figs 1 inset & 3A). The stratigraphic thickness data used in constructing the template for the restored section were taken from the following sources; Hudson & Philcox (1965), Naylor (1966), Reilly & Graham (1976), Keegan (1977), Naylor et al. (1981), Gardiner & MacCarthy ( i 9 8 t ) and Sevastopulo (1982). The Glandore High (Naylor et al. 1974) causes some southward thinning of the Carboniferous beds in the southern part of the section. Our own structural observations currently extend from Galley Head to Enniskeane (Fig. 3A). The structural style of the section north of the Cork syncline is very similar to Section 2 (compare Figs 2(A), A - D & 3(A), A - C ) . A series of large ramps just south of Kanturk transfer much of the slip on the sole thrust into the Westphalian. To the south of Macroom the structure is dominated by a series of large-scale pop-ups and triangle zones which root in a ddcollement at the base of the Old Red Sandstone (Fig. 3(A), E). The bulk shortening in Section 3 is 42% ( - 0 . 5 4 natural strain). The three sections display many common features whilst reflecting a pattern of increasing bulk strain to the west. The major d6collement

FIG. 3. Section 3 (A) is the deformed section. (B) is the restored section which is split at F'. (C) is a stereogram of slickensides (+), minor fold axes and cleavage/bedding intersection lineations (0) for the area between Galley Head and the point F. Key as in Fig. 2.

172

M. A. Cooper et al.

horizons are the base of the Devonian, the Lower Limestone Shales, Kinsale Formation, Namurian and horizons within the Westphalian. The observed but atypically high angles between bedding and thrusts suggests that some buckling may have preceded thrusting. In the areas where we have completed detailed mapping several foreland dipping 'backthrusts' have been recognized (Sections 1 and 3). These in combination with orogen dipping thrusts produce pop-ups and triangle zones (Boyer & Eiliott 1982) which are considered to be indicative of an underlying d6collement in other orogenic belts (Jacobeen & Kanes 1974; Harris & Milici 1977; Boyer & Elliott 1982)

The nature of the Variscan thrust front The Variscan thrust front in Ireland was recognized by Gill (1962) as a line from Dingle Bay to Dungarvan across which a change in structural style occurs. The concept of the thrust front has prompted much discussion since Gill's original paper. Wingfield (1968) recognized the presence of a high angle reverse fault between the Old Red Sandstone and Carboniferous limestone in the Killarney area, although Naylor & Sevastopulo (1970) suggested that this fault was vertical with a wrench displacement. Walsh (1968), working to the west of Killarney believed that the outcrop pattern did not necessitate a major faulted contact. Gill (1962) recognized the occurrence of a number of closely spaced thrusts in the Mallow area, forming an imbricate zone. However, further east there is a marked absence of major thrusts at the surface. This fact has led many authors to question the validity of the thrust front concept. More recently, the presence of high angle reverse faults along the thrust front was noted by Naylor (1978), however he commented that their displacements were likely to be small. Naylor & Sevastopulo (1979) suggested that the intensity of deformation shows a gradual decrease across the Dingle-Dungarvan line, and in common with the Variscan orogen elsewhere in Europe, the 'thrust front' is only locally marked by faults and thrusts. Matthews (1974) suggested that such faulting was locally induced by thinning of the Upper Palaeozoic succession on to shallow basement massifs. The importance of basement control was also dealt with by Gardiner (1978), who regarded the Irish Variscan front as a localized foreland feature, resulting from deformation at the northern boundary of a graben structure in the south-

ernmost part of Ireland, and thus having a direct link with the Variscan front in South Wales. Naylor et al. (1981) suggest that the front is broadly coincident with a zone in which deformation intensity decreases northwards, the major control on which is felt to be the basement. These authors also comment that major faulting is present only when the front coincides with the northern margin of the Munster basin, resulting from the adjustment of the sedimentary prism to faulted basement. It is noted that such faults should be extensional whilst those observed are reverse. This may simply be due to reactivation of small basin margin faults during deformation. The whole concept of thrust and deformation 'fronts' is a dubious one. The intensity of deformation is unlikely to decrease abruptly across any given line. It is far more likely to decay gradually into the foreland with a slow change in structural style. In other orogenic belts it is now recognized that the delimiting of a deformation front is fruitless, e.g. the Appalachians (Nickelsen 1966). The Canadian Rocky Mountain front ranges end abruptly at the McConnell thrust. Topographically this appears to be the thrust front since it is the last major thrust to outcrop (Price 1981). However, gas and oil exploration has revealed extensive thrusting east of the 'thrust front' which is difficult to recognize in the poorly exposed ground of the foothills. The three balanced sections we have presented illustrate the role of thrusting in the Irish Variscan orogen. Some of these thrusts outcrop, but are rarely exposed due to drift cover. Other thrusts do not outcrop as they lie beneath the current level of erosion; these are termed blind thrusts (Thompson 1981). Blind thrusts can be recognized by the diagnostic structures produced in the strata above them (Harris & Milici 1977; Thompson 1981), e.g. in Section 1 north of the Dungarvan syncline (Fig. 1). Between Killarney and Mallow the 'thrust front' is sharply defined by a series of thrusts with 5-7 km of displacement (Figs 2 & 3). Here the sole thrust is transferred by footwall ramps to higher stratigraphic levels; it is these thrust ramps which define the 'thrust front'. To the east of Mallow no large thrust faults have been recognized, and the 'thrust front' line follows the northern limb of the Dungarvan syncline. Section 1 illustrates the typical structure of this area, where the sole thrust stays blind at a deep level until much further north than in the sections to the west. A large ramp in the footwall of the sole thrust beneath the northern limb of the Dungarvan syncline trans-

Thrust front of the Irish Variscides fers the sole thrust to the basal Devonian d6collement. The major footwall ramp in the sole beneath the Cork syncline in Sections 2 and 3 probably connects with the ramp beneath the Dungarvan syncline in Section 1 via a large lateral ramp. This lateral ramp does not produce a surficial N - S fault, but may well be the cause of the westward termination of the Tallow syncline. The change in character of the thrust front is very well illustrated by the cumulative shortening graphs calculated from our sections (Fig. 4). These plot the cumulative shortening measured on the D e v o n i a n Carboniferous boundary against distance from the foreland pin lines. In Sections 2 and 3 an abrupt decrease in shortening occurs at the thrust front. In Section 1 shortening remains constant across the thrust front because the sole thrust is blind. A similar argument based on sole thrust footwall topography could be developed to explain the absence of thrusts at the 'thrust front' to the west of Killarney. The structural features observed along the 'thrust front' thus vary depending on the geometry and stratigraphic level of the sole thrust footwall. A similar change from outcropping thrusts to predominantly blind thrusts occurs from S to N in the foreland thrust and fold belt of the Canadian Rockies (Thompson 1981). Here, the cause appears not to be footwall topography but a facies change along strike. Northwards a higher percentage of incompetent strata occur in the stratigraphic prism than to the south. Thompson (1981) argues that this has created conditions more favourable to producing d6collement horizons that remain flat and blind,

601 5O 40 e%30

Thrustf r o n

3

2o 1

0 10 20 30Dista40 50 60 710 810 ncein krn ,

,

,

I

I

,

FIG. 4. Cumulative shortening (e%) measured on the Devonian-Carboniferous boundary, plotted against distance measured from the foreland pin-lines for Section 1 O, Section 2 II and Section 3 9 The curves are normalized to the thrust front. In Sections 2 and 3 an abrupt decrease in e occurs at the thrust front. In Section 1 e remains constant.

173

than the ramp and flat thrust trajectories through the thick carbonate and quartzite units to the south. We do not at present have any evidence for a similar facies change to explain our structural style change along strike.

Discussion Our interpretation of the 'thrust front' provides a possible explanation for the outcrop distribution of the Upper Palaeozoic in the area to the north. North of the line between Killarney and Mallow there is a large outcrop of Namurian strata with no large inliers of Silurian, Devonian and lower Carboniferous. The sole thrust is at the base of the Namurian, acting as a d6collement above which the Namurian is folded and thrust northwards (Morton 1965; Naylor 1978). The absence of inliers suggests that the sequence beneath the sole thrust is undeformed. To the east of Mallow the sole thrust is at the base of the Devonian (Fig. 1). North of the Knockmealdown Mountains the Devonian thins abruptly and the thick fine grained clastic units wedge out. This may cause the sole thrust to cut down-section stratigraphically northwards into the Silurian, whilst remaining at a constant structural level. Fold development above this d6coilement produced the anticlinal Silurian cored inliers that form prominent features to the north. Locally slip transfer to the higher levels occurs via minor thrusts such as those described by Shelford (1963). Using the inliers as a guide to the stratigraphic level of the sole thrust, it must lie within the Silurian until north of a line drawn east from Galway Bay where the inliers end. Here it must climb to a higher level, probably within the Carboniferous limestone, transferring by now a limited amount of slip. It is interesting to note that this coincides approximately with the northern boundary of the zone of concentric folding proposed by Gill (1962). The three restored sections (Figs 1B, 2B & 3B) illustrate some important features of the restored Upper Palaeozoic sedimentary prism. The Cork-Kenmare line, south of which no Carboniferous limestone is found, when restored moves 22.5 km south on section line 2 and 33~ km south on section line 3. Similar corrections can be applied to any of the other facies and thickness changes. This has implications for current views on the shape and palaeoenvironment of the Munster basin, e.g. Gardiner & MacCarthy (1981). Recent regional reviews of the Irish Variscan

174

M. A. Cooper et al.

( N a y l o r et al. 1981; G a r d i n e r & M a c C a r t h y 1981) have discussed fault control of basin margins and s e d i m e n t a t i o n . T h e c o n s t r u c t i o n of synoptic stratigraphic diagrams with e x t r e m e vertical e x a g g e r a t i o n has p r o b a b l y c o n t r i b u t e d to this concept, due to the steep gradients of stratigraphic b o u n d a r i e s that result. O u r r e s t o r e d sections indicate that such steep gradients are not present, and m a j o r b o u n d a r y faults n e e d not be inferred. M i n o r faults may well have b o u n d e d the n o r t h e r n margin of the M u n s t e r basin, reflected in the rapid facies change, but their significance has probably been o v e r e s t i m a t e d in the past (e.g. N a y l o r et al. 1981; G a r d i n e r & M a c C a r t h y 1981). B a s e m e n t block faulting is of local significance in controlling s e d i m e n t a t i o n , e.g. the G l a n d o r e high ( N a y l o r et al. 1974; K e e g a n 1977). T h e s t e r e o g r a m s of m i n o r fold axes and

cleavage b e d d i n g intersection lineations (Figs 1C, D, E & 3C) show a c h a n g e in t r e n d f r o m W N W - E S E in the east to E N E - W S W in the central part of the orogen. This arcuate n a t u r e of the o r o g e n is also well displayed by the m a j o r fold structures. It is interesting to speculate w h e t h e r the a r c u a t e g e o m e t r y is due to differential a m o u n t s of s h o r t e n i n g along strike. O u r sections suggest that this m i g h t be the case, as s h o r t e n i n g increases from 33% in the east (Section 1) to 42% in the west (Section 3). ACKNOWLEDGMENTS:We wish to thank Peter Brfick, Mike Coward and Norman Fry for useful discussions on this paper. The Geological Survey of Ireland provided financial support for D. Collins, F. Murphy and P. Trayner. University College, Cork provided financial support for M. Ford. Miss P. Hegarty is thanked for typing the manuscript.

References BOYER, S. E. & ELLIOTT, D. 1982. Thrust systems. Bull. Am. Ass. Petrol. Geol. 66, 1196-230. CAPEWELL, J. G. 1957. The stratigraphy, structure and sedimentation of the Old Red Sandstone of the Comeragh Mountains and adjacent areas, County Waterford, Ireland. Q. Jl geol. Soc. Lond. 112, 393-412. CAREY, W. S. 1962. Folding. J. Alberta Soc. Petrol Geol. 10, 95-144. COOPER, M. A., GARTON, M. R. & HOSSACK,J. R. 1983. The origin of the Basse Normandie Duplex, Boulonnais, France. J. struct. Geol. 5, 139-52. DAHLSTROM, C. D. A. 1969. Balanced cross sections. Can. J. Earth Sci. 6, 743-57. 1970. Structural geology in the eastern margin of the Canadian Rocky Mountains. Bull. Can. Petrol. Geol. 18, 332-406. ELLIOT, D. & JOHNSON, M. R. W. 1980. Structural evolution in the northern part of the Moine thrust belt, N.W. Scotland. Trans. R. Soc. Edinb. 71, 69-96. GARDINER, P. R. R. 1978. Is the Hercynian Front in Ireland a local feature? Nature, 271, 538-9. & MACCARTHY, 1. A. J. 1981. The Late Palaeozoic evolution of southern Ireland in the context of tectonic basins and their transatlantic significance. Mem. Can. Soc. Petrol. Geol. 7, 683-725. GmL, W. D. 1962. The Variscan Fold Belt in ireland. In: COE, K. (ed.) Some Aspects o f the Variscan Fold Belt, 49-64. Manchester University Press. GWINN, V. E. 1970. Kinematic patterns of lateral shortening, Valley and Ridge and Great Valley Provinces, Central Appalachians, south-central Pennsylvania. In: FISHER, G. W. et al. (eds) Studies o f Appalachian Geology--Central and Southern, 127-46. Wiley, New York. -

-

HANCOCK, P. L., DUNNE, W. M. & TRINGHAM, M. E. 1981. Variscan Structures in southwest Wales. ln" ZWART, H. J. & DORNSIEPEN,U. F. (eds) The Variscan Orogen in Europe. Geol. Mijnb. 60, 81-8. HARRIS, L. D. & MILICI, R. C. 1977. Characteristics of thin-skinned style of deformation in the southern Appalachians, and potential hydrocarbon traps. Prof Pap. U.S. geol. Surv. 1018, 40 pp. HOSSACK, J. R. 1979. The use of balanced crosssections in the calculation of orogenic contraction. J. geol. Soc. London, 136, 705-11. , NICKELSON,R. P. & GARTON, M. R. 1984. The geological section from the foreland up to the Jotun sheet in the Valdres Area, South Norway. In: GEE, D. G. & STEPHENS, M. B. (eds) The Caledonide Orogen--Scandinavia and related areas. Wiley, Chichester. In press. HUDSON, R. G. S. & PmLCOX, M. E. 1965. The Lower Carboniferous stratigraphy of the Buttevant area, Co. Cork. Proc. R. Irish Acad. 64B, 65-79. JACOBEEN, F. • KANES, W. H. 1974. Structure of Broadtop synclinorium and its implications for Appalachian structural style. Bull. Am. Ass. Petrol. Geol. 58, 362-75. KEEGAN, J. B. 1977. Late Devonian and Early Carboniferous miospores from the Galley HeadLeap Harbour region of southwest Ireland. Pollen Spores, 19, 545-73. MA3WHEWS, S. C. 1974. Exmoor thrust? Variscan Front? Proc. Ussher Soc. 3, 82-94. MACCARTHY, [. A. J., GARDINER, P. R. R. & HORNE, R. R. 1978. The lithostratigraphy of the Devonian-Early Carboniferous Succession in parts of Counties Cork and Waterford, Ireland. Bull. geol. Surv. Ireland, 2, 265-305. 1980. Facies changes in the Upper Devonian and Lower Carboniferous of south &

-

-

Thrust front o f the Irish Variscides Cork Ireland--a reassessment. Geol. Mijnb. 59, 65-77. MCCLAY, K. R. 1981. What is a thrust? What is a nappe? In: MCCLAY, K. R. PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 7-9. Blackwell Scientific Publications, Oxford. MORTON, W. H. 1965. The Carboniferous stratigraphy of the area north-west of Newmarket, Co. Cork, Ireland. Sci. Proc. R. Dubl. Soc. 2A, 47-64. NAYLOR, D. 1966. The Upper Devonian and Carboniferous geology of the Old Head of Kinsale, Co. Cork. Sci. Proc. R. Dubl. Soc. 2A, 229-49. 1969. Facies changes in Upper Devonian and Lower Carboniferous rocks of southern Ireland. Geol. J. 6, 307-28. 1975. Upper Devonian-Lower Carboniferous stratigraphy along the south coast of Dunmanus Bay, County Cork, ireland. Proc. R. Irish Acad. 75B, 317-37. 1978. A structural section across the Variscan fold belt, southwest Ireland. J. Earth Sci. R. Dubl. Soc. 1, 63-70. , JONES, P. C. & MATTHEWS, S. C. 1974. Facies relationships in the Upper Devonian-Lower Carboniferous of south-west Ireland and adjacent regions. Geol. J. 9, 77-96. -& SEVASTOPULO, G. D. 1979. The Hercynian 'Front' in Ireland. Krystalinikum, 14, 77-90. --, SEVASTOPULO, G. D., gLEEMAN, A. G. & REmLY, T. A. 1981. The Variscan Fold Belt in Ireland. In: ZWAR1, H. J. & DORNSIEPEN, U. F. (eds) The Variscan Orogen in Europe. Geol. Mijnb. 60, 49-66. NEVlLL, W. E. 1966. The geology of the North Cork (Kanturk) coalfield. Geol. Mag. 103, 423-31. NICKELSEN, R. P. 1966. Fossil distortion and penetrative rock deformation in the Appalachian Plateau, Pennsylvania. J. Geol. 74, 924-31. PENNEY, S. R. 1980. A new look at the Old Red Sandstone succession of the Comeragh Mountains, County Waterford. J. Earth Sci. R. Dubl. Soc. 3, 155-78. PHILCOX, M. E. 1964. Compartment deformation near Buttevant, County Cork, Ireland and its relation to the Variscan thrust front. Sci. Proc. R. Dubl. Soc. 2A, 1-11. PRICE, R. A. & MOUNTJOY, E. W. 1970. Geologic structure of the Canadian Rocky Mountains -

-

-

-

175

between Bow and Athabasca Rivers--a progress report. Spec. Pap. geol. Ass. Can. 6, 7-25. 1981. The Cordilleran foreland thrust and fold belt in the southern Canadian Rocky Mountains. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 427-48. Blackwell Scientific Publications, Oxford. REILLY, T. A. & GRAHAM, J. R. 1972. The historical and geological setting of the Glandore Mines, southwest County Cork. Bull. geol. Surv. Ireland, 1, 253-65. & 1976. The stratigraphy of the Roaringwater Bay area of south-west County Cork. Bull. geol. Surv. Ireland, 2, 1-13. ROEDER, D. GILBERT, O. E. & WITHERSPOON, W. D. 1978. Evolution and macroscopic structure of Valley and Ridge Thrust Belt, Tennesee and Virginia. Studies in Geology, 2, University of Tennesee, Department of Geological Science, Knoxville. 27 pp. SEVASTOPULO, G. D. 1982. Upper Carboniferous. In: HOLLAND, C. H. ( e d . ) A Geology o f Ireland, 173-89. Scottish Academic Press, Edinburgh. SHELFORD, P. H. 1963. The structure and relationships of the Namurian outcrop between Duntryleague, Co. Limerick and Dromlin, Co. Tipperary. Proc. R. Irish Acad. 62B, 255-66. SLEEMAN, A. G., REILLY, T. A. & HIGGS, K. 1978. Preliminary stratigraphy and palynology of five sections through the Old Head Sandstone and Kinsale Formations (Upper Devonian to Lower Carboniferous), on the west side of Cork Harbour. Bull. geol. Surv. Ireland, 2, 167-86. THOMPSON, R. J. 1981. The nature and significance of large 'blind' thrusts within the northern Rocky mountains of Canada. In: MCCLAV, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 449-62. Blackwell Scientific Publications, Oxford. WALLACE, P. 1968. The sub-Mesozoic palaeogeology and palaeogeography of northeastern France and the Straits of Dover. Palaeogeogr., Palaeoclim., Palaeocol. 4, 241-55. WALSH, P. T. 1968. The Old Red Sandstone west of Killarney, Co. Kerry, Ireland. Proc. R. Irish Acad. 66B, 9-26. WINGFIELD, R. T. R. 1968. The geology of Kenmare and Killarney. Ph.D. Thesis. University of Dublin. -

-

M. A. COOPER, D. COLLINS, M. FORD, F. X. MURPHY & P. M. TRAYNER, Department of Geology, University College, Cork, Ireland.

Does the Variscan front in Ireland follow a dextral shear zone? M. D. Max & J. P. Lefort SUMMARY: The line commonly taken in Ireland as marking the 'Variscan front' passes in a gentle arc from Dungarven Bay in the east to Dingle Bay in the west. This is the most convenient course for the northern margin of the Rheno-Hercynian zone through Ireland and marks a primary Variscan structural line. The line follows thrusts and steep faults in the west and a series of less well seen strike faults along the southern limb of a tight syncline in the east from Mallow to Dungarvan. Although the line is thus structurally composite, and stratigraphic evidence shows that it was locally of indirect influence during deposition of the Upper Palaeozoic sediments, both structural and geophysical evidence suggests that a dextral strike-slip component was associated with this line. Strike-slip movement along with N-S compression is shown by widespread E - W horizontal strain markers, strike faults and rotated fold traces which formed at the same time as the cleavage. Basement control of structures to the south of the line is not discernible while to the north rejuvenation of Lower Palaeozoic structures in the basement is common.

Discussion about the existence and nature of a Variscan front are of regional concern as this line has commonly been taken as marking the northern limit of the R h e n o - H e r c y n i a n zone of the Variscides (Read & Watson 1975). It has been commonly regarded as a narrow zone in Ireland where simple folding in thin Upper Palaeozoics to the north gives way to thrusting in thick Upper Palaeozoics in the south. Although this line is not everywhere a well defined feature, it is too continuous a structural line to be disregarded. Dunning (1966, 1980), for instance, refers to it as a distinct structural line in Britain and Ireland while noting that the re-evaluation of the R h e n o - H e r c y n i a n zone itself may lead to a widening of the meaning of the term, 'Variscan front'. Freshney & Taylor (1980) show that the line of the Variscan front in Pembrokeshire marks a major structure, while Badham (1982) draws attention to the probability of major strike-slip movement in the basement associated with the recognized thrusts and reverse faults. Between Mallow and Dungarvan Bay (Fig. 1) the D i n g l e - D u n g a r v a n line lies along the southern limb of a tight 80 km long near-horizontally plunging syncline where a number of strike faults are seen in the Upper Palaeozoic strata. The continuity and horizontal extent of this syncline is unique is southern Ireland since, to the south, often very tight triclinic folds are the rule (Max 1978) while to the north fold plunges in Upper Palaeozoic rocks are controlled by the position of major faults and broad domes cored by Lower Palaeozoic rocks (Hallissy 1928). It is not surprising that with these structural variations along strike the Variscan front in

places may be better defined by gravity and magnetic data than by any single geological criterion. The basement appears to be disrupted or stepped at depth and is seen by both a linear edge affect and disruptions of both gravity and magnetic patterns. In southern England where a steep dipping deep structure has been recognized from deep seismic reflection experiments (Kenolty et al. 1981), the associated juxtaposition has produced similar low amplitude geophysical anomalies to those seen in and around Ireland. Thus the D i n g l e - D u n g a r v a n line can be traced to Pembrokeshire and SE Britain, which implies that the recently suggested alternative course of the Variscan front along the Bristol Channel and the south Celtic Sea basin (Gardiner & Sheridan 1981), marks a subzone boundary within the R h e n o - H e r c y nian zone, rather than its margin.

Upper Palaeozoic sediments and basement in Ireland in relation to the Dingle-Dungarvan line Fossil, micro-fossil and pollen evidence, now allows precise regional correlation both laterally within the southern zone as well as with sediments to the north of the D i n g l e - D u n g a r van line (Naylor e t a l . 1981; Gardiner & McCarthy 1981). Stratigraphic correlations strongly suggest an extensive lateral continuity of U p p e r Palaeozoic sedimentation, which was only downfaulted to the south during sedimentation along the incipient D i n g l e - D u n g a r v a n line in its western end (Naylor et al. 1981, fig. 3). Indeed, there is a Lower Palaeozoic inlier at Mine H e a d (McCarthy 1978) to the south of 177

178

M. D. M a x

Dungarvan which is similar to the situation in Leinster to the north where Cambro-Ordovician Ribband Group sediments are unconformably overlain by Upper Palaeozoic basal conglomerates. The relationship between Lower and Upper Palaeozoic rocks thus appears to be similar on both sides of the Dingle-Dungarvan line with laterally continuous continental to shallow marine, ensialic Upper Palaeozoic sediments on both sides of the line sitting upon a Lower Palaeozoic basement. The Munster basin, containing Devonian and Carboniferous sediments (Gardiner & McCarthy 1981), appears not to have been greatly influenced by E - W basin-margin faulting during sedimentation but sedimentary zonation throughout Upper Palaeozoic times to the south (Naylor etal. 1980) consistantly follows E - W lines approximately parallel to the Dingle-Dungarvan line.

Structures to the north and south of the Dingle-Dungarvan line The original two-fold Hercynian structural N-S zonation recognized by Gill (1962) still applies

b

o

J. P. L e f o r t in his general sense but the line separating the two zones may not be a thrust front everywhere as he suggested. The 'cleavage folds' to the south of the line are generally upright or have steep-dipping axial planes except near the south side of Dingle Bay where there are shallowdipping axial planes associated with southdipping thrusts. Relatively small displacement, upright reverse and strike-slip faults are commonly seen. Folds usually lie with their axial traces lying more to the north of east than the trace of the Dingle-Dungarvan line, but folds near the line show a rotation of axial traces (Fig. la) into parallelism with the line. North of the Dingle-Dungarvan line, there appears to be a general but irregular decrease in the effects of deformation toward the north. The nearly isoclinal northerly overturned plunging folds seen in the Dingle peninsula give way to increasingly open folding until in the cliffs of north County Clare (Fig. la) bedding dips are nearly horizontal and towards the south over a broad area. The prominent regional cleavage also becomes less well developed towards the north and dies out along an arc about 75 km north of the Dingle-Dungarvan line (Max 1980). With increasing distance north

//

60

/

a

/

" DU

MHI

THIS AREA

FIG. 1. Location and general structural map. (a) Lower Palaeozoic inliers ornamented showing cleavage trace. Toothed lines, thrusts; heavy lines, thrusts; dashed line, position of eastern segment of Dingle-Dungarvan line. C, Cork Harbour; CC, County Clare; CK, Chair of Kildare; CZ, Variscan cleaved zone; D, Dublin; DA, Dunmore anticline; DB, Dingle Bay; DP, Dingle Peninsula; DU, Dungarvan Bay; G. Galtee Mountains; K, Killarney; LG, Leinster granites; M, Mallow; MHI, Mine Head inlier; R, Rosslare. (b) Diagram showing compression and shearing components.

Variscan f r o n t in Ireland of the line, deformation appears to be concentrated locally by structural controls within the pre-Carboniferous basement (Naylor et al. 1980). Over a broad area in the southern Devonian supercontinent and in virtually t h e whole of Ireland and most of Britain, however, Variscan deformation was locally manifested well north of the D i n g l e - D u n g a r v a n and south Pembrokeshire lines (Dewey 1982). Important strike faults such as that seen in the far SW (Graham & Reilly 1976) and in the Cork Harbour region (A. Sleeman pers. comm.), which are regarded as having important lateral movement, occur commonly throughout the southern Irish Variscides south of the D i n g l e - D u n g a r v a n line. Dextral rather than sinistral movement is more common. Naylor & Sevastopulo (1979) point out that strike faulting may be far more widespread than has been previously supposed and a strike-slip component in the region between Killarney and Mallow associated with the Dingle-Dungarvan line (Naylor et al. 1980) suggests that the engine of dislocation is lateral movement and it is likely that thrusting on the line is rooted in E - W transcurrent shear zones in the subjacent basement. If the density of strike faulting proves on a regional scale to be about the same as that seen in detail locally, and this is not firmly established, then the total lateral displacement inherent in even a moderate or late strike-slip movement would be on the order of tens of kilometres. The major zone of thrusting and reverse faulting between Dingle Bay and Mallow turns

179

off the Dingle-Dungarvan line, or continues in a north-easterly course in the vicinity of Mallow, where thrusts occur both south and north of the Galtee Mountains (Fig. la). Thrusting associated with the Dunmore anticline further to the NE can be followed along a structural high through Kildare (Gardiner & Sheridan 1981) into the E - W zone of Variscan cleavage passing into the Irish Sea north of Dublin. This 'S'-shaped line is approximately coincident with the south Ireland lineament of Gardiner (1975). The overall pattern of the Variscan structural trends in Ireland is seen by us as being of the greatest significance in determining the precise nature of the Dingle-Dungarvan line. The geometry is strongly reminscent in pattern to folding forming in the vicinity of a major compressional shear zone (Fig. lb). The extent of this coherent structural pattern shows that the regional Variscan strain was manifested over a broad area in Ireland. This implies a deeprooted structural framework.

Tracing the Dingle-Dungarvan line The D i n g l e - D u n g a r v a n line has a magnetic and gravity expression on land which can be traced offshore both to the west and east.

The magnetic data The magnetics in Fig. 2 have been smoothed from newly compiled magnetic data (Max,

Fie. 2. Magnetic map. 50 nT contour intervals (total field) except for 100 nT in western end. Dashed lines indicate data source area boundaries. Projected course of Variscan front shown by solid toothed line.

M. D. Max & J. P. Lefort

180

pair is generated by an only slightly disrupted anomaly on either side of the E - W feature. The magnetics are also useful in locating the possible E - W extension of the line to the west of Ireland (Lefort & Max 1984). The E - W lineament, which is the westward continuation of the Dingle-Dungarvan line, is offset by large post-Hercynian 130 ~ dextral shear zones of probable Permo-Triassic age. On the western side of the Porcupine Bank the N52~ latitude E - W lineament represents the westernmost extension of the line. This course for the Variscan front passes into the vicinity of the eastern termination of the Gibbs fracture zone. Recognition by us of the line as a major shear zone supports the view of Cherkis, Flemming & Massingill (1973) that the vertical oceanic fracture zone was generated as an extension of the Variscan front and implies that the Variscan front is more likely to be a steep shear zone penetrating the crust than a shallow dipping structure confined by thrust tectonics to very high crustal levels.

Inamdar & McIntyre 1982). There are abrupt flexures Of the magnetic lines in the approximate positions of the Dingle-Dungarvan line and the Variscan front in Pembroke (Fig. 2). West of south Wales a clear E - W trending paired magnetic high and low suggest the probable extension of the Variscan front. All the transverse magnetic structures display Caledonian trends between N50 ~ and N60~ they stop at the vicinity of the magnetic trace of the line. The most prominent of these transverse structures is the magnetic high associated with the Irish Sea land mass: the northernmost part of the anomaly extends through the magnetic high of the Lleyn Peninsula, while in the south it is the expression of the Rosslare complex. This anomaly is probably related to major occurrences of mafic bodies at depth which are locally seen at the surface as gneisses and intrustive rocks. South of the line the same magnetic high (the Kinsale ridge) appears to follow the Irish coastline offshore. There is an apparent dextral offset of about 30 km between the Rosslare and Kinsale ridges. Between the Lleyn Peninsula and south Wales, a magnetic low is probably related to the existence of the Cardigan Bay basin. This basin anomaly may be extended in the south by a weak magnetic low which is related to the Mesozoic and Cenozoic Celtic Sea basin (Whitbread 1975). Again, there appears to be a dextral offset between the magnetic lows of about 30 kin. This suggestion of a 30 km dextral offset along the line from the magnetic pattern in the Irish Sea would, of course, be a minimum estimate based on the premise that the anomaly

The gravity data The gravity compilation (Fig. 3) incorporates the Bouger anomaly map of southern Great Britain and the Irish Sea (Maroof 1974), the Bouger anomaly map of Ireland (1974), free air gravity data (Handley 1971) data from the sea area around Rosslare (smoothed by Blundell 1975) and data of Haworth & Jacoby (1982). The magnetic highs of the Lleyn and Rosslare ridges are broadly superimposed with discontinuous gravity highs, except on the north54 ~

_1~

',3o~C

Go

-53 ~

,

.

~,~

.

.

.

.

k//Z/~/~\\k\.

.

,-,

\

t( /

~

yL/r

.....

o

?" "-,..<~' ..., ,," yo ~ ; ~ S -52 ~

16 9

1;"

14o

1; ~

12o

~io

i0o

9~

8~

70

6~

~~

FIG. 3 Gravity map. Solid lines at 10 regal, dashed where projected. Projected course of Variscan

front shown by solid, toothed line. Features: C, Celtic Sea basin; I, Irish Sea basin; K, Kinsale ridge; R, Rosslare ridge; A, Anglesey; L, Llynn Peninsula; P, Pembrokeshire. Irish mainland after Murphy (1974).

Variscan front in Ireland western tip of the Cardigan Bay basin where there probably is extensive Mesozoic cover. This overlaying of anomalies reinforces our magnetic interpretation of probable mafic rocks at depth. The Kindale ridge shows a similar gravity high superimposed on its magnetic high. There appears to be the same, but a better defined, dextral offset of the NW faulted margins of the basins. The course of the Variscan front between Dungarvan and Pembrokeshire also marks a change in both the magnetic and gravity patterns to the north and south, with a deepening of the basement surface to the south. Onshore, in Ireland, two major gravity lows are separated by the E - W line. In the east a large low trending N60 ~ reflects the position of the Leinster granite and probable subjacent granitic rocks. Near Dingle Bay another gravity low is most likely related to an unseen granite (Howard 1975) which preliminary radioisotope data suggest is also of Caledonian (c. 400 Ma) age, there may be the same 30 km dextral offset as is seen in the offshore data (Figs 2 & 3). Both the magnetic and gravity interpretations suggest that the Dingle-Dungarvan line and its continuation from the Variscan front in Pembrokeshire acted as a dextral transcurrent fault.

The Variscan front in a European context The Dingle-Dungarvan line appears to be a primary shear zone and the only Variscan structural line of this importance in Ireland. The line passes into Britain and eastwards on to the continent where it forms the northern margin to the Rheno-Hercynian zone. Although the Irish Variscides are a single structural zone, Dunning (1966) shows that the British Variscides are composed of more than one structural compartment. The area from the Variscan front in the Pembrokeshire area to a Bristol Channel line, which may be marked by the Exmoor and Cannington thrusts (Gardiner & Sheridan 1981), comprises the northermost zone. This zone broadens to the west (Fig. 5) where it is represented by the whole of the Irish Variscides. The southern margin of this zone passes to the SE of Ireland (Gardiner & Sheridan 1981). Briefly, SW Britain to the south of the Bristol Channel is quite different from the Irish Variscides: (1) the Irish Variscides have no internal structural zonation as do the British Variscides; (2) the Irish Variscides comprise usually upright structures; SW Britain has a more complex tectonic history with horizontal-directed shear features (major thrusting) common over

181

broad areas (Shackleton et al. 1982); (3) the most important contrast, however, is the development of granites. No Variscides granites are seen in the Irish Variscides; SW Britain and its western continuation on to the immediately adjacent continental shelf, is characteristically a granite terrane. As the general geological situation is so different, it would be unwise to project structural syntheses directly to the Irish Variscides from SW Britain. Although there may well be major horizontal detachment thrusts below southern Ireland (Cooper et al., this volume) similar to those seen in SW Britain, these have not been mapped, but only modelled. If a shallow structural model is proven through deep geophysical investigations, these Irish structures would be composite, and carry an E - W wrench component.

Dextral shear of the Variscan front in the framework of the North Atlantic Major transatlantic shear zones showing clear geophysical signatures trend E - W between Europe, Africa and Amerca. These zones were major dextral shears during late or middle Carboniferous times, but some might be as young as Permian or Lower Triassic especially toward the south. Some thrusting is associated with them, suggesting that a compressional shear regime acted over a very large region at this time. Two main shear zones have been delineated between western Europe and the Grand Banks of Newfoundland (Lefort & Haworth 1978); the Armorican shear zone and the Cobequid-Chedabucto fault zone (Fig. 4). Between New England and Morocco, the Bay of Maine shear zone is parallel to the south Atlas shear zone which trends between New York and southern Morocco (Lefort & Haworth 1983). The Variscan front is parallel to these other shear zones and we regard the Irish and British Variscan front as a shear zone of this set; other n'finor faults may occur to the NW (Dewey 1982). At the same time that these dextral shears were active large sinistral movements may have occurred along the Appalachian Axis (Van der Voo 1981) and perhaps in Scotland (Van der Voo & Scotese 1980). Some small sinistral displacements have been recognized in the field, but this major fault set, if real is probably now buried beneath the numerous Appalachian thrusts of Alleghenian age. The main sinistral displacement actually known acted beneath the Appalachian basin, and east of New England in America (Fig. 4) and along the Zemmour fault in Africa. The general pattern delineated by these shear

182

M . D . M a x & J. P. L e f o r t

FIG. 4. Atlantic framework of strike-slip faults. Locations: 1, Great Glen Fault; 2, east of Newfoundland; 3, east of New England; 4, Appalachian basin; 5, Armorican shear zone; 6, Cobequid-Chedabucto fault; 7, Bay of Maine shear zone; 8, Atlas shear zone; 9, Zemmour fault; 10, Variscan front of Ireland and Britain. Features: 1, major transcurrent faults; 2, Irish and British Variscan front; 3, West Africa craton. zones suggest that the West A f r i c a n craton acted like an i n d e n t e r (Lefort & V a n der V o o 1981). T h e i n d e n t i n g process w o u l d have b e e n responsible for the uplift of the A n t i - A t l a s o r o g e n y of the n o r t h e r n b o u n d a r y of the craton and for the O u g a r t a o r o g e n y on its eastern side. A t the s a m e time m o v e m e n t and rifting o c c u r r e d in the Maritimes as a r e s p o n s e to the overall stress. All those events are typical of an assymetric i n d e n t e r ; the t h e o r e t i c a l m o d e l first suggested by M o l n a r & T a p p o n n i e r (1977) can

a p p a r e n t l y be used for the d e v e l o p m e n t of the A t l a n t i c Variscides. A c o m p r e s s i o n a l shear zone along the Variscan front in s o u t h e r n Ireland and P e m b r o k e s h i r e w o u l d fit very well with the g e n e r a l transatlantic p a t t e r n for U p p e r C a r b o n i f e r o u s time.

ACKNOWLEDGMENT: This paper is published with the permission of the Director of the Geological Survey of Ireland.

References BADHAM, J. P. N. 1982. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 493-504. BAILEY, R. J. 1975. The geology of the Irish continental margin and some comparison with the offshore Eastern Canada. In: YORATH et al. (eds) Canada's Continental Margins and Offshore Petroleum Exploration. Mere. Can. Soc. Petrol. Geol. 4, 313-40.

BLUNDELL,D. J. 1975. The geology of the Celtic Sea and southwestern Approaches. In: YORATH et al. (eds) Canada's Continental Margins and Offshore Petroleum Exploration. Mere. Can. Soc. Petrol. Geol. 4, 341-52. CHERKIS, N. Z . FLEMMING,H. S. &; MASSINGILL,J. V. 1973. Is the Gibbs Fracture Zone a westward projection of the Hercynian Front into North America? Nature, 245, 113-5.

Variscan front in Ireland DEWEY, J. F. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. London, 139, 371-412. DUNNING, F. W. 1966. Tectonic Map of Great Britain and Northern Ireland. Institute of Geological Sciences, England. 1980. United Kingdom, geotectonic position. In: BORDAS and 26th Int. Geolog. Congress (eds). Geology of the European Countries. 331-7. Graham & Trotman, London. FRESHNEY, E. R. & TAYLOR, R. 1980. The Variscides of southwest Britain. In: BORDAS and 26th Int. Geolog. Congress (eds). Geology of the European Countries, 379-87..Graham & Trotman, London. GARDINER, P. R. R. 1975. Plate tectonics and the southern Irish Caledonides. Proc. R. Dubl. Soc. 5, 385-96. & MACCARTHY,I. A. J. 1981. The late Palaeozoic evolution of southern Ireland in the context of tectonic basins and their transatlantic significance. In: KERR, J. W. & FERGUSSON,A. J. (eds) Geology of the North Atlantic Borderlands. Mere. 9 Can. Soc. Petrol. Geol. 7, 683-725. & SHERIDAN, D. J. R. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan Front. J. struct. Geol. 3, 317-31. GILL, W. D. 1962. The Variscan Fold Belt in Ireland. In: COL, K. (ed.) Some Aspects of the Variscan Fold Belt, 49-64. Manchester University Press. GRAHAM, J. R. & REILLEY, T. A. 1976. The stratigraphy of the area around Clonakilty Bay, south County Cork. Proc. R. Ir. Acad. 76, 379-91. HALL1SSY, T. 1928. Geological Map of Ireland. HANDLEY, R. L. 1971. A geophysical study of the Porcupine Seabight. Unpublished Ph.D. thesis. University of Birmingham, 179 pp. HAWORTH, R. J. & JACOm, R. D. 1982. Geophysical correlation between the geological zonation of Newfoundland and the British Isles. In: HATCHER, R. D., WILLIAMS,H. & ZIETZ, 2. (eds) Contributions to the Tectonics and Geophysics of Mountain Chains. Spec. Pap. geol. Soc. Am. 158, 25-32. HOWARD, D. W. 1975. Deep seated intrusives in Co. Kerry. Proc. R. Ir. Acad. 75, 173-83. KENOLTY, N., CHADWICK,R. A., BLUNDELL, D. J. & BACON, M. 1981. Deep seismic reflection survey across the Variscan Front of southern England. Nature, 193, 451-3. LEFORT, J. P. & HAWORTH, R. T. 1978. Geophysical study of basement fractures on the western European and Eastern Canadian shelves: transatlantic correlations and late Hercynian movements. Can. J. Earth Sci. 15, 397-404. & HAWORTH, R. J. 1984. Geophysical correlation between basement features in Northwest Africa and North America, and their control over structural evolution. Soc. geol( mineral. Bretagne (in press).

183

& MAX, M. D. 1984. Development of the Porcupine Seabight: the direct relationship between early oceanic and continental structures. J. geol. Soc. London (in press). & VAN DER VOO, R. 1981. A kinematic model for the collision and complete suturing, between Gordwanaland and Laurussia in the Carboniferous. J. Geol. 89, 537-50. MACCARTHY, I. A. J. 1978. The lithostratigraphy of the Devonian-Early Carboniferous succession in parts of County Cork and Waterford, Ireland. Bull. geol. Surv. lrel. 2, 265-305. MAROOF, S. I. 1974. A Bouguer anomaly map of southern Great Britain and the Irish Sea. J. geol. Soc. London, 130, 471-4. MAX, M. D. 1978. Tectonic control of offshore sedimentary basins to the north and west of Ireland. J. Petrol. Geol. l, 103-10. -1980. Tectonic map of Ireland. In: New Atlas of Ireland. Royal Irish Academy. -

-

--,

INAMDAR, D. D. & MCINTYRE, T. 1982.

Compilation of magnetic map: The Irish Continental shelf and adjacent areas. Rep. geol. Surv. Irel. 82/2. MOLNAR, P. ~k TAPPONIER, P. 1977. Relation of the tectonics of Eastern China to the India-Eurasia collision. Application of ship-line field theory to large scale continental tectonics. Geology, 5, 212-6. MURPHY, T. 1974. Gravity anomaly map of Ireland. Communs. Dublin Inst. advd. Stud. D, 32. NAYLOR, D., PHILLIPS, W. E. A., SEVASTOPULO, G. D. & SYNGE, F. M. 1980. Ireland. In: BORDAS and 26th Int. Geolog. Congress (eds). Geology of the European Countries, 133-81. Graham & Trotman, London. -& SEVASTOPULO, G. D. 1979. The Hercynian front in Ireland. Krystalinikum, 14, 77-90. --, , SLEEMAN, A. G. & REmLY, T. A. 1981. The Variscan fold belt in Ireland. Geologie Mijnb. 60, 49-66. READ, H. H. & WATSON, J. 1975. Introduction to Geology Vol. 2. Earth History. Macmillan, London. SHACKLETON, R. M., RIES, A. C. & COWARD, M. P. 1982. An interpretation of the Variscan structures in S.W. England. J. geol. Soc. London, 139, 533-41. VAN DERVOO, R. 1981. The position of Great Britain with respect to the North American Craton in the Palaeozoic (abstract). Trans. Am. geophys. Un. 62, 264. & SCOTESE, C. R. 1980. Great Glen Fault. 2000 km sinistral displacement during the Carboniferous (abstract). Trans. Am. geophys. Un. 61, 220. WHITBREAD, D. R. 1975. Geology and petroleum possibilities west of the United Kingdom. In: WOODLAND, A. W. (ed.) Petroleum and the Continental Shelf ofNorth West Europe, 1, 45-73. Applied Science Publishers, Barking, Essex.

M. D. MAX, Geological Survey of Ireland, 14 Hume Street, Dublin 2, Ireland. J. P. LEFORT, CNRS Centre Armoricain L'Etude Structurals des Socles, Institut de G6ologie de l'Universit6 de Rennes I, Campus de Beaulieu, 35042 Rennes Cedex, France.

Variscan structures in the Upper Palaeozoic rocks of west central Ireland D. W. Coller SUMMARY: The western half of central Ireland is part of a low strain zone on the northern margin of the Variscan orogen in which the structures in the Upper Palaeozoic 'cover' rocks appear largely to be controlled by ENE-WSW and NNE-SSW trending structures in the Caledonian, and possibly the pre-Caledonian basement. Preceding the main ductile phase of Variscan deformation, there was an important period of mid-Dinantian faulting which is closely associated with major base metal mineralization. New structural evidence suggests that the main movement on these faults involved an important component of dextral transcurrent shear. The main ductile deformation, between post-Westphalian and pre-Upper Permian in age, is characterized by kilometric scale open folds and mainly ENE-WSW trending vertical transcurrent ductile-brittle shear zones. The model of deformation is considered to be one of regional, heterogeneous transpression where the main compression, or pure shear component was approximately N-S, and the simple shear component was an ENE-WSW dextral shear. A major NNE-SSW trending sinistral shear zone, the Fergus shear zone, appears to be the western limit of the dextral shear component. The post-Caledonian cover sequence in central Ireland lies unconformably on Ordovician and Silurian metasediments and consists of Upper Old Red Sandstone successions (Upper Devonian to Tournasian age), of dominantly fluvial clastic sediments, generally less than 300 m thick (Emo 1978; Holland 1981). These pass up through a thin marine transgressive sequence into thick marine carbonates, up to 1500 m thick in the River Shannon area east of Foynes, of Lower Carboniferous age (Sevastopulo 1981). These limestones are overlain in the western part of the area by thick successions of Namurian shales and sandstones comprising the Shannon trough (Sevastopulo 1981). The pattern of Caledonian faults and ductile structures, exposed in inliers in the cores of Variscan anticlines (Fig. 1), suggests that they formed in a broad zone of strike-parallel dextral simple shear, the axis of which is the proposed trace of the Iapetus suture which extends eastnorth-eastwards, to north-eastwards through Silvermines and Navan (Phillips, Stillman & Murphy 1976; Phillips, Flegg & Anderson 1979; Sanderson et al. 1980). There is a close parallelism in strike trend between the principal Variscan structures in the Upper Palaeozoic cover rocks and the Caledonian structures, particularly the cleavage, in the basement (Fig. 1). The period of Variscan deformation, as for the Caledonian in central Ireland, appears to be characterized by an ENE-trending dextral shear regime approximately parallel to the Caledonian structural trend. This paper describes Variscan structures of the western half of central Ireland which corn-

prises part of a relatively low strain orogenic margin, in which a N - S compression and an E - W to E N E - W S W dextral transcurrent shear couple are resolved as the main deformation components. The tectonic regime is characterized by ductile-brittle transcurrent (strike-slip) shear zones, major open folds and heterogeneous vertical cleavage. The age of the main ductile deformation is between post-Westphalian B and pre-Upper Permian. The chronology and pattern of the minor structures associated with the Silvermines fault, a major strike fault and shear zone, suggest that the period of mid-Dinantian faulting, which is closely associated with base metal mineralization, involved a significant E - W dextral shear component. This early deformation may correlate with that resulting in similar style structures described in the Dinantian rocks of the Pennines, England (Arthurton, this volume). A major N N E trending sinistral shear zone, the Fergus shear zone, appears to mark the western boundary of the central Ireland zone transpression regime (Fig. 1). To the west of the Fergus shear zone, the pattern of parallel trains of minor folds in the Namurian sandstones and shales suggest they are the result of an approximately N - S compression, with little evidence of major shearing along discrete zones as seen to the east, although more detailed mapping of the structures in this area is required. East of the Fergus shear zone, linear shear zones bound less deformed tectonic units, such as the largest defined unit, the Nenagh block (Fig. 1), which is characterized by large-scale open folds, no cleavage and poorly developed vein systems.

185

186

D. W. Coller

FIG. 1. Variscan tectonic units in west central Ireland. Stippled areas are Lower Palaeozoic inliers showing generalized Caledonian cleavage (dashed lines). White is undifferentiated Old Red Sandstone and Carboniferous. Source of data for the Caledonian mainly Sanderson et al. (1980), also A. Weir, J. Archer, A. Flegg. Minor structures are best developed in the widely exposed Dinantian limestone lithologies. A chronological summary of these structures and their relationship to major structures is presented in Fig. 2. Most of the vein structures are not described in detail in this paper, but are included to illustrate the sequence of mineralized fracture development.

Major cover faults and their early movement history Large, steep faults, commonly 3 0 - 5 0 km in length, are important tectonic elements of central Ireland (Figs 3 & 4). Most of the major faults trend E N E - W S W and are parallel to the principal Caledonian structures in the base-

FIG. 2. Chronology of Variscan structures and carbonate veins. En echelon veins are vertical and related to minor shears; intrabed veins are shallow dipping veins within bed units; bedding surface veins separate bed units; principal extension veins are vertical, planar veins.

Variscan structures in west central Ireland ment. Significant changes in the Caledonian strike to NNE-SSW occur in the Knockshigowna inlier (Fig. 4) and in the N W margin of the Slieve Aughty inlier (Sanderson et al. 1980), where the main Variscan structural trend, in the form of the Fairy Hill fault (Brfick 1982) and the Fergus shear zone respectively (Figs 3 & 4), is also NNE-SSW, suggesting there is a basement control to the faulting and shearing in the Upper Palaeozoic cover. The principal movements on the faults occurred durmg the main Variscan deformation phase or later; the faults generally being the axial region of high strain transcurrent shear zones, as in the case of Silvermines and several major faults in the Slieve Aughty inlier (see next section). There is evidence, however, as in the case of the Silvermines fault zone and in some of the faults in the Slieve Aughty inlier, that there was an important period of pre-Westphalian fault movement which influenced sedimentation. The main faults commonly have a vertical separation of several hundred metres, in many cases throwing Lower Palaeozoic greywackes and slates, and Old Red Sandstone against Lower Carboniferous limestones, e.g. the Sil-

187

vermines fault, Tynagh fault, and the main faults in the Slieve Aughty inlier. Upper Devonian and Tournasian (Lower Dinantian) clastic sedimentation appears to have been locally controlled by some of these faults as in Slieve Aughty (Emo & Grennan 1982), the Galty Mountains (Curruthers pets. comm.) and the Devils Bit Mountains (Feehan 1980), whilst there is also growing evidence, particularly from studies of major base metal deposits hosted in limestones, that many of these major faults were also active later in the Dinantian, for example the Tynagh fault (Moore 1975), the Silvermines fault (Taylor & Andrew 1978; Coller in prep.), and faults in the North Curlew mountains (Sevastopulo 1981). Structures associated with the Silvermines fault zone, such as clockwise rotation of cleavage and minor transcurrent shears, suggest that the main fault movements were dextral transcurrent. Also, the overall geometric pattern and the net displacement on the numerous faults at Silvermines is, as has been stated (Rhoden 1958; Taylor & Andrew 1978), compatable with an ENE trending dextral transcurrent shear couple. The large net vertical displace-

FIG. 3. Map of west central Ireland showing major geophysical features and Variscan cover structures. Dash-dot lines, aeromagnetic lineaments; thin lines (with barbs), lineaments and anomalies from Bouguer anomaly data. Thick solid lines, faults; thick dashed lines, shear zones; thin lines, generalized cleavage and fold trends. The stippled areas are Lower Palaeozoic inliers. All geophysical lineaments and anomalies were interpreted from aeromagnetic maps of Ireland (Williams 1981) and the Gravity Anomaly Map of Ireland (Dublin Institute of Advanced Studies 1974). Faults were compiled from numerous sources.

188

D. W. Coller

FIG. 4. Structural map of west central Ireland showing Variscan folds, faults and shear zones. The stippled areas are Lower Palaeozoic inliers. FSZ, Fergus shear zone; TF, Tynagh fault; SF, Silvermines fault; GF, Gortdrum fault; QSZ, Quin shear zone; FHF, Fairy Hill fault; SA, Slieve Aughty; D, Devilsbit Mountains; K, Knockshigowna. m e n t on m a n y of the main E N E - W S W faults is due to a large n o r m a l (extensional), as well as a horizontal ( t r a n s c u r r e n t ) m o v e m e n t c o m p o nent, which is c o n s i d e r e d to be the result of t r a n s c u r r e n t shearing in an overall extensional r e g i m e (transextension). Also, local z o n e s of extension a n d n o r m a l fault m o v e m e n t s are

common features of t r a n s c u r r e n t zones ( T c h a l e n k o & A m b r a s e y s 1970; M o o r e 1979). F r o m detailed study of the faults and r e l a t e d structures at Silvermines it has b e e n possible to reconstruct the likely evolution of the fault complex, s u m m a r i z e d in Fig. 5, a n d show that the early, as well as the m a i n fault m o v e m e n t s

FIG. 5. Proposed sequence of fault development at the eastern termination of the Silvermines fault, at Silvermines. Half-arrows indicate the sense of shear; full arrows, extensional (normal) faulting; and barbs, reverse faulting, a represents the regional compressive stress during the main deformation. Stages: (a) earliest dextral shear movements possibly generating N E - S W open folds; (b) initiation and early stages of propagation, by extension and dextral shear, of en echelon WNW ('oblique extension faults'). Formation of the main orebodies, B zone and G zone (stippled). Initiation of a principal ENE fault (S), en echelon to the main Silvermines fault; (c) continued development of WNW faults and associated slumping in a dextral shear regime. Development of subordinate extensional (normal), and Riedel shear faults; (d) reactivation of faults during the main deformation. Some of the fault data are based on Taylor & Andrew (1978).

Variscan structures in west central Ireland during the main deformation phase, were largely dextral transcurrent. The pattern, and evolution of early faulting at Silvermines is considered to be closely analogous to that observed in a simpler, smaller-scale setting at the termination of minor transcurrent fault systems (Fig. 6), where WNW trending faults at Silvermines are correlated with the series of en echelon fractures. Fundamental in drawing this analogy is the relationship between the mineralization and faulting at Silvermines. Mineral zoning for the two main syngenetic ore bodies at Silvermines, indicates that mineralization was influenced by W N W - E S E faults (Taylor & Andrew 1978), and more specifically that zones close to the centres of these faults provided channel-ways for mineralizing fluids (Taylor pers. comm.). Fault activity also controlled the development of major soft sediment folding, close to the time of mineralization (Coller in prep.). The chronology of fracture development in the small-scale transcurrent fault termination, a dilation area (Fig. 6), is recorded by the growth of crystal fibres, such that they show a series of en echelon fractures which propagated incrementally by extension (normal faulting) and synthetic dextral shearing (transcurrent faulting). The pattern of faults at Silvermines at the time of mineralization (Lower Dinantian), suggests the area was at the eastern termination of the Silvermines fault zone, where early W N W - E N E faults developed in a shear generated dilation zone (Fig. 5a-c), analogous to the en echelon fractures in Fig. 6. Detailed study of the WN W - E S E fault which controls the B zone orebody shows that the central part of the fault,

\\\\\\

\@\\\'\ \\ ,,

x'x.

~pn

\\\\

~

\ x\ \\ .. \ \" ,

i

Fro. 6. Diagram showing the geometry of en echelon, vein-filled fractures in the termination of a dextral fault system. The orientations of the fractures and mineral fibres are based on field measurements of several, mirror image sinistral fault systems developed in Namurian sandstones at KiP kee, Co. Clare. The scale is approximate. The insets show detail of the traces of mineral fibres which record oblique opening of the fractures by extension and shear.

189

in the region of the main mineralizing fluid channels, has the maximum normal and trans, current displacement, and displacement decreases in both directions along the fault, suggesting that the fault initiated at this central point and propagated by extension and shear, as indicated by alternating horizontal and downdip slickensides. The style of propagation of the WNW faults, by extension and shear, is analogous to the history of growth recorded by the mineral fibres in the en echelon fractures in the smaller-scale fault termination (compare Figs 5 & 6). This style of faulting is referred to here as 'oblique extensional faulting', the faults lying between the orientations o f maximum extension and maximum shearing (synthetic Riedel shears) for this simple regime. The main inference from this detailed relationship is that this early faulting developed due to a dextral shear movement along the main Silvermines fault zone, probably by reactivation of a basement structure, possibly related to the Iapetus suture. There is more detailed and wider confirmation at Navan, in the eastern half of central Ireland, for major early Variscan fault movements of approximately the same age, in the form of a major unconformity near the base of the Chadian (Dinantian) which truncates a major fault (Andrew & Aston 1982) whose early movement is also transcurrent (Phillips pets. comm.). A wider correlation of this early fault activity could perhaps be made between central Ireland and northern England, where slightly later Chadian (Dinantian) age dextral transcurrent shearing in the Ribblesdale fold belt, along a similar strike trend has been described by Arthurton (this volume). If this early proposed dextral shearing is an important regional strain, then the E N E - W S W dextral shear couple may explain the common occurrence in central Ireland of N W - S E trending faults, many of them normal faults, which would have formed parallel to the planes of maximum extension. Parallel to these faults is a N W - S E magnetic lineament extending between Foynes and Kilmallock (Fig. 3) which most likely represents a major dyke related to the nearby Visean volcanic complex of Limerick.

The main Variscan deformation The tectonics related to the main ductile phase of Variscan deformation can be divided in the west of central Ireland into three main areas: the west Clare block, the Fergus shear zone and the main central Ireland zone (Fig. 1).

190

D. W. Coller

The west Clare block

This block, west of the Fergus shear zone, consists largely of Namurian sandstone and shale lithologies in the south, and Upper Dinantian (Visean) limestones in the north. The main structures are trains of subparallel E - W to E N E - W S W trending, gently plunging minor folds with wavelengths up to 500 m. In profile the folds are predominently angular, and pass northwards from anticlinal box folds with more rounded synclines, around the River Shannon estuary, into progressively more open folds, commonly widely spaced monoclines. In the most northerly part of the block, the Upper Dinantian limestones are gently undulating with dips rarely exceeding 10 ~ In the south, the folds are generally associated with a steep, approximately axial planar cleavage which is preferentially developed in the steep limb and hinge regions of folds. Northwards, this cleavage becomes progressively less well developed, until most of the Upper Dinantian limestones in the very north are uncleaved. The formation of the main folds in this block appear to relate to a simple N - S pure shear compression, with little evidence for any significant shear component to the deformation as obvious further east. In addition, a later N - S compression is deduced from extensive, vertical, conjugate sets of en echelon vein arrays and minor shear fractures which post-date the folds. In this family of brittle structures are N N E - S S W trending minor sinistral faults whose main movement may relate to shearing on the Fergus shear zone to the east.

magnetic lineament along its northern mapped limit, and a steep gravity gradient along its southern extent (Fig. 3). This deeper level of the zone is evident to some extent in the Lower Palaeozoic rocks in the N W of the Slieve Aughty inlier where there is a marked change in strike of the steep Caledonian cleavage and folds (Emo 1978) parallel to the shear zone. This NNE trend in the Caledonian structure is also seen in the Stokestown inlier (Morris 1979) and may be a continuation of the same zone. Another feature of the magnetic pattern (Fig. 3) is the abrupt termination eastwards against the Fergus shear zone of, what appears to be the large Galway granite complex, further suggesting that this shear zone in the cover is possibly a structure in the basement. South of the Slieve Aughty inlier the Fergus shear zone broadens, and appears to terminate in the area of the River Shannon around Foynes (Figs 3 & 4). This southern termination of the shear zone may be related to the large positive gravity anomaly centred on Foynes (Fig. 4), which may represent an area of more basic material in the basement which has acted as a rigid block (Phillips pers. comm.). A summary of the structures comprising the southern termination of the shear zone is shown in Fig. 8(d). The pattern of reactivated vertical, extension veins is consistent with a N N E - S S W sinistral shear couple, whilst possible synthetic, N N E - S S W splay faults may explain the anticlockwise swings in the axial traces of the folds just west of Foynes, south of the River Shannon. The central Ireland zone

The Fergus shear zone

At the eastern margin of the west Clare block, E - W trending folds are tightened, and rotated anticlockwise into a N N E - S S W orientation by either syn- or post-fold sinistral movement along the Fergus shear zone (Fig. 4). A minimum shear strain estimate of y = 0.9 based on the anticlockwise rotation of an early, regional E S E - W N W set of extension veins, gives a sinistral displacement across the shear zone of the order of 10 km. The Fergus shear zone is a broad zone of shear strain, varying between 5 and 15 km wide, tapering north-eastwards, and is at least 70 km in length extending from the River Shannon at Foynes, northwards along the NW margin of the Slieve Aughty Mountains and may continue as far as the Stokestown inlier (Fig. 1). A deeper expression of the Fergus shear zone is suggested by a strong aero-

The central Ireland zone is characterized by two sets of regional structures: major, kilometric scale folds and steep, ductile-brittle shear zones.

Kilometric scale folds A series of upright, open, E N E trending folds, with 5 - 1 0 km wavelengths, have gentle plunge culminations and depressions. Lower Palaeozoic metasediments are frequently exposed in the cores of anticlines, forming the main inliers of the central Ireland zone (Fig. 4). A single steep cleavage is heterogeneously developed in the Dinantian limestones, often axial planar to minor folds, but commonly transecting the axial traces of the kilometric scale folds. Strike swings between N - S and E N E - W S W and variation in intensity of the cleavage is largely related to the main shear zone development (see s e c t i o n on ductile-

191

Variscan structures in west central Ireland

brittle shear zones). Similarly, minor folds are often oblique to the major folds and are more commonly related to ductile shear zones (see Fig. 8a). The Nenagh block (Fig. 1) is interpreted as a relatively undeformed, and tectonically stable area within the central zone. The main stucture is a broad, open syncline (Briick 1982), whose axial trace strikes N N E - S S W , oblique to the regional trend ( E N E - W S W), but parallel to the main Variscan faults and the principal Caledonian structures exposed in the Knockshigowna inlier (Figs 1 & 5). The limestones, which mainly comprise the Nenagh block, are gently dipping, uncleaved and have poorly developed vein systems, representing a low strain area when compared to much of the central Ireland zone. This relatively low strain area is coincident with a large, broadly N N E - S S W trending negative gravity anomaly (Fig. 3), which is thought to be an underlying Caledonian intermediate pluton at approximately 3 km depth (Williams pers. comm). A large E N E - W N W positive magnetic in the same region is thought not to relate to the same possible pluton but to a structure deeper in the basement. It is possible that the body reflected by the gravity anomaly acted as a rigid block during the Variscan deformation causing a decrease and perhaps the refraction of the maximum compressive stress to W N W - E S E , normal to the axis of the body and the main Caledonian strike resulting in the lower strain and anomalous trend of the syncline. Southwards, the axial trace of the main syncline is rotated clockwise through almost 70 ~ by syn- to post-fold dextral shearing on the Silvermines fault zone. In addition, the fold becomes tightened, and the axial trace truncated by the Silvermines fault, such that lower Dinantian limestones on the northern limb of the fold are in contact with Silurian rocks to the south. The eastern termination of the Silvermines fault zone against the Nenagh block in the Silvermines area, during the Dinantian, is possibly support for the rigid behaviour of this segment of the central Ireland zone. A comparable structural pattern is developed to the south, where the Slieve Felim anticline rotates clockwise into the N N E - S S W trending Gortdrum fault (Figs 4 & 8b). It is interesting to note that important ore bodies occur in this type of setting, which may also include Tynagh (Figs 4 & 8b). Ductile-brittle shear zones The pattern of minor folds and cleavage east of the Fergus shear zone is consistent with structures resulting from E N E - W S W hetero-

geneous dextral shear, parallel to the main Caledonian strike, in a transpressive regime, in which there was a N N W - S S E regional pure shear compression, compatible with the main Variscan fold development. A synopsis of the structures comprising the shear zone regime is shown in Fig. 7. Relatively high strain evidenced by intense pressure solution cleavage and moderate, strike parallel minor folds, is concentrated in steep, broad, E N E - W S W trending, dextral transcurrent shear zones (Figs 3 & 4). Many of these shear zones are associated with sharp changes of gradient in both the regional gravity and magnetic fields, as indicated by the lineaments in Fig. 3. There is a gradual decrease in strain northwards across the central Ireland zone, however, relatively strong cleavage and minor folding is present within shear zones as far north as Tynagh. The Quinn shear zone is the best example of a largely ductile shear zone (Figs 4 & 8a). It is approximately 3 km wide and trends E N E - W S W , subparallel to, and slightly offsetting the axial trace of a major syncline, the east Clare syncline. On the margins of the shear zone a weakly developed spaced cleavage stikes between N - S and N E - S W , markedly oblique to the shear zone margin. The main shear zone is largely defined by a well developed cleavage striking subparallel to the shear direction.

~,

v. ~=~

zo~

~

--/

~

..-z~_ ~

?

.......

......

\\

"-'~'- 1'41NOR Ff.LD

FIG. 7. Generalized diagram showing the geometry of secondary structures associated with the transpressive deformation regime in the central Ireland zone. Regional compressive stress (pure shear), a, is based on the trend of the major folds. The field of orientations of observed minor shears, dextral (d) and sinistral (s), are coincident with R and R Riedel shears respectively, for an ENE trending dextral shear couple. V represents a set of rotated extension veins.

192

D. W. Coller

FiG. 8. Ductile-brittle shear zones in west central Ireland. (a) Quin shear zone and associated structures. Approximate limits of high strain are shown by dash-dot lines. Stipple, Lower Palaeozoic; ORS, Old Red Sandstone; SR, R and SpR are sub-reef, reef and supra-reef limestones respectively, comprising the Lower Carboniferous. (b) Gortdrum fault-shear zone after Tyler (1979). Open stipple, Old Red Sandstone; close stipple, white and block shading are lithostratigraphic units of the Lower Carboniferous. (c) Major transcurrent faults and associated ductile deformation in Slieve Aughty, after Emo (1978). Faults, thick lines; bedding trace in Old Red Sandstone and Lower Carboniferous (undifferentiated, white), thin lines. Variscan cleavage, bars with ticks. Stippled areas, Lower Palaeozoic inliers with Caledonian cleavage, dashed lines. (d) Southern termination of the Fergus shear zone. Fields of minor dextral (d) and sinistral (s) shear sense of reactivated extension veins are shown with respect to the approximate finite strain ellipse for the shear zone. In the inset, P, principal fault; R and R, synthetic and antithetic Riedel shears; E, extension faults; C, cleavage and minor folds.

Variscan structures in west central Ireland Small-scale en echelon folds are developed largely parallel to the cleavage, within and on the margins of the shear zone, having a similar clockwise rotation to the cleavage into the zone. There is, however, commonly an obliquity between the cleavage and axial traces of minor folds, which may be explained as a feature of rotational strain related to the shearing, where folds may have initiated at a high angle to the shear direction (Sanderson et al. 1980). In the area of intersection of the margins of the Quinn and Fergus shear zones the minor folds are typically basin and dome style. These are considered to be interference folds due to superimposed, or possibly coeval, shear movements on the two zones. From the moderate intensity of cleavage in the limestones within the higher strain zones, and the relatively small stratigraphic displacement associated with the Quinn shear zone, the simple shear component of the finite strain cannot be very large. A n estimate of the minimum shear strain, based on the rotation of N W - S E trending regional joint sets, is ~ = 1. This order of shear strain is incompatible with cleavage and minor folds forming subparallel to the shear zone without a significant component of pure shear across the zone, as in a transpressive regime (Harland 1971). It is considered that this pure shear component is the regional, N N E - S S W compression which gave rise to the major folds, rather than a more local compression generated by the shear zones themselves, typically in areas of converging transcurrent faults, as described by Moore (1979). The above relationships suggest that the main fold and shear zone developments were part of a continuous deformation with an approximately N - S compression and an increasing component of simple shear resulting in one regional cleavage. Many of the other E N E trending shear zones are associated with major brittle faults, all of which show a dextral transcurrent sense of movement. In Slieve Aughty, gently folded bedding in the Old Red Sandstone and Lower Dinantian rocks swings clockwise, and steepens across broad zones adjacent to E N E - W S W faults, locally developing a shear-parallel cleavage (Fig. 8c). Some of these faults extend westwards from Slieve Aughty, into Upper Dinantian limestones, largely as broad zones of cleavage. Mapping in Slieve Aughty by E m o (1978) shows that a number of slightly oblique en echelon folds are rotated into, and truncated by

193

these major E N E faults indicating that brittle deformation did supercede much of the ductile deformation (Fig. 4). Planar vertical extension veins, believed to be related to the folding, are widely developed throughout the Dinantian limestones, and they commonly exhibit horizontal recrystallized mineral fibres due to transcurrent shear reactivation on the veins. It is possible in many cases to determine the sense of shear movement of the veins. Fields of sinistral and dextral shears for the central Ireland zone are orientated about NNW and E N E respectively, which is broadly compatible with the expected antithetic and synthetic shear directions for a regional E - W dextral shear couple (Fig. 7). This pattern of minor shears changes to NNE sinistral, and NW dextral in the region of the Fergus shear zone, which is compatible with a change to a NNE sinistral shear couple. The precise age of these shear movements is uncertain, although t h e y generally post-date the cleavage.

Discussion The history of deformation in the west of Central Ireland would tend to support the general hypothesis outlined by Reading (1980), that this marginal zone of the Variscan orogen evolved as a major dextral strike-slip belt. Early Variscan deformation is characterized by a period of basement controlled faulting, dominated by major E N E dextral transcurrent (strike-slip), and compatible, subordinate NW extensional (normal) faults. It is suggested that this early fault regime was one of transtension (transcurrent shear in an extensional regime) and controlled the complex basin development during the Upper Devonian and Lower Carboniferous. Following continued basin development in the Upper Carboniferous, the main Variscan event involved ductile deformation in a transpressive regime, with E N E dextral shearing along linear zones, largely reactivated older faults, during a N N E - S S W regional compression. ACKNOWLEDGMENTS:I would like to thank numerous colleagues for their helpful criticisms and comments during the preparation of this paper, in particular W. E. A. Phillips, D. J. Sanderson, G. D. Sevastopulo, M. F. Critchley, J. Dolan, N. Haughey and R. Later. I wish to acknowledge the help given by the Geological Survey of Ireland and the Dublin Institute of Advanced Studies in providing the geophysical data.

194

D. W. Coller

References ANDREW, C. J. & ASHTON, J. H. 1982. Mineral textures; metal zoning and ore environment of the Navan orebody, Co. Meath, Ireland. In: BROWN, A. G. (ed.) Mineral Exploration in Ireland Progress and Developments 1971-1981,35-46. Irish Association for Economic Geology. BROCK, P. M. 1982. The regional lithostratigraphical setting of the Silvermines zinc--lead and the Ballynoe barite deposits, Co. Tipperary. In: BROWN, A. G. (ed.) Mineral Exploration in Ireland Progress and Developments 1971-1981, 162-70. Irish Association for Economic Geology. DUBLIN INSTITUTEADVANCED STUDIES, 1974. Gravity anomoly map of ireland. Bull. Dubl. Inst. Advd. Stud. Geophys. 32. EMO, G. T. 1978. The Lower Palaeozoic and Old Red Sandstone of Slieve Aughty, Counties Clare and Galway. Unpublished Ph.D. Thesis. University of Dublin. -& GRENNAN, E. F. 1982. A review of mineralization in the Old Red Sandstone and its significance to Irish exploration. In: BROWN, A. G. (ed.) Mineral Exploration in Ireland Progress and Developments 1971-1981, 27-34. Irish Association for Economic Geology. FEEHAN, J. 1980. Alluvial fan sediments from the Old Red Sandstone of Devilsbit Mountain, County Tipperary, J. Earth Sci. R. Dubl. Soc. 3, 179-94. HARLAND, W. B. 1971. Tectonic transpression in the Caledonian Spitzbergen. Geol. Mag. 108, 27-42. HOLLAND, C. H. 1981. Devonian. In: HOLLAND, C. H. (ed.) A Geology oflreland, 121-46, Scottish Academic Press, Edinburgh. MOORE, J. McM. 1975. Fault tectonics at Tynagh Mine, Ireland. Trans. Instn rain. metall. 84, B141-5. - 1979. Tectonics of the Najd Transcurrent Fault System, Saudi Arabia. J. geol. Soc. London, 136, 441-54. MORRIS, J. H. 1979. The geology of the western end of the Lower Palaeozoic Longford-Down Inlier,

Ireland. Unpublished Ph. D. Thesis. University of Dublin. PHILLIPS, W. E. A., FLEGG, A. M. & ANDERSON, T. B. 1979. Strain adjacent to the Iapetus Suture in Ireland. In: HARRIS, A. L., HOLLAND, C. H. & LEAKE, B. E. (eds) The Caledonides o f the British Isles--reviewed. Spec. Publ. geol. Soc. London, 8 , 257-62. Scottish Academic Press, Edinburgh. --., STILLMAN, C. J. & MURPHY, T. 1976. A Caledonian plate tectonic model. J. geol. Soc. London, 132, 579-609. READING, H. G. 1980. Characteristics and recognition of strike-slip fault systems. In: BALLANCE, P. F. & READING, H. G. (eds) Sedimentation in Oblique-slip Mobile Zones. Spec. Publs int. Ass. Sediment. 4, 7-26. Blackwell Scientific Publications, Oxford. RHODEN, H. N. 1958. Structure and economic mineralization of the Silvermines District, Co. Tipperary, Eire. Trans. lnstn min. Metall. 68, 67-93. SANDERSON, D. J., ANDREWS,J. R., PHIl.LIPS, W. E. A. & HUTrON, D. n . W. 1980. Deformation studies in the Irish Caledonides. J. geol. Soc. London, 137, 289-302. SEVASTOPULO, G. D. 1981. Lower Carboniferous & Upper Carboniferous. In: HOLLAND, C. H. (ed.) A Geology o f Ireland, 147-87. Scottish Academic Press, Edinburgh. TAYLOR, S. & ANDREW, C. J. 1978. Silvermines orebodies, Co. Tipperary, ireland. Trans. lnstn Min. Metall. 87, B 111-24. TCHALENKO, J. S. & AMBRASEYS, N. N. 1970. Structural analysis of the Dasht-e-Boyaz (iran) earthquake fractures. Bull. geol. Soc. Am. 81, 41-60. TYI.ER, P. 1979. The Gortdrum Deposit. In: BROWN, A. G. (ed.) Prospecting in Areas of Glaciated Terrain. Excursion Handbook. Irish Association for Economic Geology. WILLIAMS, C. E. 1981. Aeromagnetic Maps o f Ireland. Parts of 1:250,000 sheets 1-5. Geological Survey of Ireland.

D. W. COLLIER, Department of Geology, Trinity College, Dublin 2, Ireland.

The Alleghenian orogeny in eastern North America N. Rast SUMMARY: The Alleghenian-Variscan-Hercynian orogeny in eastern North America has five structural domains: (1) Newfoundland-Maritime (northern) platform; (2) northern Appalachian Maritime-New England Variscan belt; (3) Alleghenian platform; (4) central-southern Appalachian (Alleghenian) orogen; (5) Ouachita belt. Connections between these domains are generally faulted or unexposed and the domains are heterogeneous. There is a major discontinuity between the northern Appalachians and central-southern Appalachians. In the northern Appalachians the Variscan orogeny is distinct from the Acadian, but not so in the central-southern Appalachians. The Alleghenian overthrusting of the southern Appalachians is still to be traced to the north where the whole of southern Nova Scotia may be totally allochthonous. Also, the northern but not the southern belt appears to have suffered in Devonian to Carboniferous times a sinistral strike-slip movement of about 2000 km. By taking this movement into account it is suggested that the four Appalachian domains have been assembled in Devonian-Lower Carboniferous time by transcurrent faulting and deformed in the late Carboniferous-Permian times by subduction-related processes. The geology of the Ouachita belt is different from the southern Appalachians and their unexposed connection is still enigmatic. The timing of orogenic events in the two domains also differs. It is possible that the Ouachitas were deformed by collision with the South American plate, while the Appalachians were deformed by collision with the African plate.

The initial continuity of the Variscan (Hercynian) orogen of Europe and the Alleghenian orogen of North America was first proposed by Bailey (1929). I have recently examined the stratotectonic principles and problems involved in such a transatlantic correlation (Rast 1983), but research in North America has been so fast and so far-reaching that it is apt to inquire into the present status of the structure of Carboniferous rocks of eastern North America, as well as the orogenic and plate tectonic mechanisms that went into their deformation. At the time when Bailey proposed his correlations it was thought that the northern Appalachians were affected principally by the mid-Ordovician (Taconic or Taconian) and the Middle Devonian (Acadian) orogenies, while the central and southern Appalachians suffered the Alleghenian deformation (Fig. 1). Thus, with respect to the British Isles Bailey equated the T a c o n i a n - A c a d i a n orogen with the British Caledonides and the Alleghenian orogen of the central-southern Appalachians with the British Variscides--the two crossing each other. Even in North America until recently two very different conceptions of the Appalachian orogenic sequence existed. In the northern Appalachians the emphasis lay on the Acadian orogeny, while in the central-southern Appalachians the socalled Appalachian (late Palaeozoic) revolution was stressed (e.g. King 1959), although by 1977 King spoke of a separate phase of Allegheny

orogeny. The changeover from the northern to central-southern Appalachians occurs just north of 41 ~ parallel, where it is obscured partly by the Meso-Cenozoic coastal plain deposits and partly by the extensive development of Triassic half-grabens (Fig. 2). Thus, the bulk of the northern Appalachian deformed rocks trending almost n o r t h - s o u t h into the Sound of Long Island are separated from the presumably onstrike equivalents of Virginia and North Carolina by several hundred kilometres of unknown relationships. Correlations are difficult, partly because of the different levels of erosion and the fact that the highly metamorphosed strata in the central and southern Appalachians have yielded very few fossils. Even the continuity of central and southern Appalachians bristles with correlation problems, although in this paper for the sake of simplicity the central and southern Appalachians are grouped together. Throughout this region, however, dated Carboniferous rocks are recognized solely in the north-western zones adjacent to the craton. The absence of Carboniferous rocks from the interior of the belt renders the dating of structures there very difficult. Lastly, separated by an extensive tongue of the Meso-Cenozoic coastal plain deposits, in Mississippi and Louisiana (referred to here as the Mississippian embayment), there is an inlier of an orogenic belt known as the Ouachitas (Fig. 1). This comes to the surface again some 197

198

N. R a s t

FIG. 1 Position and divisions of the Variscides in eastern North America. Fronts of northern, central and southern and Ouachitan belts are indicated by double lines. Letters indicate Canadian provinces and American states. Provinces: N--Newfoundland, NB--New Brunswick, NS--Nova Scotia, PEI--Prince Edward Island, Q--Quebec, NB, NS and PEI = Canadian Maritimes. States: northern Appalachians: C--Connecticut, M--Maine, MA--Massachusetts, NY--New York, RI--Rhode Island; central Appalachians: P--Pennsylvania, V--Virginia, WV--West Virginia; southern Appalachians: A--Alabama, G--Georgia, K--Kentucky, NC--North Carolina, SC--South Carolina, TN--Tennessee. Ouachitas: AR--Arkansas, L--Louisiana, MS--Mississippi, O--Oklahoma, T--Texas, Mar--Marathon inlier. I--Illinois basin, II--Kansas-Nebraska basin, Ill--Michigan basin, IV--Ozark dome. Dashed line is the edge of the coastal plain--its northward deflection south of Illinois basin in the Mississippian embayment.

800 km to the SW as the Marathon inlier. The connection between the O u a c h i t a - M a r a t h o n belt and the central-southern Appalachians is hidden under the Mississippian embayment. Thus, although Thomas (1973), on the basis of deep borehole data, thinks that there is a connection, there are insufficient data to demonstrate an incontrovertible tectonic continuity. There are very substantial depositional and structural differences between the southern Appalachians and the Ouachitas which, therefore, must be considered as a separate segment. The three very different on-trend segments of the Alleghenian-Variscan belt of North America are not homogeneous across strike. In the northern Appalachian segment Carboniferous rocks can be divided into undeformed strata of New Brunswick, Prince Edward Island and

parts of Nova Scotia to the NW, while variably deformed and, in part, regionally metamorphosed strata lie to the SE. Deformed and metamorphosed Upper Carboniferous strata have long been known from R h o d e Island (Quinn 1971; Skehan et al. 1979). The deformation and metamorphism in Maritime Canada and R h o d e Island were attributed until recently to localized phenomena (cf. Maritime disturbances, Poole 1967), to disturbances along major faults (Webb 1969; Eisbacher 1970) or to the effects of igneous intrusions (Quinn & Moore 1968). The first change in approach was introduced by Rast & Grant (1973a, 1977) and Rast et al. (1976) who suggested that the eastern northern Appalachians can be considered as a continuation of the Variscan orogenic belt of Europe. Thus, two tectonic domains can now

Alleghenian orogeny in America

199

FiG. 2. The relationship of the central and northern Appalachians. AF--Allegheny front, AS--Allegheny synclinorium, BF--Bloody Bluff Fault, BR--Blue Ridge, CN--Clinton Newbury Fault, LL--Logan's Line, NAF--northern Appalachian front (limit of significant deformation), t--Triassic grabens. Indicated states: Md--Maryland, P--Pennsylvania, V--Virginia. be recognized in the northern Appalachians-the western including Newfoundland, most of New Brunswick, Quebec and Maine, where rocks are entirely undeformed or only show deformation associated with major faulti n g - a n d the eastern where orogenic deformation is discerned (Fig. 3). In the central-southern Appalachians a very

similar situation exists. From western Pennsylvania and New York down to Alabama and Mississippi lie the undeformed or weakly deformed strata of the Appalachian and Blac k Warrior basins (Fig. 4), while separated from these to the SE are the extensively deformed Carboniferous rocks of Virginia, Tennessee and eastern Alabama. The two tectonic domains

FIG. 3. Northern Appalachians--platform and deformed belt. Mississippian--dashed, Pennsylvanian--stippled, old massifs in the Maritimes--crossed. Faults: B--Belleisle, BF--Bloody Bluff, C--Cabot, CC--Cobequid-Chedabuct, CN--Clinton-Newbury, H--Honeyhill, HB--Hare Bay, N-F--Nurembega-Fredericton, VF--Variscan front. Basins: N--Narragansett, CS--central syncline, granites in black, thrusts with black triangles.

N. Rast

200

FIG. 4. Southern-central Appalachians. kmb--King Mountain belt, pm--Pine Mountain thrust,

granites--black, gabbros--crossed. C-C'--COCORP seismic lines, G-G'--Geological Survey seismic lines. Major overthrusts shown by black triangles.

can be called the Alleghenian platform and the Alleghenian orogen respectively. Lastly, the Ouachita segment, although it is amenable to division internally because of existing disagreements, would at present be treated as a single tectonic domain.

Northern Appalachian platform The northern Appalachian platform occupies the area of unmetamorphosed and little deformed Carboniferous strata found between Newfoundland and Maine. Its outlines further south are indeterminate since there are no known outcrops of Carboniferous strata, but it undoubtedly continued and joined the Alleghenian platform of New York and Pennsylvania. To the east a provisional line of separation can be drawn on the basis of deformation of Carboniferous rocks (Fig. 5).

Lithology and succession Throughout the area (Fig. 3), the Carboniferous succession, which on the basis of the flora contained and lithologic comparisons with known areas, is chronostratigraphically divided into the Mississippian and the Pennsylvanian. The lower part of the Mississippian in places conformably and elsewhere unconformably

overlies Devonian strata (Rust 1981). The Carboniferous succession is variable from place to place, but in general it consists of variegated Mississippian deposits with, in places, a predominance of red beds and Pennsylvanian red and grey sandstones and conglomerates. In south-western New Brunswick Mississippian bimodal volcanic rocks (basalts and ignimbrites) are well developed and have been described by Ruitenberg (1963) and Van de Poll (1963, 1967, 1972). The precise age of these volcanics is badly known, but speculatively (at present) both Tournaisian and Namurian ashflows seem to be present. To the SE in New Brunswick and Nova Scotia, and also to the NE under the Gulf of St Lawrence considerable deposits of Upper Visean halite are encountered, but in general these deposits are much better developed among the more deformed Carboniferous strata of Nova Scotia, where they show rapid variations in thickness, reaching in places as much as 900 m (Van de Poll 1972). The Pennsylvanian strata, especially well represented in central New Brunswick in the broad central New Brunswick syncline and that are at present being actively investigated by the Department of Natural Resources of New Brunswick, are mainly grey and red beds of a considerable thickness. Howie & Barss (1975) estimate almost 2 km of thickness in the deeper

Alleghenian orogeny in America

201

FIG. 5. Tectonic map of Maritime Carboniferous deformed belt. Direction of derivation of sediments (adopted from Van de Poll 1973; Rust 1981) is indicated by arrows. Data after Van de Poll (1973), Legun & Rust (1982), Fralick & Schenk (1981) and Rast (1983). part of the central New Brunswick syncline, but in the central part of the Gulf of St Lawrence the thickest sequence is mainly of Mississippian sediments reaching over 8 km. In parts of the Pennsylvanian sediments coals of commercial quality are known. The Pennsylvanian strata pass gradually into Permian red beds exposed in Prince Edward Island. Structure

In exposures, in general, neither Mississipplan nor Pennsylvanian rocks are appreciably deformed and dips of the Pennsylvanian sediments, except on the flanks of the central New Brunswick syncline, do not exceed 10 ~ Major exceptions are observed where Mississippian strata are caught up within major faults as lenticles in which case their dips are steep and in places they develop crude cleavage (Rast 1984). In Newfoundland a considerable degree of structural disturbance is developed in the NW of the Cabot fault (Fig. 3) (Knight 1975). In Maine a small lenticle interpreted as a graben is found in association with the Frederic-

ton-Norumbega fault (Larrabee, Spencer & Swift 1965; Doyle 1979). Although the dip here is not very steep this is similar to other Mississippian fault-bound lenticles. These observations suggest that Mississippian strata have been deformed, presumably in connection with faulting, prior to the deposition of the Pennsylvanian and to post-Pennsylvanian gentle folding, which is assumed to be the platformal manifestation of the Alleghenian orogeny. Cataclastic effects associated with the Alleghanian orogeny and later events have been analysed by Ruitenburg et al. (1977). To the NW of the central New Brunswick syncline the Alleghanian deformation rapidly disappears. In northern New Brunswick the Mississippian (?) Bonaventure formation is entirely horizontal and its equivalent in Quebec (Cannes de Roche formation) is only very gently folded. Rust (1981) has suggested that the Mississippian rocks of northern New Brunswick and Quebec were deposited penecontemporaneously with Carboniferous transcurrent faulting. Prior to that, association with graben for-

N. Rast

202

mation was claimed (Belt 1968). Kelley (1967) had in fact proposed that the thick Carboniferous strata of south-eastern New Brunswick were formed in epieugeosynclinal grabens, an idea that was further pursued by Howie & Barrs (1975). While, in principle, either association with strike-slip or normal faults is possible, at present the choice between the two is difficult. Perhaps the presence of lenticles of Mississippian and earlier sediments in the faults is more consistent with a strike-slip origin, indeed the Fredericton-Norumbega fault has an undoubted right-lateral strike-slip displacement along it. The derivation of the sediments in the northern Appalachian platform presents an interesting situation. In the western north-western part of the platform the sediments have an entirely local provenance (Rust 1981) and are variable. The cross-bedding in Mississippian strata reflects both the fans emerging from uplifted blocks and flow parallel to major valleys. In the central New Brunswick syncline however, Van de Poll (i 973) has recognized in most of the Pennsylvanian rocks a generalized direction of sediment derivation principally from the SW and west (Fig. 5), indicating that the syncline was also a basin and the sediment transport was easterly. Westerly derivation has been confirmed recently by Legun & Rust (1982) from northern New Brunswick.

Allegheny platform The Allegheny platform is a broad tract stretching from New York (Fig. 4) to Alabama and Mississippi. Internally it can be divided into major basins, domes and graben-like troughs. The Appalachian basin (Dennison 1977) and the Black Warrior basin (cf. Thomas 1972) form the two most important depositional troughs, the western part of which is broken up by smaller but still considerable structures (Donaldson & Shumaker 1981).

Lithology and succession The general sedimentation is similar to the northern Appalachian platform, in so far as the Mississippian is partly marine and the marine facies is represented by carbonates of Visean-Namurian age. But, while the carbonate depositing sea in the Northern Appalachian platform lay to the east and SE of the landmass in relation to the central-southern Appalachians the sea was to the west and NW of the landmass.

Pennsylvanian rocks of the Allegheny platform are essentially fluvial or deltaic deposits (Ferm 1970, 1971) and their relation to the Mississippian is still debatable, although in places there is fair evidence for an unconformity separating the two sequences. Thus Mississipian clastics and carbonates are succeeded by Pennsylvanian clastics, in part red beds but in part coal measures. The alleged Permian (Dunkard) rocks overlying the Pennsylvanian may in fact be Upper Stephanian. Much recent research on the stratigraphy and sedimentation of this region is described in the United States Geological Survey Professional Paper 111 O-A to L (1979). The Mississippian and Pennsylvanian of the Allegheny platform show considerable variations in thickness, since in the Appalachian basin sources of derivation were both from the rising Appalachian chain and from the North American craton (Donaldson & Shumaker 1981). In general, however, the combined thickness does not much exceed 1 km except within the Appalachian mountains. In addition the thickness is partly controlled by the salients and recesses of the Appalachian orogenic belt (King 1977, p. 62; Thomas 1977). Thomas, in particular, advances an attractive hypothesis that during the continental collision that resulted in the Alleghenian orogeny salients developed particularly thick sedimentary clastic wedges when the collision first occurred. Carboniferous sediments of the Appalachian basin together with those in the Appalachian chain are commonly referred to as molasse. In the sense that molasse is a post-orogenic deposit (Van Houten 1981) this is not valid, since deposits of this kind existed in the Appalachian region from Upper Devonian to the P e r m i a n - - a duration of some 100 x 106 yr. In an overall structural context, this problem will be referred to again.

Structure The Allegheny platform constitutes the easternmost south-eastern edge of that part of the United States that King (1977) termed the Interior Lowlands. The central part of that vast area is characterized by large basins such as: the Michigan basin, the Illinois basin, the Kansas-Nebraska basin, etc. (Fig. 1) separated by intervening arches and at least in one case (Ozark dome) a major uplift. In this region the rocks, although domed and depressed, are not folded. Towards the SE the arches become sharper (Cincinnatti) and the domes smaller (Lexington, Nashville, Pocono) and to the NE

Alleghenian orogeny in America in Pennsylvania and New York fairly strongly folded strata form a major Allegheny synclinorium, that follows the trend of the Appalachian basin (Fig. 2). The Appalachian basin underlies the geomorphic province known as the Appalachian plateau which is immediately adjacent to the Appalachian orogen. In this area signs of Alleghenian deformation are detected in tight folds, which involve complex and in places chaotic minor structures (Wheeler 1978), that are nevertheless thought to be tectonic rather than sedimentary, although the latter also abound in the area and are associated with slumps. In general only one generation of tectonic folds is seen, but in Pennsylvania and New York fold interference due to refolding can be recognized (P. Geiser, personal communication), the first with east-west orientation and the second being more N E - S W (Engelder & Geiser 1981). Such structures affect both Mississippian and Pennsylvanian rocks and clearly represent late phases of the Alleghenian orogeny. This can also be deduced from sedimentological studies. Moore & Clark (1970) have described Tournaisian (Lower Mississippian) flysch from Kentucky. Perry & de Witt (1977) outlined Mississippian debris flows from West Virginia. Lastly, Bruno (1982), following a pioneer work of Krynine (1948) on the analysis of Mississippian and Pennsylvanian rock fragments from West Virginia and Virginia has indicated a progressive denudation of presumably rising uplands within the Appalachian belt, since in all these cases the current bedding implies derivation from a south-westerly provenance. A derivation from the Canadian shield (North American craton) is proposed for the sediments of the NW side of the Appalachian basin (Edmunds et al. 1979). There are, however, other orientations of deformation, which at present are represented by surface faulting and for which very old sequences of deformational events, going back to the Cambrian, are suggested (Donaldson & Shumaker 1981) on the basis of subsurface data. Of these the so-called Rome trough is in part subparallel to the Appalachian orogen, but has diverging cross-cutting lines of faulting--the 38th parallel trend, the West Virginia trend and the 40th parallel trend (Drake & Woodward 1963). It is also claimed that displacements on these lines changed in time and normal throw to the south prevailed in middle or late Palaeozoic times. Shumaker (1975) made a good case for relating faulting in the mid-continent to orogenic events. The continuation of the 40th parallel trend coincides with the changeover of the Appalachian regional

203

strike from the northern to central Appalachians and the 38th parallel trend marks the change from the central to southern Appalachians.

Northern Appalachian Variscan belt Until comparatively recently it was thought that the Alleghenian deformation had not much affected the Carboniferous successor basins of the northern Appalachians. It is true that deformed and metamorphosed Carboniferous rocks were described from Rhode Island (for bibliography see Quinn 1971) and Gussow as far back as 1953 recognized overthrusting in New Brunswick, which he attributed to late Mississippian deformation. It was Rast & Grant (1973a, 1977) who first proposed that the orogenic belt did exist in the northern Appalachians and that it can be directly related to the Variscan of Europe. In particular these authors suggested that the edge of overthrusting in southern New Brunswick could be extended into the Cobequid-Chedabucto fault zone that transects Nova Scotia. This fault zone was assumed to be a connecting fault with the Variscan front of Ireland, Wales, England and northern France (Rast 1983). The fact that the overthrust front of southern New Brunswick terminated in such a transcurrent (Fig. 3) fault was not inherently surprising since as known for a long time (cf. Rodgers 1970) a major overthrust of this kind (Pine Mountain of Tennessee and Kentucky) similarly terminates in a transcurrent fault (Jacksboro). In this sense the front represents laterally continuous but differing strain responses to orogenic stress.

Succession The Carboniferous succession (Tournaisian to Permian) in the deformed area of southern New Brunswick and Nova Scotia is again divided into Mississippian and Pennsylvanian. The total thickness of these sequences is variable and these variations are attributed to the existence of penecontemporaneous grabens and horsts (Kelley 1967, 1970). Kelley in particular maintained that the thickest deposits were formed within the so-called Fundy epieugeosyncline where in places the succession reached over 5 km, but that the Pennsylvanian deposits were unconformable on the Mississippian. The succession in Nova Scotia consists of Devonian and Lower Mississippian (Horton) clastics overlain by Visean (Windsor) carbonates and finally by Upper Mississippian (Namurian) and Penn-

204

N. Rast

sylvanian red beds and fluvial deposits. The carbonates are thickest in Nova Scotia and thin into southern New Brunswick. In New Brunswick, both mafic and felsic volcanics are involved with the Mississippian sediments (Rast etal. 1978; Strong, Dickson & Pickerill 1979) although in Nova Scotia there are few Carboniferous volcanics (Kelley 1970, p. 289). Several internal unconformities are recognized (Howie & Barss 1975). Yet the most important unconformities are sub-Visean (Kelley 1970; Keppie 1980) and sub-Pictou (Westphalian C) of which the latter marks a completely new sedimentation that was in part lacustrine. The Pennsylvanian alone in places reaches 3.5 km. Nevertheless thicknesses vary from one lithostratigraphic unit to the other and no completely unified depocentre can be detected in the area. Throughout the region Carboniferous conglomerates and sandstones predominate and occur on several levels. Coals are known, especially from the Pictou of northern Nova Scotia. The deposition of conglomerates is often attributed to penecontemporaneous faulting (Fralick & Schenk 1981), although this had probably terminated by Westphalian B times. Therefore Westphalian C to Permian sediments, grouped together as the Pictou Group (Bell 1938), overran the whole region and continued from the platform into the epieugeosyncline. Because of the complexity of original topography a highly variable provenance of sediments existed from place to place and from time to time (cf. Fralick & Schenk 1982). Because of its marine nature and abundance of fossils the Visean has been particularly comprehensively studied (Bell 1929, 1940; Schenk 1967, 1970, 1976). The latest interpretation is that it represents a cyclically deposited sequence, alternating between marine (carbonates), shallow dessication playa (evaporites) and terrestrial (conglomerates) environments. The evaporites are particularly significant since the interpretation of the tectonics of the region often arose from the suggestion that the movement of salt produced local structures (Howie & Barss 1975). To the NE of the Canadian mainland Carboniferous sediments are known from Newfoundland where again the clastic equivalents of Upper Devonian-Lower Mississippian are known as the Anguille Group that is overlain by lateral equivalents of Windsor Group rocks (Codroy Group). Pennsylvanian (?) deposits in Newfoundland are mainly conglomeratic and are associated with the Cabot fault. Similar, submarine, outcrops of Pennsylvanian (?) are also known from an area north of Newfound-

fl.. ./~

o

Kms,o

FIG. 6. The geology of south-eastern New England (after Skehan & Murray 1979b). NPG--Narragansett Pier granite, 1--Narragansett basin, 2--Norfolk basin, 3--Woonsocket basin, 4--Scituate basin, 5--Worcester outcrop, 6--Pin Hill outcrop, B--Boston, C--Connecticut, NH--New Hampshire, M--Maine, MA--Massachusetts, RI--Rhode Island, CN--Clinton-Newbury fault. land (St Anthony Basin), although no deposits of the Pictou Group are known at present. Such coarse deposits are interpreted as fault-related conglomerates in this area. To the SW in New England the Carboniferous deposits of the orogenic belt occur in a series of basins, of which the largest, Narragansett, occupies a large area of the states of Rhode Island and Massachusetts. It has smaller subsidiary basins (Norfolk, Woonsocket and Scituate and also further north where there are small outcrops near Worcester and Pin Hill). The geology of these areas has recently been reviewed by a series of authors in a comprehensive guidebook edited by Cameron (1979) and in the U.S. Geological Survey Professional Paper 1110-A (cf. Fig. 6). Until recently sediments of the Boston basin were considered as Carboniferous, but recent data by Kaye & Zartman (1980) have led to the re-evaluation of these rocks as Precambrian. The deposits of the Narragansett basin (Mutch 1968; Skehan & Murray 1979b) form an exceedingly thick sequence (c. 4 km) of conglomerates, sandstones, shales and felsic volcanics often strongly deformed and in part

Alleghenian orogeny in A m e r i c a metamorphosed. Similar but even coarser sediments occur in the Woonsocket and north Scituate basins. Hepburn (1979) also correlates the metaconglomerates and anthracites of the Worcester outcrop with the Narragansett basin. It is, in the absence of fossils, somewhat doubtful whether the conglomerate at Pin Hill is Carboniferous, but it seems likely to be so (Thompson & Robinson 1976). The deposits of the Narragansett basin have been accurately dated as Westphalian B to Stephanian (Lyons 1979) and thus correspond to a part of the Pictou of the Canadian Maritimes. However, while Pictou deposits were derived from south-westerly sources the Narragansett sediments were derived from the NE sources. Thus, in Uppe~ Carboniferous times there was a land barrier intervening between the Maritimes and southern New England. Furthermore, some proximate derivation is also suspected since the coarser sediments of Narragansett are often arkosic (Quinn 1971) and there are often coarse conglomerates near the edges of the basin. Whether the basin was originally fault conditioned or not is at present debateable (Skehan & Murray 1979a, p. 10).

Structure and plutonism An attempt at the general correlation of the internal structures of the Variscan-Alleghenian deformed rocks in the northern Appalachians has been made already (Rast & Grant 1977; Rast, Skehan & Grant 1976; Rast 1983; Mosher & Rast, this volume). Internally the strata of New Brunswick and Narragansett basin, as well as parts of Nova Scotia show polyphase deformation, low angle thrust faults, injection of granitoids and regional metamorphism (Rast et al. 1978; Skehan & Murray 1979b). However, on a major scale in New Brunswick and northern Nova Scotia (Currie 1977) large-scale low lying north-westerly facing thrusts determine the position of the Variscan front, while in New England the front allegedly runs along the Clinton-NewburyBloody Bluff System of faults (Skehan & Murray 1980a) (Fig. 3) with movement on these reversed faults being to the SE. I do not think that a recent claim by Goldstein (1982) that the Bloody Bluff fault is in effect a normal fault is valid. However, across the ClintonNewbury fault, i.e. to the west of it (Fig. 6) lies the Worcester outcrop. The rocks here have two cleavages, bear manganiferous garnets and also Westphalian plants (Lyons 1976). Adjacent to these rocks there occurs an early Carboniferous granite that has yielded an

205

isotopic age ofc. 345 Ma (Zartman et al. 1965). Thus the Clinton-Newbury fault is not a continuation of the Variscan-Alleghenian front and the front should be located somewhat to the west of it. The age of the Variscan-Alleghenian orogeny in New England-Canada can be estimated as postWestphalian C in New Brunswick and postStephanian in New England. Isotopic age determinations that are available for New England yield a Lower Permian date of c. 270-260 Ma (cf. Skehan & Murray 1980a). More recent work by Dallmeyer (1982) suggests a prolonged cooling extending to 230 Ma. No accurate Alleghenian dates are at present available from Canada. The Carboniferous dates of some granites in eastern Newfoundland (Bell & Blenkinsop 1977) have been strongly disputed (Dallmeyer, Blackwood & Odom 1981), but the fact that sediments with Westphalian C flora are deformed in southern New Brunswick suggests a late Carboniferous to Permian age. During the deformation and metamorphism which, according to Dallmeyer (1982) lasted no more than 40 • 106 yr, two important fold phases associated with thrusting can recognized (F 1 and F:) (Rast & Grant 1973a,b). The first involved, both in New Brunswick and in New England, transport to the NW. Yet, while in New Brunswick F 1 was the major episode and F z folds are smaller in amplitude and the thrusts minor, in New England the F 2 folds are major and the related thrusts are large scale and associated with well-developed mylonites. The transport on these later structures is to the SE and they represent back-thrusting. A similar sequential folding and resulting confrontation of thrusts and therefore transport directions was first noted by Roberts & Sanderson (1971) from Cornwall (England), and an elaborate description of this was recently produced by Hobson & Sanderson (1983). In both cases--Cornwall and northern Appalachians-polyphase relationships are clear-cut. However, Hancock, Dunne & Tringham (1983) describe from South Wales box-like folds with opposing facing directions that seem to be conjugate. It seems possible that even the sequential folds are kinematically related. The metamorphism in Rhode Island is of regional, Barrovian type and was originally thought to be an aureole of the Narrangansett Pier granite in south-western Rhode Island and Connecticut. How far the regional metamorphism extends into Connecticut is at present difficult to determine, since it is difficult to separate it from the Acadian high grade effects.

206

N. Rast

However, it certainly continues west to the Honeyhill fault (Dallmeyer 1982). The Narragansett Pier granite is a composite intrusion that has yielded a Permian age (Kocis & Murray 1977; Kocis et al. 1978). Whether the metamorphic effects in New Brunswick are Barrovian or not is difficult to ascertain since the metamorphic grade is not high enough. Small granitic bodies have been described from New Brunswick (Ruitenberg et al. 1977; Rast et al. 1978). In relation to structure, metamorphism is generally syntectonic to post-tectonic and various phases of it can be recognized (Skehan & Murray 1979b). The granites are partly or completely post-tectonic. Thus, the Variscan-Alleghenian disturbance of the northern Appalachians is noticeable at the eastermost edge of the mainland of Canada and the U.S.A. The event involved appreciable overthrusting with the principal direction of transport being to the NW. The presence of granitic bodies and areas of regional metamorphism of medium pressure type imply a collisional event. What actually collided with North America is at present a mystery. There are claims that North African (Moroccan, Algerian) late Carboniferous strata are similar to North America (Van Houten 1976). Even the orientation of the stress system in North Africa (Michard & Pique 1980) appears to have been similar to that of the northern Appalachians, while Simpson, Bothner & Shride (1980) speculate that the Clinton-Newbury fault system continued into the South Atlas fault of north-west Africa. Therefore, the only possible candidate for Variscan collision with North America is Africa, and it is this collision that generated the overthrust structures of the northern Appalachians. There remains, however, the interpretation of events that were post-Middle Devonian (Acadian) and pre-Pennsylvanian. Accurate observations have shown that in general the Mississippian deposits in the northern Appalachians were more deformed than the Pennsylvanian. Van de Poll (1972) indicated that in New Brunswick there are numerous conglomerates throughout the Mississippian and underlying Devonian succession, starting with the Memramcook Formation, then the Hopewell Formation and ending with the Boss Point Formation, in each case indicating some tectonic disturbances in the orogenic belt. Similar indications of movement reflected in conglomerates is reported from Newfoundland (Knight 1975; Fong 1975). Gussow (1953) pointed out that the Mississippian rocks of south-eastern New Brunswick were far more

deformed than the Pennsylvanian. Even in the platformal part of New Brunswick Mississippian strata in fault lenticles show signs of fairly severe deformation. The fact that in Nova Scotia, Windsor or other Visean sediments are often angularly unconformable on earlier Carboniferous and Upper-Middle Devonian strata suggests considerable and continuous episodes of movement from the Middle Devonian to Pennsylvanian (in fact Westphalian B). The traditional interpretation has been that the Mississippian movements, that can collectively be correlated with the Sudetic orogenic stage of Europe, are due to the formation of horsts and grabens. A completely new concept has been advanced in the last few years on the basis of palaeomagnetic determinations. The precursor of these studies was produced by Morris (1976). He pointed out that the palaeomagnetic poles of the Ordovician to Lower Devonian of the British Isles differed significantly from those of North America, although he thought that Upper Devonian poles were similar. Therefore, he proposed a very large (1800 km) sinistral displacement of the British Isles with respect to the North American plate in the Middle Devonian times, thus causing the Acadian orogeny. At the time, information on Upper Devonian-Mississippian poles was insufficient. Later on Kent & Opdyke (1978, 1979) and Van der Voo, French & French (1979) revived the notion by suggesting some 15 ~ latitude displacement not only of the British Isles (cf. Van der Voo & Scotese 1981), but of a large continental fragment adhering to eastern North America, which is known as the Avalonian belt or Acadia displaced terrane (Kent 1983). The actual size of the displaced fragment is not known and various speculations are advanced. Van der Voo & Scotese (1981) suggest the Great Glen Fault as a candidate for this, although the notion has not been entirely accepted (Donovan & Meyerhoff 1982; Parnell 1982). Kent (1983) suggests that the Variscan front is connected to the Hare Bay fault of Newfoundland as a line of movement. Certainly there are few obvious faults to the NW of it (Ludman 1981) that can be used as a megashear (also cf. Brown & Kelley 1980). Rast (1983) suggests that the movements could have occurred not on a single fracture, but on a dispersed series of faults, each with a considerable, but not commensurate with continental scale, displacement. Under such conditions, the large closely spaced transcurrent faults can coexist with intervening grabens or so-called pull apart holes (Scrutton 1979). One difficulty of this system is that the majority of stratigraphically rising lines appears to have dextral direc-

Alleghenian orogeny in A m e r i c a tions of movement (cf. Fralick & Schenk 1981). Lefort & Van der Voo (1981) proposed an ingeneous model of east-west Middle Carboniferous dextral faulting being succeeded and partly overprinted by late Carboniferous faults that have a N E - S W trend and that may be either dextral or sinistral. However, there are very few faults in the Canadian Maritimes along which sinistral movements are demonstable. This is partly the result of the fact that polyphase movements occurred on such faults. For instance, the major Belleisle fault of New Brunswick has strong indications of late Carboniferous reversed movements (Brown & Helmstaedt 1970; Garnett & Brown 1973). However, their interpretation of the deformation found in the proximity of the fault arose from the assumption of a single long-term heterogenous strain. The situation along the Belleisle fault (Fig. 3) is complex and lenticles of rock of various determinable ages ranging from Cambrian to Carboniferous occur along the main trace (Potter, Hamilton & Davies 1979) with apparent downthrow being sometimes to the NE and at times to the SE. This suggest that movements along the fault were episodic and that transcurrent motions of either sinistral or dextral type are possible, especially since the overall formational plunge to the SE of the fault is fairly uniform, if gentle, to the NE (Rast 1979). Motions earlier than late Pennsylvanian on the fault are not inconsistent with sinistral shear which may have started in the late Precambrian (Rast & Dickson 1982). The Belleisle fault therefore represents a very major boundary. It so abruptly and finally cuts off magnetic anomalies of the Avalon terrane (Haworth & Lefort 1979, fig. 6) that transcurrent movement along it appears not only possible but, as judged from the magnetic anomalies map of Zietz et al. (1980), it is of the order of at least 250 km. If, as seems possible, the fault is partly equivalent to the Cabot fault in Newfoundland (Fig. 3), movements of a similar magnitude occurred there as well. At present, therefore, it is possible to accept the sinistral translation of the Avalon terrane in late Devonian-Carboniferous times as suggested by Kent (1970, 1983) and others. How all this reflects on the Great Glen fault situation should be tackled by workers in the United Kingdom.

Central and southern Appalachian orogenic belt The central and southern Appalachians were long known to have been affected by the Alleghenian-Variscan orogeny (cf. Rodgers 1970).

207

Throughout the early research, however, the degree of allochthony and the nature of crossstrike correlations was in dispute. In the southern Appalachians, three zones of varying succession, deformation and metamorphism have been recognized that NW to SE are known as: the Valley and Ridge province, the Blue Ridge and the Piedmont (Rodgers 1970; Rast 1983; Fig. 4). Of these the well-dated Carboniferous strata occur only in the Valley and Ridge province. In the central Appalachians the situation is more complex and intervening between the Valley and Ridge Province and the Blue Ridge is the so-called Great Valley province (Perry 1978). The successions and structures within these provinces vary from place to place. In this account the rocks of the Great Valley are considered together with the Valley and Ridge.

Succession The central and southern Appalachians form a wide belt that is some 2000 km long and therefore any generalizations regarding succession and structure formed suffer from imperfections. There are, however, general features that can be recognized. The NW boundary to the belt is known as the Allegheny front. Successions can be easily correlated across this between the platform and the Valley and Ridge province, although in the latter area the aggregate thickness is much greater. The sequence in the Valley and Ridge province starts with a thick succession at Lower Palaeozoic sediments that, without any appreciable break, follow up into the Devonian and Carboniferous strata. The latter, although commonly involved in overthrusting, yield fossils and have been studied in detail from Pennsylvania to Alabama. The succession near to the Allegheny front contains both Mississippian and Pennsylvanian sediments, but in general to the SE only faultbound Mississippian rocks are preserved (cf. Englund 1979; Milici et al. 1979), although in Georgia (Thomas & Cramer 1979) such is not the case. The Pennsylvanian deposits form the main coal-bearing sequence. The Mississippian sediments consist of lower carbonates and overlying elastics with the latter prograding from SE to NW as is also the case with the Pennsylvanian (Smith 1979). The Valley and Ridge province to the SE is bounded by a major fault system that is known as the Blue Ridge in the south central Appalachians, as the Great Smoky Mountains fault in Tennessee, as the Cartersville fault in Georgia and as the Talladegha fault in Alabama. In the area to the SE of this fault system there are no Carboniferous

208

N. Rast

sediments, but the ground has numerous granite and gabbro intrusions to which Carboniferous ages have been adduced. As pointed out by Sinha & Zietz (1982) these intrusions are essentially in place and possibly represent a Carboniferous island arc. The ages of these plutons fall between 330 and 260 Ma and both deformed and undeformed intrusions are recognized. There is a record of strongly deformed gneissoze granites within the group (Snoke, Kish & Secor 1980). Therefore, there is little doubt that although no Carboniferous sediments are present in the Blue Ridge and Peidmont provinces, Alleghanian deformation was just as unimportant as in the Valley and Ridge. Structure The southern Appalachian structure is dominated by overthrusting. This has long been recognized, but from time to time contrasted with the central Appalachians where, although the belt as a whole is much narrower, folding is supposed to predominate (King 1977, p. 46). Furthermore in the south some thrusting took place even to the NW of the Valley and Ridge province (Wiltschko 1981). The orogenic front, which is in the south fault-determined, becomes a fold (Allegheny front of Maryland and Pennsylvania) and then entirely dies out as a single surface structure. Thus the Alleghenian thrusting becomes less significant to the north. In the central part of the southern Appalachians the critical research has been conducted on the nature of the Allegheny front in Pennsylvania and New York by Geiser (1980), Engelder & Geiser (1981) and further into the Blue Ridge by Mitra & Elliott (1980). In Pennsylvania and New York the model proposed by Engelder & Geiser (1981) consists of major blind dislocations (d6collements) which project to the NW and north at the level of Silurian salts generating brittle fracture in limestones and fold deformation and cleavage in the higher Upper Devonian shales. These dislocations are assumed to be Alleghenian. The allochthonous sheets behaved brittly at the base and were ductile at the top. Therefore, a contemporaneity of early ductile effects and certain joints and tear faults is claimed, while the later joints originated as a result of uplift and erosion, leading to relaxation. The reorientation of joints was related to reorientation of movement on the plane of blind slip (d6collement). Mitra & Elliott (1980) correlate cleavage developed in the Blue Ridge of northern Virginia, Maryland and parts of Pennsylvania with deformation that is produced in thrust sheets

passing over ramps. This model carries the presupposition that the Blue Ridge was largely allochthonous. The notion of allochthony was specially confirmed by major projects of seismic reflection studies in the central part of the southern Appalachians by the so-called C O C O R P (Cook et al. 1979) group and the Geological Survey (Harris et al. 1981) profiles (Fig. 4). Although seismic profiles are interpretative, there seems to be little doubt that they reflect the allochthony of most of the southern Appalachians on a master d6collement. In fact, in an interpretation of the Alabama traverse Neathery & Thomas (1983) follow the same precept as does recent work in north Central Virginia (Glover & Costain 1982). There are two basic problems arising out of the above findings. The first is how far does the allochthony continue a cross-strike and secondly how far does it run along strike, either into the northern Appalachians or into the Ouachita belt? There are several different (Fig. 7) interpretations as to what happens in the southern Appalachians west of the Brevard fault. The interpretation favoured by Cook et al. (1979, 1981) involves the rooting of the overthrusts under the coastal plain, with the surface of basal d6collement being the Cambro-Ordovician sediments under the Valley and Ridge, Blue Ridge and eastern Piedmont, but the situation further SE is complex and is attributed to a number of possibilities. A second possibility (Hatcher & Zietz 1980) is that the Blue Ridge and inner Piedmont allochthons root at the junction of the inner and eastern Piedmont in the so called King's Mountain belt. Lastly, Harris & Bayer (1979) and Harris et al. (1981) have an interpretation in which the d6collement continues under the coastal plain and even the continental shelf. All these interpretations involve displacements of a few hundred kilometres. Thus, a substantial displacement must be accepted. However, the claim that all the thrusting has been thin skinned cannot be entertained since large slices of Precambrian (Grenville?) rocks are involved. Cook et al. (1979) and Harris & Bayer (1979) assumed that the allochthony identified in the southern Appalachians continues into the northern Appalachians. They did not take into account the geology of the northern Appalachians or the timing of the events. Williams (1980) has pointed out that there are basic dissimilarities between the geology of the southern and northern Appalachians, and namely (among others): (1) there is no Valley

A l l e g h e n i a n o r o g e n y in A m e r i c a

V&R

BR

B

209

P

-a

V&R

V&R

BR

S

BR

P

8

P

FIG. 7. Models of conceptual geometry in the southern Appalachians: (a) according to Cook et ai. (1979, 1980); (b) according to Hatcher & Zietz (1980); (c) according to Harris & Bayer (1979). M--Moho discontinuity, B--Brevard fault zone, BR--Blue Ridge, O--multiple reflections suggesting obduction, P--Piedmont, V&R--valley and ridge. Thick lines indicate the position of major d6collements, arrows direction of movement. Continental crust in crosses. Distance between platform and coastline is approximately 500 km+. Transitional crust in dashed ornament. and Ridge province in the northern Appalachians; (2) dated Silurian, Devonian and Carboniferous rocks occur across the orogen and impose limitations on thrusting; (3) the westernmost allochthonous rocks in the northern Appalachians are affected by the 7aconian deformation and are bounded by the qaconian front; (4) no basal masterd6collement can be projected across Newfoundland; (5) there is a gradual eastward increase in metamorphism in the metamorphic belt of Newfoundland and this by implication is even against the idea of Taconian wholesale thrusting in the interior of the orogen. The replies that were given by Cook et al. (1980) and Harris & Bayer (1980), did not dispose of Williams' objections. There was a point made that the profiles cut by SOQUIP across the Taconian orogen of Quebec yielded a style similar to that in the southern Appalachians. Yet since the date of overthrusting has not been determined the similarity is no more than the similarity of style. Secondly, it was suggested that surface arrangements of rocks in an allochthonous belt form incomplete indications of the structure at depth. While this is so, the allochthony of the southern Appalachians was largely hypothesized on the basis of surface exposures (Jonas 1929), and no such hypothesis evolved in the northern Appalachians because the surface evidence was insufficient. Furthermore, as already described, almost entirely undeformed Carboniferous strata are exposed from south central New Brunswick to the Gasp6 Peninsula

of Quebec (a half of the width of the orogen). The only evidence for Alleghanian deformation is found in Maritime Canada and therefore it is this belt that should be correlated with the southern Appalachian thrust conditioned Alleghenian orogen. The relationship of thrusting to metamorphism in the southern and northern Appalachians is very different. In the north Lower-Middle Palaeozoic strata, while metamorphosed, are not affected by much postmetamorphic thrusting, whereas in the south (Harris etal. 1981) isograds are displaced by the thrusts of the Alleghenian orogeny. The only exception to the general rule in the north is in the Maritimes, where not only the Carboniferous sediments, but also the Precambrian basement, is involved in allochthonous displacements. Hatcher & Odom (1980) recognized that in the southern Appalachians there are tectonic slides of Taconian and Acadian events, but that these, especially to the NE are cut by Alleghenian thrusts, which themselves can be of several ages, beginning with the Goat Rock-Bartlett Ferry (c. 380 Ma) and Brevard (356---20 Ma ages) and ending with 260 Ma events involving metamorphism (Kish et al. 1978). Sinha & Zietz (1982) date late Carboniferous-Permian events as lying in the range of 240-270 Ma. Lastly, the palaeomagnetic evidence for large-scale transcurrent movement in the northern Appalachians is not identifiable in the south. Therefore, all in all, at present, the

210

N. R a s t

extension of Alleghenian structures from the southern to the northern Appalachians in a straightforward sense of relating the Allegheny front to the Logan line of the northern Appalachians (Fig. 3) is impossible;the only possibility being that the Variscan front of New Brunswick-New England correlates with the Allegheny front, but that it does not mean that there has necessarily been a continuous fracture, rather that the front is marked by severe but varying strain. In addition, the late Acadian to early Variscan deformation leading to transcurrent movement of the Acadia in the southern Appalachians took place to the east of the present-day coast of North America. If this is so, then the along strike correlation of the Avalon terrane of the northern and southern Appalachians is not valid. The Alleghenian orogeny sensu stricto was entirely later than the transcurrent movement and involved the formation of an island arc (Sinha & Zietz 1982) and presumably a collision with Africa. By that time the Avalon terrane of the northern Appalachians was already in the same position relative to North America as at the present time.

cian-Lower Mississippian shales, micrites, bedded chert and a peculiar silicified carbonate mud known as novaculite. These sediments have been traced widely along the strike. The flysch in this zone is very thick and some of it is possibly structurally repeated (Lillie et al. 1982). There are in places small pods of serpentine within the succession (Sterling & Stone 1961). The topmost Carboniferous sediments are represented by a post-orogenic molasse. In the southernmost zone flysch continues and has included in its beds of rhyodactitic tufts. The deformation is relatively slight, but the gently dipping and folded strata are unconformably overlain by Upper Pennsylvanian carbonates--effectively undeformed platformal sediments. The unconformity marks the Ouachitan orogeny. Viele (1979) interpreted the sequences of the three zones such that the northern zone represents the shelf-slope deposits of the North American plate, the southern zone a separate microcontinent (Llanoria), and the intervening Benton uplift as the zone of suturing--a collision between North America and Llanoria.

Structure The O u a c h i t a s Traditionally the Ouachitas including the Marathon region were treated as a separate orogenic belt from the Appalachians although in recent years Thomas (1972, 1977) persuasively advocated a buried continuity between the two systems. However, there are general correlational and structural difficulties. These are discussed below.

Succession The succession and structure of the Ouachitas in recent years has been described by King (1975). In Arkansas three main east-west trending zones can be recognized: the northern (Arkoma basin), the central (Benton uplift) and the southern (Marzarn-Trap Mountains). The succession in these three zones is different. In the northern A r k o m a basin Lower and Middle Palaeozoic carbonates are overlain unconformably by the essentially Lower Pennsylvanian clastic sediments that thicken into the basin and upward into massive turbidites, interpreted as flysch and wildflysch. To the south of the latter, in the Benton uplift, there is a zone in which the flysch overlies a very thin sequence of Ordovi-

The structure of the best studied profiles in Arkansas supports Viele's (1979, 1982) contention as indeed does the recent C O C O R P transect (see Brewer 1984). The rocks of the Benton uplift zone are strongly deformed and their northern margin is marked by a well developed 'frontal' north-directed thrust belt. Thrusting on a smaller scale continues into the A r k o m a basin, but dies out relatively rapidly to the south. In the Benton zone deformation is polyphase and the initial thrusts are followed by a reversal of vergence and also backthrusting. In this sense the sequence is similar to the northern Appalachians. There is little igneous activity in the orogen and Carboniferous granites are not recorded. In this sense the Ouachitan belt is different from the Appalachian. Metamorphism is generally low grade and can be attributed to two separate episodes, one at about 380 Ma and another at 3 2 0 - 2 5 0 Ma (Denison 1982) and is therefore similar to the southern Appalachians. Thomas (1982) suggests that the Ouachitan belt is separated from the Appalachians by a palaeotransform fault, and therefore the differences in tectonic style and sedimentation between the southern Appalachians and the Ouachitas should be attributed to events taking place in relatively

Alleghenian orogeny in America distant domains. Furthermore, the main deformation in the Ouachitan belt took place in the Middle Pennsylvanian times while in the southern Appalachians it is end-Pennsylvanian suggesting that the collisional events that presumably gave rise to the two orogens occurred in different periods with the Ouachitan being earlier than the Alleghenian.

Summary of Variscan events The Appalachian-Ouachitan belts of North America are interpreted as collisional orogenic belts. The collisional events began in both at approximately the same time (late Devonian-early Mississippian). Both orogenic belts represent portions of shelf-slope sediments and of the advancing plates that acccreted to the North American plate. The two belts can be related to the European Armoricanides and Hercynides and the three segments collectively designated as the Variscan orogen. The differences in the chronology of the three major segments of the Variscan orogeny can be interpreted as collisions with different plates. The collision with the North European craton probably took place with numerous microplates between the African and European cratons (Badham 1981; Matte 1983). The collision with the Appalachian orogen took place between the African and North American plates with the Avalonian (Acadia) fragment intervening. Hence Schenk (cf. 1978 for full bibliography) long maintained that southern Nova Scotia is a fragment of Africa. Since it is separated from New Brunswick by the Variscan frontal thrusts it may be entirely allochthonous. The intervention of the Avalonian fragment accounts for the differences between the northern and southern Appalachians. Lastly Llanoria may have been a fragment of South America that at the time was adjacent to the southern United States. Irving (1977) in fact thought that the Appalachian belt was generated by such a collision; computerized modelling (cf. Ziegler et al. 1979) tends to deny it, although the resulting maps for the late Westphalian do juxtapose South America against the Gulf of Mexico. The late Westphalian to Lower Permian times also saw the development of late Variscan, post-thrusting faults, recognized in Maritime Canada (Lefort & Van der Voo 1981) and geophysically related by tracing these faults across the fitted present-day continental shelves of the North Atlantic (Lefort & Haworth 1978;

211

Haworth 1981; Lefort & Haworth, this volume). These structures represent late shearing of essentially right lateral sense (Arthaud & Matte 1977) on the plate scale, but left lateral on the local scale as they affect late Carboniferous sediments. In terms of the stratigraphical chronology developed in the northern Appalachians these faults are post-Stephanian. The interactions of major plates produced megascale compressive structures in the Variscan belt because the collisions were between blocks of continental dimensions. Thus the sutures were covered and the remnants of oceanic plates were hidden under the overthrusts and do not occur on the surface. Cook & Oliver (1981) have suggested that the major trend-parallel gravity anomaly in the Piedmont of Georgia, South Carolina and North Carolina, where a gradient of 20 mgal exists, represents the hidden ancient edge of the continent. Their profile shows a large number of inclined reflectors. I suggest that this struc-

N Ar~e,r'~c~

I~

PP

FIG. 8. Plate tectonic interpretation and the development of thrusting shown in time-sequence diagrams. M--Mississippian, MP--Middle Pennsylvanian, PP--late Pennsylvanian-Perrnian. Solid triangle--direction of movement, thick lines with solid triangles--Carboniferous overthrusting, lines with open triangles--Ordovician-Acadian overthrusting, transcurrent faulting shown by arrows.

212

N. Rast

ture is an o b d u c t e d Variscan ophiolite m a r k i n g the b u r i e d s u t u r e (Fig. 7), as h a d b e e n first p r o p o s e d by H a t c h e r & Zietz (1980). T h e s e q u e n c e of events thus w o u l d involve (Fig. 8): (1) T r a n s c u r r e n t coast parallel faulting in Mississippian times g e n e r a t i n g clastics of 'molasse' type. L l a n o r i a as a part of S o u t h A m e r i c a plate m o v e s t o w a r d N o r t h A m e r i c a and starts shedding flysch. (2) T h e d e v e l o p m e n t and a possible collision of the early P e n n s y l v a n i a n volcanic arc, w h e n t r a n s c u r r e n t m o v e m e n t s w e r e r e p l a c e d by subduction, with s o m e c o n c o m i t a n t o v e r t h r u s t i n g in the A p p a l a c h i a n s a n d the m a i n o r o g e n i c m o v e m e n t in the O u a c h i t a s w h e r e L l a n o r i a collided with the N o r t h A m e r i c a n plate by M i d - P e n n s y l v a n i a n times. (3) T h e collision of the A f r i c a n a n d N o r t h A m e r i c a n plates in the n o r t h e r n a n d s o u t h e r n

Appalachians in the late Carbonifero u s - P e r m i a n time, a possible tensional separation of L l a n o r i a f r o m South A m e r i c a owing to the c h a n g e in regional o r i e n t a t i o n of direction of p l a t e m o v e m e n t .

ACKNOWLEDGMENTS AND APOLOGIA; I acknowledge the help of Rebecca Meacham who chased the references and prepared the manuscript. Typing was done by Rebecca Meacham and Connie Irvine. The preparation was partly financed from generous donations from Standard Oil of California (Chevron). The scope of the region is so vast that inevitably many papers quoted are of a collative nature. To all whose research I have left unacknowledged I offer my apologies, but in the context of this account there was only limited space.

References ARTHAUD, F. & MATYE, P. 1977. Late Palaeozoic strike-slip faulting in southern Europe and northern Africa: result of a right-lateral shear zone between the Appalachians and the Urals. Bull. geol. Soc. Am. 88, 1305-20. BADHAM, J. P. N. 1981. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 495-506. BArnEY, E. B. 1929. The Palaeozoic mountain systems of Europe and America. Rep. Br. Ass. Advmt Sci. 811-4. BELL, K. & BLENKINSOP, J. 1977. Geochronological evidence of Hercynian activity in Newfoundland. Nature, 265, 616-8. BELL, W. A. 1929. Horton-Windson District, Nova Scotia. Mem. geol. Surv. Can. 155. 1938. Fossil flora of Sydney coal field, Nova Scotia. Mere. geol. Surv. Can. 215, 334 pp. -1940. The Pictou coal field, Nova Scotia. Mere. geol. Surv. Can. 225, 160 pp. BELT, E. S. 1968. Post-Acadian rifts and related facies--Eastern Canada. In: ZEN, E. et al. (eds) Studies o f Appalachian Geology, 95-113. Wiley, New York. BRowN, L. & HELMSqAEWI, H. 1970. Deformation history in part of the Lubec-Belleisle zone of southern New Brunswick. Can. J. Earth Sci. 7, 748-67. & KELLEY, W. M. 1980. Paleomagnetic results from Northern Maine--reinterpretations. Geophys. Res. Lett. 7, 1109-11. BRUNO, A. C. 1982. Derivation of synorogenic clasties in the Mississippian and Pennsylvanian strata of southern West Virginia. Unpublished M.S. Thesis. University of Kentucky. 137 pp. -

-

CAMERON, B. (ed) 1979. Carboniferous basins of. southeastern New England. Guidebk Field Trip No. 5, Ninth int. Congr. Carbonif. Strat. Geol. American Geological Institute. 170 pp. COOK, F. A., ALBAUGH,D. S., BROWN,L. D., KAUFMAN, S., OLIVER, J. E. & HATCHER, R. D. 1979. Thin-skinned tectonics in the crystalline southern Appalachians: COCORP seismic-reflection profiling of the Blue Ridge and Piedmont. Geology, 7, 563-7. & 1980. Reply. Geology, 8, 213-4. --, BROWN, L. D., KAUFMAN,S., OLIVER, J. E. & PETERSON,R. A. 1981. COCORP seismic profiling of the Appalachian orogen beneath the Coastal Plain of Georgia. Bull. geol. Soc. Am. 10, 738-48. & OLIVER, J. E. 1981. The late Precambrian-early Palaeozoic continental edge in the Appalachian orogen. Am. J. Sci. 281, 993-1008. CURRIE, K. L. 1977. A note on post-Mississippian thrust faulting in north-western Cape Breton Island. Can. J. Earth Sci. 14, 2937-41. DALLMEYER, R. D. 1982. 4~ ages from the Narragansett Basin and southern Rhode Island basement terrane: their bearing on the extent and timing of Alleghanian tectonothermal events in New England. Bull. geol. Soc. Am. 93, 1118-30. , BLACKWOOD,R. F. & ODOM, A. L. 1981. Age and origin of the Dover Fault: tectonic boundary between the Gandar and Avalon zones of the northeastern Appalachians. Can. J. Earth Sci. 18, 1431-42.

A l l e g h e n i a n o r o g e n y in A m e r i c a DENISON, R. E. 1982. Ages of Ouachita Metamorphism. Abstr. Progr. geol. Soc. Am. 14, 475. DENNISON, J. M. (compiler) 1977. Bibliographic index of Appalachian Basin, theses and dissertations through 1975. Appal. geol. Soc. Charleston, W. Va., 295 pp. DONALDSON, A. C. & SHUMAKER, R. C. 1981. Late Paleozoic Molasse of Central Appalachians. In: MIALL, A. D. (ed.) Sedimentation and Tectonics in Alluvial Basins. Pap. geol. Ass. Can. 23, 99-124. DOYLE, R. G. 1979. Rocks of Maine inferred to be Carboniferous. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Massachusetts, Rhode Island and Maine. Prof. Pap. U.S. geol. Surv. IlIO-A, A24-5. DRAKE, C. L. & WOODWARD, H. P. 1963. Appalachian curvature, wrench faulting and offshore structures. Trans. N. Y. Acad. Sci. ser. 2,26, 48-63. EDMUNDS, W. E., BERG, T. M., SEVON, W. D., PIOTROWSKI, R. C., HEYMAN, L. & RICKARD, L. V. 1979. Pennsylvania and New York. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States. Prof. Pap. U.S. geol. Surv. 1110-B, B1-33. EISBACHER, G. H. 1970. Deformation mechanics of mylonitic rocks and fractured granites in Cobequid Mountains, Nova Scotia, Canada. Bull. geol. Soc. Am. 81, 20 009-20. ENGELDER, T. & GEISER, P. 1981. On the use of regional joint sets as trajectories of paleostress fields during the development of the Appalachian Plateau, New York. J. geophys. Res. 85, 6319-41. ENGLUNDE, K. J. 1979. The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Virginia. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States. Prof. Pap. U.S. geol. Surv. 1110-C, C1-21. FERM, J. C. 1970. Allegheny deltaic deposits. In: MORGAN, J. P. (ed.) Deltaic Sedimentation Modern and Ancient. Spec. Pubis Soc. econ. Paleont. Miner., Tulsa, 15, 246-55. -1971. Carboniferous depositional environments in northeastern Kentucky. Geological Society of Kentucky. 29 pp. FONG, C. C. K. 1975. Geological mapping--Deer Lake-White Bay Carboniferous basins. In: Report of Activities, 1975. Rep. 76-1, 19-24. Mineral Development Division, Province of Newfoundland, Department of Mines and Energy. FRALICK, P. W. & SCHENK, P. E. 1981. Molasse deposition and basin evolution in a wrench fault setting: the Late Paleozoic, eastern Cumberland basin, Maritime Canada. In: MIALL, A. D. (ed.) Sedimentation and Tectonics in Alluvial Basins. Spec. Pap. geol. Ass. Can. 23, 77-98. GARNETr, J. A. & BROWN, R. L. 1973. Fabric variation in the Lubec-Belleisle zone of southern New Brunswick. Can. J. Earth Sci. 10, 1591-9.

213

GEISER, P. A. 1980. Cleavage in Lower and Middle Devonian rocks of the Hudson and Delaware River Valleys: its implication for Appalachian tectonics. Abstr. Progr. geol. Soc. Am. 12, 37. GLOVER, L. I I I & COSTAIN, J. K. 1982. Vibrosies reflecting seismic structure along the Blue Ridge and Piedmont James River profile, north central Virginia. Abstr. Prog. geol. Soc. A m . 14, 497. GOLDSTEIN, A. G. 1982. Geometry and kinematics of ductile faulting in a portion of the Lake Char mylonite zone, Massachusetts and Connecticut. Am. J. Sci. 282, 1378-1405. Gussow, W. C. 1953. Carboniferous stratigraphy and structural geology of New Brunswick, Canada. Bull. Am. Ass. Petrol. Geol. 37, 1713-1816. HANCOCK, P. L., DUNNE, W. M. & TRINGHAM, M. E. 1983. Variscan deformation in southwest Wales. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 47-73. Hilger, Bristol. HARRIS, L. D. & BAYER, K. C. 1979. Sequential development of the Appalachian orogen above a master decollement--a hypothesis. Geology, 7, 568-72. &~ 1980. Reply. Geology, 8, 214. --, HARRIS. A. G., DE WITT, W. (JR) & BAYER, K. C. 1981. Evaluation of southeastern overthrust belt beneath Blue Ridge-Piedmont thrust. Bull. Am. Ass. Petrol. geol. 65, 2497-505. HATCHER, R. D. & ODOM, A. L. 1980. Timing of thrusting in the southern Appalachians, USA: model for orogeny? J. geol. Soc. London, 137, 321-7. -& ZlETZ, I. 1980. Tectonic implications of regional aeromagnetic and gravity data from the southern Appalachians. In: WONES, D. R. (ed.) The Caledonides in the USA. IGC Proj. 27, Caledonide Orogen, Va. Poly. Inst. & St. Univ., Mem. Dept. geol. Sci. No. 2, Blacksburg. 235-44. HAWORTH, R. T. 1981. Geophysical expression of Appalachian-Caledonide structures on the continental margins of the North Atlantic. Mere. Can. Soc. Petrol. Geol. 7, 429-46. & LEEORT, J. P. 1979. Geophysical evidence for the extent of the Avalon zone in Atlantic Canada. Can. J. Earth Sci. 16, 552-67. HEPBURN, J. C. 1979. Deposits near Worcester, Mass. In: The Mississippian and Pennsylvanian (Carboniferous) System in the United States--Massachusetts, Rhode Island and Maine. Prof. Pap. U.S. geol. Surv. l l l O - A , 15-20. HOBSON, D. M. & SANDERSON, D. J. 1983. Variscan deformation in southwest England. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles, 108-30. Hilger, Bristol. HOWlE, R. D. & BARSS, M. S. 1975. Paleogeography and sedimentation in the Upper Paleozoic, eastern Canada. In: YORATH, C. J., PARKER, E. R. & GLASS, D. J. (eds) Canada's Continental Margins. Mem. Can. Soc. Petrol. Geol. 45-58. IRVING, E. 1977. Drift of the major continental blocks since the Devonian. Nature, 270, 304-9. JONAS, A. 1929. Structure and metamorphic belts of

214

N. R a s t

the Southern Appalachians. Am. J. Sci. 24, 228-43. KAYE, C. A. & ZARTMAN, R. E. 1980. A Late Proterozoic A to Cambrian age for the stratified rocks of the Boston Basin, Massachusetts, U.S.A. In: WONES, D. R. (ed.) The Caledonides in the USA. IGC Proj. 27: Caledonide Orogen, Mere. Dept geol. Sci. No. 2, 257-62. Va. Poly Inst. & St. Univ., Blacksburg. KELLEY, D. G. 1967. Some aspects of Carboniferous stratigraphy and depositional history in the Atlantic provinces of Canada. Spec. Pap. geol. Ass. Can. 4 , 2 1 3 - 2 8 . 1970. Carboniferous and Permian. In: Geology of Southeastern Canada: Geology and Economic Minerals o f Canada. Mere. Can. geol. Surv. 229-304. KENT, D. V., 1983. Paleomagnetic evidence for Post-Devonian displacement of the Avalon Platform (Newfoundland). J. geophys. Res. 87, (B10), 8709-16. KENT, D. V. & OPDYKE,N. D. 1978. Paleomagnetism of the Devonian Catskill red beds; evidence for motion of the coastal New England-Canadian Maritime region relative to cratonic North America. J. geophys. Res. 83, 4441-50. & -1979. Paleomagnetism of the Barnett Formation (Texas) and its bearing on Carboniferous tectonics of eastern North America (abstr.). Los, 60, 746. KEPP1E, J. D. 1980. An introduction to the geology of Nova Scotia. In: RAST, N. (ed.) The Northern Appalachian Geotraverse: Quebec-New Brunswick-Nova Scotia. Field Trip Guidebk geol. Ass. Can. 107-15. KING, P. B. 1959. The Evolution of North America. Princeton University Press. 190 pp. 1975. Ancient southern margin of North America. Geology, 3, 732-4. 1977. The Evolution o f North America (revised edition). Princeton University Press. 197 pp. KISH, S., FULLAGAR, P. D., SNOKE, A. W. • SECOR, D. T. R. (JR) 1978. The Kiokee belt of South Carolina (part I): evidence for late Paleozoic deformation and metamorphism in the southern Appalachian Piedmont. Abstr. Progr. geol. Soc. Am. 10, 172-3. KNIGHT, I. 1975. Geological mapping, Carboniferous rocks, Cape Anguille Mountains, southwest Newfoundland. Prov. Newfndld Miner. Dev. Div. Rep. 75-1, 31-41. KOCIS, D. E., HERMES, O. D. & CAIN, J. A. 1978. Petrologic comparison of the pink and white facies of the Narragansett Pier granites, Rhode Island. Abstr. Progr. geol. Soc. Am. 10, 71. & MURRAY, D. P. 1977. Re-evaluation of Late Paleozoic igneous activity and accompanying contact metamorphism in southeastern New England. Abstr. Progr. geol. Soc. Am. 9, 296-7. KRYNINE, P. D. 1948. The megascopic study and field classification of sedimentary rocks. J. Geol. 56, 130-65. LARRABEE, D. M., SPENCER, C. W. & SWIFT, D. J. P. 1965. Bedrock geology of the Grand Lake area, -

-

Aroostook, Hancock, Penobscot, and Washington counties, Maine. Bull. U.S. geol. Surv. 1201-E, 38 pp. LEFORT, J.-P. & HAWORTH, R. T. 1978. Geophysical study of basement fractures on the western European and eastern Canadian shelves: transatlantic correlations and late Hercynian movements. Can. J. Earth Sci. 15, 397-404. & VAN DER Voo, R. 1981. A kinematic model for the collision and complete suturing between Gondwanaland and Laurassia in the Carboniferous. J. Geol. 89, 537-50. LEGUN, A. S. & RUST, B. R. 1982. The Upper Carboniferous Clifton Formation of northern New Brunswick: coal-bearing deposits of a semi-arid alluvial plain, Can. J. Earth Sci. 19, 1775-85. LILLIE, R. J., NELSON, K. D., DE VOOYD, B., OLIVER, J., BROWN, L. D. & KAUFMAN, S. 1982. Crustal structure of the Ouachita Mountains, Arkansas, from COCORP seismic profiles and regional gravity data. Abstr. Progr. geol. Soc. Am. 14, 547. LUDMAN, A. 1981. Significance of transcurrent faulting in eastern Maine and location of the suture between Avalonia and North America. Am. J. Sci. 281, 463-83. LYONS, P. C. 1976. Implications of coral flora and stratigraphy of the Narragansett Basin, New England on plate tectonic models. Abstr. Progr. geol. Soc. Am., 8, 989-90. 1979. Biostratigraphy of the Pennsylvanian of Massachusetts and Rhode Island. In: The Mississippian and Pennsylvanian Systems in the United States--Massachusetts, Rhode Island, and Maine. Prof. Pap. U.S. geol. Surv. 1110-A, A20-4. MAT-rE, P. 1983. Two Geotraverses across the Ibero-Armorican Variscan Arc of Western Europe. J. Am. geophys. Un. in press. MICHARD, m. & PIQUE, A. 1980. The Variscan Belt in Morocco: structure and developmental model. In: WONES, D. R. (ed.) Caledonides in the USA. IGC Proj. 27: Caledonide Orogen, Mem. Dept. geol. Sci. No. 2, 317-22. Va. Poly. Inst & St. Univ., Blacksburg. MILICI, R., BRIGGS, G., KNOX, L. M., SITrERLY, P. D. & STATLER, A. T. 1979. The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Tennessee. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States. Prof. Pap. U.S. geol. Surv. lllO-G, G1-38. MITRA, G. & ELLIOYr, D. 1980. Deformation of basement in the Blue Ridge and the development of the South Mountain cleavage. In: WONES, D. R. (ed.) The Caledonides in the USA. IGC Proj. 27: Caledonide Orogen, Dept. geol. Sci. Mem. No. 2, 307-12. Va. Poly. Inst. & St. Univ., Blacksburg. MOORE, B. R. & CLARKE, M. K. 1970. The significance of a turbidite sequence in the Borden formation of Eastern Kentucky and southern Ohio. Spec. Pap. geol. Ass. Can. 7, 211-18. MORRIS, W. A. 1976. Transcurrent motion determined paleomagnetically in the Northern

A l l e g h e n i a n o r o g e n y in A m e r i c a Appalachians and Caledonides and the Acadian Orogeny. Can. J. Earth Sci. 13, 1236-43. MUTCH, T. A. 1968. Pennsylvanian nonmarine sediments of the Narragansett basin, Massachus e t t s - R h o d e Island. In: KLEIN, G. DE V. (ed.) Late Paleozoic and Mesozoic Continental Sedimentation, Northeastern North America. Spec. Pap. geol. Soc. Am. 106, 177-209. NEATHERY, T. & THOMAS, W. 1983. Geodynamic transect of the Appalachian orogen in Alabama. J. Am. geophys. Un. in press. PARNELL, J. 1982. Comment: comments and reply on Paleomagnetic evidence for a large (2000 km) sinistral offset along the Great Glen fault during Carboniferous time. Geology, 10, 605. PERRY, W. J. (JR) 1978. Sequential deformation in the Central Appalachians. Am. J. Sci. 278, 518-42. -& DE WITT, W. (JR) 1977. A Field Guide to Thin-Skinned Tectonics in the Central Appalachians. Am. Ass. Petrol. Geol.--Soc. econ. Paleont. Miner. Ann. Convention, Washington DC, Field Trip 4, 54 pp. POOLE, W. H. 1967. Tectonic e-volution of Appalachian regions of Canada: geology of the Atlantic region. Spec. Pap. R. Soc. Can. 4, 9-51. POTTER, R. R., HAMILTON, J. B. & DAVIES, J. L. 1979. Geological Map, New Brunswick, 2nd ed. New Brunswick Department of Natural Resources. QUINN, A. W. 1971. Bedrock geology of Rhode Island. Bull. U.S. geol. Surv. 1295, 1-65. 8(: MOORE, G. E. 1968. Sedimentation, tectonism, and plutonism of the Narragansett basin region. In: ZEN, E. et al. (eds) Studies of Appalachian Geology: Northern and Maritime. 269-79. Wiley, New York. RAST, N. 1979. Precambrian metathiabases of southern New Brunswick--the opening of the Iapetus Ocean? Tectonophys. 59, 127-37. 1983. Variscan orogeny. In: HANCOCK, P. L. (ed.) The Variscan Fold Belt in the British Isles. 1-19. Hilger, Bristol. 1984. The northern Appalachian traverses in the Maritimes of Canada. J. Am. geophys. Un. in press. & D1CKSON, W. L. 1982. The Pocologan mylonite zone. In: ST JULIEN, P. & BELIND, J. (eds) Major Structural Zones and Faults of the Northern Appalachians. Spec. Pap. geol. Ass. Can. 24. • GRANT, R. 1973a. Transatlantic correlation of the Variscan-Appalachian Orogeny. Am. J. Sci. 273, 572-9. &~ 1973b. The Variscan Front in Southern New Brunswick. In: RAST, N. (ed.) Geology o f New Brunswick, Field Guide to Excursions. NEIGC, 4-11. &~ 1977. Variscan-Appalachian and Alleghanian deformation in the Northern Appalachians. Colloquium volume: contributions to International Colloquium on the Variscan System. Rennes, France, 1974. Colloques. int. Cent. natn. Rech. scient. No. 243,583-6. --, --, PARKER, J. S. D. & TENG, H. C. i978.

215

The Carboniferous in deformed rocks west of Saint John, New Brunswick. In: LUDMAN, A. (ed.) Guidebk for Field Trips in Southern Maine and Southwestern New Brunswick. 70th Ann. Meeting NEIGC, 162-73. , SKEHAN,J. W. (S.J.) & GRANT, R. H. 1976. The Variscan-Alleghanian-Appalachian orogenic movement in North America and Europe (abstr.). Geol. Soc. Am. Northeastern Section 11th Ann. Meeting, Southeastern Section 25th Ann. Meeting, 251-2. ROBERTS, J. L. & SANDERSON,D. J. 1971. Polyphase development of slaty cleavage and the confrontation of facing directions in the Devonian rocks of north Cornwall. Nature, 230, 87-9. ROOGERS, J. 1970. The Tectonics o f the Appalachians. Wiley, New York. RUITENBERG, A. A. 1963. Potential ore mineralization and alteration of the Mount Pleasant tin prospect, Charlotte County, New Brunswick. Unpublished thesis. University of New Brunswick, Fredericton, Canada. ~, MCCUTCHEON, S. R., VENUGOPAL, D. V. & PIERCE, G. A. 1977. Mineralization related to Post-Acadian tectonism in southern New Brunswick. Geosci. Can. 4, 13-22. RUST, B. R.-1981. Alluvial deposists and tectonic style: Devonian and Carboniferous successions in eastern Gaspe. In: MIALL, A. D. (ed.) Sedimentation and Tectonics in Alluvial Basins. Pap. geol. Ass. Can. 23, 49-76. SCHENK, P. E. 1967. The significance of algal stromatolites to paleoenvironmental and chronostratigraphic interpretations of the Windsorian stage (Mississippian). Spec. Pap. geol. Ass. Can. 4,229-45. 1970. Regional variation of the flysch-like Meguma Group (Lower Paleozoic) of Nova Scotia compared to recent sedimentation off the Scotian shelf. Spec. Pap. geol. Ass. Can. 7, 127-53. - 1976. A regional synthesis (of Nova Scotia geology). Marit. Sed. 12, 17-24. - 1978. Synthesis of the Canadian Appalachians. In: TOZER, E. T. & SCHENK, P. E. (ed.) Caledonian-Appalachian Orogen o f the North Atlantic Region. Pap. geol. Surv. Can. 7 8 - 1 3 , 1 1 1 - 3 6 . SCRUTrON, R. A. 1979. On sheared passive continental margins. Tectonophys. 59, 293-305. SHUMAKER, R. C. 1975. Relationships of Paleozoic taphrogeny and orogeny in the midcontinent region. Bull. Am. Ass. Petrol. Geol. 59, 1739. SIMPSON, R. W., BOTHNER, W. A. & SHRIDE, A. F. 1980. Offshore extension of the ClintonNewbury and Bloody Bluff fault systems of Northeastern Massachusetts. In: WONES, D. R. (ed.) Caledonides of the USA. IGC Proj. 27: Caledonide Orogen. Dept geol. Sci. Mere. 2, 229-34. Va. Poly. Inst. & St. Univ., Brunswick. SINHA,A. K. & ZIETZ,I. 1982. Geophysical and geochemical evidence for a Hercynian magmatic arc, Maryland to Georgia. Geology, 10, 593-6. SKEHAN, J. W. (S.J.) & MURRAY, D. P. 1979a. Geology of Narragansett Basin, Southeastern Mas-

216

-

N. Rast

sachusetts and Rhode Island. In: CAMERON,B. (ed.) Carboniferous Basins of Southeastern New England. Field Guidebk. Trip No. 5, 9th int. Congr. Carboniferous Stratigraphy and Geology, 7-35. &~ 1979b. Narragansett Basin of Massachusetts and Rhode Island. In: 7he Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Massachusetts, Rhode Island and Maine. Prof. Pap. U.S. geol. Sure. 1110-A, A2-13. & 1980a. A model for the evolution of the eastern margin (EM) of the Northern Appalachians. In: WONES, D. R. (ed.) Caledonides in the USA. IGC Proj. 27: Caledonide Orogen. Dept geol. Sci. Mern. No. 2, 229-34. Va. Poly. Inst. & St. Univ., Blacksburg. & ~ 1980b. Geologic profile across southeastern New England. Tectonophys. 69, 285-320. - - , HEPBURN, J. C., BILLINGS, M. P. LYONS, P. C. & DOYLE, R. H. 1979. The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Massachusetts, Rhode Island and Maine. In: The Mississippian and Pennsylvanian (Carboniferous) Systems of the United States. Prof. Pap. U.S. geol. Surv. 1110-A, A1-30. SMITH, W. E. 1979. Pennsylvanian stratigraphy of Alabama. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Alabama and Mississippi. Prof. Pap. U.S. geol. Surv. lllO-I, 123-36. SNOKE, A. W., KISH, S. A. • SECOR, D. I. 1980. Deformed Hercynian granitic rocks from the Piedmont of South Carolina. Am. J. Sci. 280, 1918-34. STERLING, P. J. & STONE, C. G. 1961. Nickel occurrences in soapstone deposits, Saline County, Arkansas. Econ. Geol. 56, 100-10. STRONG, D. F., DICKSON, W. L. & PICKERILL, R. K. 1979. Chemistry and prehenite-pumpellyite facies metamorphism on calc-alkaline Carboniferous volcanic rocks of southeastern New Brunswick. Can. J. Earth Sci. 16, 1701-85. THOMAS, W. A. 1972. Mississippian stratigraphy of Alabama. Alabama geol. Surv. Mong. 12, 121 pp. 1973. Southwestern Appalachian structural system beneath the Gulf Coast Plain. Am. J. Sci. 273-A, 372-90. 1977. Evolution of Appalachian-Ouachita sediments and recesses from reentrants and promontories in the continental margin. Am. J. Sci. 277, 1233-78. 1982. Regional geology and aeromagnetic map of the Southern United States. Abstr. Progr. geol. Soc. Am. 14, 631. & CRAMER, H. R. 1979. The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States--Georgia. In: The Mississippian and Pennsylvanian (Carboniferous) Systems in the United States. Prof Pap. U.S. geol. Sure. 1110-H, H1-37. -

THOMPSON, J. B. (JR) & ROBINSON, P. Geologic setting of the Harvard conglomerate, Harvard Massachusetts. In: CAMERON, B. (ed.) Geology of Southeastern New England, 345-51. Guidebk for Field Trips to the Boston Area and Vicinity, NEIGC, Science Press, Princeton. VAN DE POLL, H. W. 1963. Carboniferous, volcanic and sedimentary rocks of the Lower Shin Creek area, Sinberry Co., New Brunswick. Unpublished thesis, University of New Brunswick, Fredericton, Canada. 1967. Carboniferous volcanic and sedimentary rocks of the Mount Pleasant area, New Brunswick. Can. Dept Invest. 3. Mineral Resources Br., Dept. Nat. Resources of New Brunswick, Fredericton. 1972. Stratigraphy and economic geology of Carboniferous basins in the Maritime Provinces. 24th int. Geol. Congr. Field Excursion A60, 1-90. 1973. Stratigraphy, sedimentation, dispersal and facies analysis of the Pennsylvanian Pictou Group in New Brunswick. Marit. Sed. 9, 72-7. VAN DER VOO, R. & SCOTESE, C. 1981. Paleomagnetic evidence for a large (2000 km) sinistral offset along the Great Glen fault during Carboniferous time. Geology, 9, 583-9. ~, FRENCH, A. N. & FRENCH, R. B. 1979. A paleomagnetic pole position from the folded Upper Devonian Catskill red beds and its tectonic implications. Geology, 7, 345-8. VAN HOUTEN, F. I . 1976. Late Variscan nonmarine deposits, northwestern Africa: implications for pre-drift North Atlantic reconstructions. Am. J. Sci. 276, 671-93. -1981. The odyssey of molasse. In: MIALL, A. D. (ed.) Sedimentation and Tectonics in Alluvial Basins. Pap. geol. Ass. Can. 23, 35-48. VIELE, G. W. 1979. Geologic maps and cross sections, eastern Ouachita Mountains, Arkansas: Map Summary. Bull. geol. Soc. Am. 90, 1096-9. 1982. Interpretation of aeromagnetic anomalies South Central United States. Abstr. Progr. geol. Soc. Am. 14, 638. WEBB, G. W. 1969. Paleozoic wrench faults in Canadian Appalachians. In: KAY, M. (ed.) North Atlantic Geology and Continental Drift. Mere. Am. Ass. Petrol. Geol. 12, 754-91. WHEELER, R. 1978. Slip lines from Devonian Millboro Shale, Appalachian plateau province: review and extensions of discfold analysis. Am. J. Sci. 278, 497-517. WILLIAMS, H. 1980. Comment. Geology, 8, 211-2. WILTSCHKO, D. V. 1981. Thrust sheet deformation at a ramp: summary and extensions of an earlier model, ln: MCCLAY, D. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. Lond. 9, 55-9. Blackwell Scientific Publications, Oxford. ZARTMAN, R. E., SNYDER, G., STERN, T. W., MARVIN, R. F. & BUCHNAM,R. C. 1965. Implications of new radiometric ages of eastern Connecticut and Massachusetts. Prof. Pap. U.S. geol. Sure. 525D, D1-10. -

-

-

-

Alleghenian orogeny in America ZIEGLER, A. M., SCOTESE, C. R., MCKERROW, W. S., JOHNSON, M. E. & BAMBACH,R. K. 1979. Paleozoic paleogeography. Ann. Rev. Earth planet. Sci. 7, 473-502.

217

ZtETZ, I., HAWORTH, R. T., WILLIAMS, H. & DANIELS, D. L. 1980, MagneticAnomaly Map of the Appalachian Orogen. Memorial University of Newfoundland Map No. 2, 1:1,000,000.

NICHOLASRAST, Hudnall Professor and Chairman, Department of Geology, 252 Bowman Hall, University of Kentucky, Lexington, Kentucky 40506, USA.

Geophysical evidence for the extension of the Variscan front on to the Canadian continental margin: geodynamic and palaeogeographic consequences Jean-Pierre Lefort & Richard T. Haworth SUMMARY: Dextral transcurrent movement and thrusting induced by transcurrent movement have been identified along a roughly east-west line across the continental shelf NE of Newfoundland at approximately 52~ The location of that boundary correlates well on a pre-Mesozoic fit of the North Atlantic with the terminus of the Variscan front on the European margin. Seismic reflection, refraction, gravity and magnetic data show a structural succession across the boundary that correlates with that across the front in Europe. This North American continuity of the Variscan front has subsequently acted as the locus for Mesozoic tectonics. Two main hypotheses have been proposed for the location of the Variscan front in North America. Most proponents believe that the front, after crossing the Porcupine Bank, turns south-westwards towards Newfoundland. Some authors such as Hurley (1968) and Cherkis, Fleming & Massingill (1973) have proposed that it crosses Newfoundland, but most think that it veers southwards before reaching the island because the deformation of Variscan age which is known onshore does not have the same style (Belt 1969) as that known in southern England (Sanderson & Dearman 1973). Such an hypothesis is supported by Kay (1969), Bard, Capdevila & Matte (1971), Rast & Grant (1973), Dunning (1980) and Gardiner & Sheridan (1981). This proposed path for the front is attractive because the feature is then concentric with the Ibero-Armorican arc (which represents the Variscan rejuvenation of a Precambrian feature-Lefort 1979) and parallels the sigmoidal shaped Precambrian ridges of the Grand Banks of Newfoundland (Haworth & Lefort 1979). A few authors prefer a direct, westward, extension of the front crossing all the Appalachian trends in the Southern Labrador Sea (Ziegler 1975; Park 1980; Lefort 1980). Such a preference implies that Newfoundland and the Grand Banks were part of the Variscan 'terrane' even if they did not suffer the same deformation as southern England in the Carboniferous. In order to discriminate between these two hypotheses, geophysical and geological data in the Southern Labrador Sea have been reexamined for evidence of correlation with the recently re-interpreted location of the Variscan front on the Porcupine Bank (Lefort & Max 1984).

Mississippian and Pennsylvanian deformation in the Southern Labrador Sea The seismic reflection data studied by Haworth, Grout & Folinsbee (1976a) and Haworth et al. (1976b) have been re-interpreted together with the data of Grant (1972), concentrating particularly on the Upper Palaeozoic deformation and structural boundaries. The area is characterized by northward trending geological zones that are generally younger towards the east (Haworth et al. 1976a). The zones are usually separated by faults, thrusts or unconformities. The ages of the zones are based on drill core samples from the various seismic units recognized. The westernmost zone (west of 55 ~ 10'W, at 52" N) is characterized by a Helikian (Middle Proterozoic) or older basement consisting of gneiss, schist, granite and gabbro. It is covered by late Hadrynian (Upper Proterozoic) and Cambrian quartzites, slates, limestones, dolomites and locally minor basalts. In some places Upper Cambrian to Middle Ordovician limestones, dolomites, shales and slates are known. All of these rocks comprise the western autochthonous unit (Fig. 1). An allochthonous unit centred on Hare Bay includes Hadrynian to Middle Ordovician quartzose greywackes and slates, basalts, tufts, schists and amphibolites which lie structurally beneath peridotite sheets. Lower Ordovician basalts, tufts and black slates, or Middle Ordovician m61ange are locally at the base of, or within the allochthon. These units are truncated in the east by the seaward extension of the Taylors Brook fault (Belt 1969). East of this fault lies what we will call the eastern para-autochthonous unit. 219

J.-P. L e f o r t & R. T. H a w o r t h

220 56*

55"~0'

54*30'

55*

54*

53030 .

53"

t

/ /

/ /

/

./

53*

'\\

//

/

/

\~\\

x\ \

x \o "

x

MP

"~

HE?

/

/

r

a"

~

~,/ /

.. O~ tN ~

~

,

'

~

,

"

'

,

///"

HE

i

j

t

~5~.<~

1 \

b

,,j

-

v .... ~

?

/ re

,r/' 9

'.

', o'

_.'

HE

GROAIS

'h~~x

i

55.30.

~ - ~ DISRUPTION --IF---# FLEXURE

"., o~ "

56*

./

//,

55"

i

5,'3o'

~"

~

;,-

LIMIT OF MESOZOIC COVER

53:30'

13.

FIG. 1. Geology of the southern Labrador Shelf adapted from Haworth et al. (1976a). Letters A - J identify those folds within the Mississippian-Pennsylvanian strata that correlate across the thrust/transcurrent fault at approximately 52~ and the discontinuity at approximately 51~ Locations of seismic sections illustrated in Fig. 2 are indicated by Roman numerals. TBF = Taylors Brook fault; CP = Conche Peninsula; CRP = Cape Rouge Peninsula; H = Helikian and older; He = Hadrynian and Cambrian; HO = Hadrynian to Middle Ordovician; EO = Upper Cambrian to Middle Ordovician; OS = Ordovician or Silurian; D = Devonian; MA = Mississippian; MP = Mississippian and Pennsylvanian; KT = Cretaceous and Tertiary. Heavy line with teeth = thrust; dash-dot ornament -- fold axis in Hadrynian terranes; thin continuous dotted line = possible extension of Hadrynian terranes in the Strait of Belle Isle; dash and square line = anticline; dash and cross line = syncline; the cover is located east of the boundary between cover and basement.

North o f H a r e Bay, a narrow belt o f psammitic and pelitic schists (at 5 2 ~ 55 ~ locally basic and derived f r o m basalts, or chloritic schists derived f r o m basaltic rocks and s o m e t i m e s including meta-gabbros and meta-peridotites, is part o f the eastern

p a r a - a u t o c h t h o n o u s unit. T h e contact o f this unit with the western a u t o c h t h o n o u s unit is a thrust which faces west ( H a w o r t h et al. 1 9 7 6 a ) . South o f Hare B a y , at 5 1 ~ 55~ the same unit is f o u n d isolated within Mississippian terrane.

The Canadian continental margin A north-north-westerly trending ridge following the 55~ meridian at 51~ consists of Ordovician or Silurian felsic and mafic volcanics with minor sedimentary rocks. It is locally intruded by Devonian (?) granite rocks. This ridge is also part of the eastern para-autochthonous unit. South of Hare Bay at the same longitude, the same ridge appears to rise close to the bedrock surface, although the outcrop in this area was interpreted by Haworth et al. (1976a) to be Codroy or Windsor sediments (Upper Mississippian) on the basis of rock drill cores obtained farther south. The interpretation of the subsurface extension of the ridge is based mainly on the lack of seismic reflectors in this area, similar to the situation of the northern portion of the ridge. Furthermore, the post-Appalachian uplift recognized south of Hare Bay (Williams 1975) seems to extend eastward producing an early Palaeozoic inlier on Groais Island (Baird 1966) and shallow mafic terrains of probable Ordovician or Silurian age at 54~ 55~ (Haworth et al. 1976a). Grey, green and red conglomerates, sandstones, siltstones and shales of Mississippian age (Anguille, Crouse Harbour and Cape Rouge Formations) are all located west of 55~ The seismic records over these rocks, close to their contact with the older formations, shows that they are clearly transgressive in the west and possibly faulted in the east, except at 51~ where the eastern boundary shows a clear transgressive contact which we will discuss below. The grey, red or brown sandstones conglomerates and shales (mainly Barachois group) of Mississippian and Pennsylvanian age are always located east of 55~ and its western contact is now recognized to be always transgressive on the Ordovician or Silurian ridges, at least in the area south of 51~ Mississippian and Pennsylvanian folds

Seismic data do not show any thickening of the Carboniferous strata towards the centre of their synforms. Such structures are therefore the result of folding, rather than downwarping at a sedimentary depocentre, and have been controlled by the pre-Carboniferous basement trending northwards in this region (Haworth & Jacobi 1983). An irregular sequence of synclines and anticlines is found from west to east, north of a line running slightly north of east through 52~ 5489176 (Fig. 1, A - G ) . All these structures trend north or NW. South of the roughly E - W

221

trending line the same irregular sequence has been found with the same orientation, except that the flexure A gives way in the south to the Ordovician or Silurian mafic ridge described earlier. The southern set of structures and other short associated features stop at about 51 ~ There is a systematic dextral offset of approximately 13 km between these northern and southern sets of folds. South of 51~ a new succession of folds (Fig. 1, H-J) trending north or NW may be observed; here the Ordovician or Silurian mafic ridge again shows a dextral offset of approximately 13 km. In addition to the interruption of the folds and ridges by the E - W boundary, the general sigmoidal path of the fold axes between 51 ~ and 53~ suggests a general dextral shear component for this area since or during late Carboniferous time. West of the Ordovician-Silurian ridge, the lower Mississippian terrain shows other folds trending northwards, but because the same structural succession cannot be recognized north and south of 51~ the amount of displacement cannot be calculated. Two different sets of magnetic lineations occur in the area of these fold axes. The broad magnetic high at approximately 54~ veers north-eastward at its northern end, crosscutting many Carboniferous fold axes. This high does not appear to deviate where the Mississippian and Pennsylvanian axes are bent or disrupted. To the east of this broad high, there is often a close relationship between the magnetic lineations and the fold axes: magnetic lows are coincident with the Carboniferous synclines, so that the magnetic trends themselves veer in the same direction as the folds exhibiting the same sigmoidal shape. Furthermore, all magnetic trends recognized between 52 ~ and 54 ~ show the same disruption north of 51 ~ or 51 ~ significance of these two different sets of magnetic lineations, caused respectively by deep and shallow structures, will be discussed in a later section. 090 ~ and 080 ~ trending discontinuities in the Mississippian and Pennsylvanian terranes

The existence of an E - W trending discontinuity just south of Hare Bay was discussed by Haworth et al. (1976a). The line was seen as providing a limit to certain Carboniferous trends, as well as being the northern limit of a major gravity high. This line may mark the northern boundary of the area of postAppalachian uplift recognize south of Hare Bay. The line may also be interpreted as a

FIG. 2. Seismic reflection records across the interpreted extension of the Variscan front on the southern Labrador Shelf, locations indicated in Fig. 1. Section I is from Karlsen 75-018 (2300/229 to 0220/230), section II is from Sackville 69-041 (0000/204 to 0115/204) and section III is from F. L. Blair 67 (1210/241 to 1250/241). Section I complete vertical scale: 480 ms (TWTT), section II, complete vertical scale: 855 ms (TWTT), section III, complete vertical scale: 500 ms (TWTT), complete horizontal scale 22 km. b: basement; c: cover; q: Quaternary; sb: seabed; arrow: thrust or sharp change in style of folding.

The Canadian continental margin dextral transcurrent fault zone because of the systematic dextral offset of the Carboniferous ridges east of 55~ There is, however, no clear offset in the magnetic expression of the Hadrynian volcanics trending 010 ~ at the mouth of Hare Bay. It is difficult to say if, or how, this shear zone links with the northern extension of the Taylors Brook fault: too many faults are seen on the seismic records to be sure of the correlation. However, the detailed geology of the Conche Peninsula and Cap Rouge Peninsula south of Hare Bay (Baird 1966) shows that post- or syn-Carboniferous dextral transcurrent faults oriented 040 ~ are known in the area, and that the Taylors Brook fault tends to veer northeastward in the Biche Arm, north of Cap Rouge harbour. This fault had dextral movement in Carboniferous time (Webb 1969). The offshore E - W discontinuity is probably not only a transcurrent zone because seismic reflection profiles indicate that the Carboniferous section also increases in thickness from south to north across this line. This is probably also the reason why the lower Mississippian basin is not in fault contact at its eastern boundary with the Ordovician-Silurian ridge at 51 ~10'N. The northernmost transverse discontinuity is an arch trending obliquely between 52~ 53~ and the Belleisle Strait. In the west it has been identified on the basis of its offset of the magnetic pattern associated with Hadrynian volcanics. North of 51~ the discontinuity correlates with a major fault on the seismic records. At 51~ the fault is no longer seen, but may be inferred at greater depth by the presence of a large asymmetric fold with its steeper flank facing north. At 52~ (Fig. 2, section III) the discontinuity is shown by a clear thrust which brings steeply dipping terrane from the south on to flatter strata to the north. At 52~ (Fig. 2, section II) the proposed thrust is crossed by a seismic reflection line in a location at which the reflection character changes abruptly. At 52~ (Fig. 2, section I) there is an abrupt change in the wavelength of the folds, the southernmost of them being the tightest. The line of discontinuity therefore shows a progressive evolution from what appears to be transcurrent faulting in the west to thrusting in the east. The latter, we feel, is induced by transcurrent movement. The thrusting itself does not have the same expression or deformational style everywhere. Where thrusting is very well expressed (at 52~ Fig. 2, section III), the ramp of the thrust is at a shallow angle. Usually the inferred thrust line separates folds which have a shorter wavelength in the

223

south than in the north and many of the asymmetric folds south of the line verge north suggesting northerly directed shear. Because of its arcuate trend we suspect that this line of discontinuity may link with the thrust, recognized east of Belleisle Strait, which may be the northern extension of the Taylors Brook fault. In summary, the southern discontinuity is apparently a dextral transcurrent fault with no evidence for thrusting, although downwarping to the north is quite probable. Because there is no evidence for the offset of the volcanic zone across the mouth of Hare Bay, its connection with the Taylors Brook fault, which had dextral movement at the same time, is quite probable. The northern discontinuity clearly shows transcurrent movement with evidence for thrusting in the east, probably linked to the thrust which is known to the east of the Strait of Belleisle. The folds and transverse discontinuities discussed in this section have also provided control over the subsequent deposition of Cretaceous strata (Fig. 2, sections I and II). The western limit of the cover shows a clear recess at 52~ when crossing the only Carboniferous anticline known in that region. At 52~ a deep recess follows the thrust front, and at 51~ there is a perturbation in the general N - S trending edge of the Cretaceous strata. This suggests either that Carboniferous features were rejuvenated during Mesozoic uplift associated with the opening of the North Atlantic, or implies that the topographic highs related to these basement structures were not completely eroded by the Middle Mesozoic. It is tempting to correlate these structures in the southern Labrador Sea with Variscan features in southern Ireland and England because the Carboniferous folding and the shear and thrust zones occur close to the projection of the Variscan front into the area.

The Variscan front in southern England and Ireland Although the Variscan front is often represented as a well delineated and continuous feature in Britain and Ireland (Dunning, 1966), this impression is currently disputed because it does not agree with many field observations. It is usually defined on the basis of major thrusting, but the difficulties with this approach have been assessed by Matthews (1974, this volume), Dunning (1977) and Hancock, Dunne & Tringham (1981). These authors note that in several places there is a choice of predominant thrust boundaries with the probability that thrusting was merely a function of the local

224

J.-P. Lefort & R. T. Haworth

shallowing of pre-Variscan basement or the relative incompetence of strata within the foldbelt, and that the thrusts do not necessarily mark the northern boundary of Variscan orogenic deformation. Indeed, thrusting is not developed everywhere along the conventionally defined front. The possibility of defining the northern limit of development of slaty cleavage as the location of the front was evaluated by Matthews (1974) and discarded because of distributional inconsistencies in SW Britain and Ireland (Max 1979). It is clear that many of the difficulties of definition depend on basement/cover relationships. If the exposed north-western part of the Variscan foldbelt includes localized, fault-controlled basins influenced by pre-existing basement structures, as suggested by Gardiner (1978), the existence of a structural basin might imply that the Variscan front forms its boundary (Gardiner & Sheridan 1981). The magnetic map of southern Britain (IGS 1965) clearly shows a pair of E - W trending magnetic highs and lows south of Wales that can be correlated with the Variscan front, but the gravity data (Blundell 1975) show only a weak correlation for the same area. In contrast, in the Celtic sea, the gravity data indicate a marked disruption in the anticipated location of the front between Ireland and Wales, probably because of Mesozoic and Cenozoic reactivation of basement features. In Ireland, the front has its gravity expression as a line separating a major E - W low in the south from different trending anomalies in the north. Magnetic data released recently (Max, Inamdar & McIntyre 1982) suggest the same path for the Front (Max & Lefort,

this volume). Westward, on the continental shelf, the location of the front on the Irish shelf and the Porcupine Bank has been hypothesized (Riddihough & Max 1976; Lefort 1980), but all such hypotheses discounted the evidence for oceanic crust in the Porcupine Seabight. The recent delineation of 'oceanic' magnetic stripes in the Porcupine Seabight now permits the pre-Mesozoic location of the Bank to be determined, providing a better framework for the delineation of the front offshore (Lefort & Max 1984). In general, the gravity and magnetic expressions of the front are so weak as not to contribute significantly to the definition of its location. The general clockwise bending of Carboniferous structures (although partly controlled by Caledonian basement onshore) north of the conventionally defined location of the front in Wales and Ireland is associated with folds which

veer south-westward south of the front, so that it has been suggested that the front had a dextral transcurrent component (Max & Lefort, this volume). Such an interpretation is supported by the existence of a 30 km dextral offset of the 060 ~ trending magnetic and gravity highs and lows recognized in southern Ireland and the Irish Sea. Other E - W transcurrent shears of about the same age are also known in the Bristol Channel (Freshney & Taylor 1980), in Britain (Sanderson, this volume) and even in NW England (Arthurton, this volume). It is therefore quite logical to find similar features in the Mesozoic and Cenozoic basins of the Celtic Sea, controlled by basement trends and offset by the Variscan front (Dunning 1980). The existence of dextral offset along the front can therefore be used as a criterion for its definition even where it does not show any gravity or magnetic expression.

The Variscan front and associated basins across the N o r t h A t l a n t i c The SW England area (Figs 3 and 4, line AB) Between early Devonian and NamurianWestphalian time, a large E - W oriented basin occupied the Cornubian area. This basin was bounded to the north by the western extension of the Midland craton and to the south by the Armorican uplands. The boundaries in the north and south suffered some displacement during the Upper Palaeozoic but, although the marine limit consequently changed, the axis of the basin has always been a depocentre. The Cornubian basin is strictly limited by deep-seated E - W faults that were probably active during the deposition of the sediments. It is probable that the transition from lowintensity folding to overfolding and overthrusting depended on the position of the southern edge of the northern pre-Variscan basement, which seems to be fairly shallow beneath south Wales. This basement plunges immediately south of Wales and the hinge appears to mark the linear development of the Waulsortian reefs and the southward increase in Carboniferous thickness. This increase in thickness continued under the Bristol Channel until truncated by the postulated major fault zone in the Bristol Channel. Section AB (Figs 3 and 4) summarizes the general structure between Wales and the English Channel (after Freshney & Taylor 1980). From south to north this section shows: a

The Canadian continental margin

/-"~

t ~

/

~ ~,

225

.--.-I 0

~. 0 "~ ~ ; ~

~-~. ,-~--S&~~ ~-~-'7__.~ ~

4//

\!i /

-~ ~~i(~ ~ f

,f / ~ : '

--,.. /<, "-.-'d;~'&", ~ "...

o'~' 3 =

,,, ,r"

_1

%

\

\ ,~ ~I I~ ,"~ 0

I

I-.IA o~

#

24,/

ooD z

o

~

~-~

=I

9

NeN

~ g

./~i

/ ~

e

s~

.~o >._.. g

-~.~ ~~

,,,!..~ ~

6 ~ x ~ o~ "= .~

-~

226

J.-P. Lefort & R. T. Haworth CORNWALL H,F

AO

IRISH

A,

IO0

B

200

SEA

}

IO0

200

I

!22 -- _ 7 - - - ~

---7 "? q~

POST

VARISCAN STRATA

514 +

+

PALEOZOIC STRATA

+

+

+ + BASEMENT

+

i

+

+

+:

+

+

4 0 0 B~

500

%

.... 7 - - _

585/?

_

+

+

+

+

+

+

_

+

+ +

+

+ +

5~9

+

65 NU

NB

PORCUPINE

BANK

A2 o ......

~OO

o

+ 2I + 'NB? SOUTH A3

EASTERN

0 0

?

200

_3.oo

B2

+

+

+

+

NU

+

:

+

+

TR

+

~"---

? SU +

LABRADOR iO0

200

490

300 +

---

TR

+

+

+

NB

I

EASTERN ~

-

O? "2-~ r~495--466k

2 4[

Boo

800

B~

+

NU SOUTH A4

s?o

+

587-~--

+ N8

TR

LABRADOR ;

I00

~

565---

+ NU

SU

o ~

200

_

180----

~

i ? ....... - :-+ -- -- ~ . . . . .

,+ '

+

+

48o---

+ TR

\

300 Z09

5 0 0 B4 ]84

,ZOO

468--~-

6 0~624~'+ + +

-- ~ qe ~..

+

+

+ SU

~ ~

"+~ ~

~ "~57

--

-272 550

492

FIG. 4. Structural sections aligned along the Variscan front in Europe and its interpreted continuation on the Canadian margin. Section AB is inferred from surface geology, section A1B 1 from seismic refraction, section A2B 2 from refraction and reflection, section A3B 3 from reflection and section A4B4 from refraction data. References to the data sets used in constructing the sections are provided in the third section of the text. For all sections the common elements, northern basin (NB), northern basement uplift (NU), Variscan front (HF), thrust ramp (TR) and southern basement uplift (SU) are indicated and discussed in the text. For section AB the representative seismic refraction velocities are 1.80 km s -1 for post-Variscan, 4.66 km s -1 for Palaeozoic and 5.65 km s -1 for basement strata. Vertical scale in kilometres. post-Variscan basin, a Precambrian b a s e m e n t high, the main basin primarily filled by D e v o n ian and Carboniferous strata (this basin being thrust to the north and to the south over the ba sem en t highs), the so-called Variscan front, a basement high and a small Silesian basin more or less superimposed on this basement high. A large number of short seismic refraction lines which have been run in the southern Irish Sea have been summarized by Hall (1978). A l t h o u g h these lines do not show much detail in the structure of the upper crust, they suggest that the b a s em en t to the central Palaeozoic basin is deeper towards Lands End than b e t w e e n Wales and Ireland where a z o n e with a seismic velocity of 6.25 km s -1 is found close to the surface near the Variscan front (Fig. 4). Northwards, strata with that velocity d e e p e n again and the presence of overlying strata with a vel-

ocity of 5.4 km s -] suggests that the northern structure is similar to the structure of the Silesian basin as seen in the geological section of south-western England (Figs 3 and 4, line A1B1). Southern Ireland area (Fig. 3) South of this area, on the basis of the data from onshore and offshore wells, Gardiner & Sheridan (1981) have argued that the Munster basin was not con n ected with the progressively deeper marine environments of the Variscan troughs in the south, as supported by Naylor & Sevastopulo ( 1 9 8 0 ) and Matthews ( 1 9 7 4 ) for D e v o n i a n and early Carboniferous. However, although it is probably true that the 060 ~ trending Celtic S e a - W a l e s platform acted as a high b e t w e e n the Munster and south-western Eng-

The C a n a d i a n c o n t i n e n t a l margin

land basins during Devonian and early Carboniferous time, it is not sufficient proof that those basins were not connected during the late Carboniferous. Also, it is known that after early Carboniferous time deep-water muddy sequences accumulated in the south Irish basin and on the Celtic basin itself (Higgs, in Gardiner & Sheridan 1981). In short, we can assume that early Carboniferous sedimentation was probably controlled by old Caledonian features, but a new structural regime developed after that time independent of the previous trends. That is why we still believe in the existence of a large E - W oriented basin, as hypothesized by Naylor & Sevastopulo (1980) at least for late and Upper Carboniferous time. The classical Variscan front runs approximately parallel to the northern side of the Munster basin which is quite a steep feature in western Ireland (Naylor & Sevastopulo i980). That is probably why the front is better expressed in that region. The western extension of the southern limit of the British Cornubian basin is not known south of the Irish Sea. The only linear feature that might indicate an extension of this boundary is a narrow belt of short-wavelength magnetic anomalies related to a probably Jurassic source (Caston et al. 1981; Lefort & Max 1984). Such a linear feature, 200 km long, could well be related to the reactivation of a major basement boundary such as a basin edge.

Porcupine area (Figs 3 and 4, line A2B2) Although no basement core samples have been taken on the Porcupine Bank, the basement features can be deduced from geophysical data. The Porcupine Seabight has to be closed before any correlation can be achieved because the Bank is interpreted to have drifted northwestward following 'oceanic' transform faults trending 130 ~ (Lefort & Max 1984). Two main E - W features are known on Porcupine Bank. The southernmost feature represents the extension of the narrow belt of short-wavelength magnetic anomalies south of Ireland. South of this, no seismic reflector can be seen beneath the Mesozoic cover (Bailey 1975) suggesting the existence of a metamorphic or plutonic basement in the area. The northernmost feature is an E - W lineated magnetic feature at 52~ (Riddihough & Max 1976), although new magnetic data (Max et al. 1982) suggest that the feature might be more correctly located at 52~ where Lefort & Max (1984)

227

tentatively locate the front. This linear feature correlates with the extension of the front west of Ireland. The line is also more or less coincident with the 53 ~ flexure described by Bailey (1975). We suggest that the shallow seismic lines of Bailey show the Variscan front (a shear or a thrust) while the gravity or magnetic data a little farther south would be related to the northern boundary of the basin seen in the east. The existence of seismic reflectors at depth only between the two E - W lines would favour this interpretation. On the other hand, N - S oriented deep seismic refraction line run on the Bank (Whitmarsh et al. 1974) has shown that the 6 km s-1 velocity basement dips southward and rises to the surface near 53~ supporting the existence of a Palaeozoic basin to the south. However, the reflectors which could be related to the existence of a northern basin (see the Cornubian Section) are located far in the north (54~ 12~ and there is a large area between 53~ and 53~ that is devoid of seismic reflectors, suggesting that the 'northern uplift' (see Fig. 4) has been eroded during the Mesozoic uplift of the bank before the opening of the North Atlantic.

Orphan Knoll area (Fig. 3) The late Palaeozoic position of the deep structures of the northern part of the Grand Banks of Newfoundland and the southern Labrador Sea has been compared with those recognized on Porcupine Bank, southern Ireland and SW Britain (Fig. 3). In the Orphan Knoll area, few geological and geophysical data are available. However, the free air gravity anomaly map of offshore eastern Canada (Shih, in Keen & Hyndman 1979) shows that the northern boundary of Orphan Knoll is bounded by a deep low which is probably caused by a thick Mesozoic basin. The boundary corresponds with a steep fault, clearly oriented E - W which is an extension of the southern linear feature recognized on Porcupine Bank (see above). Because of its position and orientation, this basin could well be related to a basement feature with a similar orientation. Furthermore, a DSDP well (site 111) drilled on the northern flank of the Knoll shows that the Bajocian sediments contain coal of almost certain Palaeozoic age (Ruffman & van Hinte 1973). This coal, which shows a germanium content similar to that of South Wales, suggested initially that the coal originated in that area, but our correlations

228

J.-P. L e f o r t & R. T. H a w o r t h

support the possibility that it could well be of Canadian origin. If this were the case then because the coal is of high anthracite rank, Variscan metamorphism should exist beneath the eastern part of the Grand Banks of Newfoundland. This is reinforced by the existence of small biotites in a Carboniferous core drilled west of this zone. The possibility of a deep basin in this area has also been suggested by Jansa & Mamet (1984).

The South Labrador Shelf area (Figs 3 and 4, lines A3B 3 and A4B 4 The outer portion of the shelf area has been surveyed by the oil companies and data between 48~ and 54~ have been released (Curt & Laving 1977). This published section, based on seismic reflection records, shows the following from south to north beneath the Mesozoic and Cenozoic cover: a basin in the south, a narrow basement uplift, a large basin which Cutt & Laving think to be filled with more than 4570 m of Visean and Namurian strata, a large northern uplift and a northern basin. This structural succession is similar to that found on section AB in SW England. The only difference is in the width of the northern uplift, which is identical to that of the northern uplift of Porcupine Bank. This enlargement of the structure can be explained by erosion during the doming of the continental crust prior to the actual opening of the North Atlantic. Moreover, Curt & Laving (1977) show a planar contact between the middle basin and the northern uplift close to 52~ which suggests the existence of a tectonic contact. This contact is located a few kilometres east of the location where the Mississippian and Pennsylvanian series of northern Newfoundland show a marked change in deformational style (see section I). It must also be emphasized that the diapirs found along the axis of this basin are a link between the diapirs found north of Newfoundland (Haworth et al. 1876a) and those found on the Porcupine Bank (Bailey 1975). The E - W trending synform traced continuously to the Canadian margin from SW England, is locally affected by a graben which parallels the continental margin of Labrador (the Belleisle sub-basin of Cutt & Laving 1977). This is probably a younger feature, now covered directly by early Eocene sediments, related to the opening of the North Atlantic. As mentioned earlier, the 51~ discontinuity seems to be the place where shallow seismic reflection data indicate a thickening of

the Mississippian and Pennsylvanian strata. This is also indicated by the seismic refraction data of Sheridan & Drake (1968) between lines 152 and 148 (see section A4B 4 in Fig. 4) where a thick basin of strata with a seismic velocity of 4.8 k m s -1 is seen beneath the Cretaceous cover, thus extending the European and Orphan Knoll structures. The seismic refraction line shows a structural succession similar to that in Cornwall with a southern uplift that predates Cretaceous time (it is not associated with the Atlantic opening), a large basin in the middle, a northern uplift and a northern basin. Furthermore, the thrust delineated by our seismic reflection records appears to be located close to the southern edge of the northern uplift and related to it in exactl~ the same position as in England. The western termination of the trans-Atlantic Carboniferous basin has been delineated by Wade et al. (1975) since their St Anthony basin includes Carboniferous strata (Grant 1974). The southern part of this basin displays an E - W trending boundary that trends towards the northern edge of Orphan Knoll. At depth, the refraction data show that the southern flank of the pre-Mesozoic synform is steep while the northern edge is shallow and dipping gently south. _ _

Conclusions on the northern Variscan c o r r e l a t i o n s in t h e N o r t h A t l a n t i c a r e a There is no suggestion, other than the speculation of Jansa & Mamet (1984) that the same E - W trending structural basin of Carboniferous age extends from Cornwall to the southern Labrador Sea. Indeed, we do not know whether the pre-Mesozoic basin of southern Labrador is mainly filled with Carboniferous strata, or whether it is a half-graben (the southern margin of the synform being far steeper than the northern one) filled by a different Palaeozoic formation covered by Carboniferous sediments. However, this does not affect our general interpretation for the Variscan front because the thrust recognized north of Newfoundland developed on the northern boundary of a Palaeozoic or Carboniferous synform, as in England; and because this thrust is related, as in Ireland, to dextral wrenching following or generating the Carboniferous folding. Furthermore, the occurrence of small biotites and coal metamorphosed during the Variscan orogeny suggests that we have, in Canada, the equivalent of European Variscan activity. The thrust-

The Canadian continental margin ing seems to have been controlled by the change in structure from a synform in the south to the shallow basement in the north. If our interpretation is correct, and if the same elongated basin is found, it suggests that the R h e n o - H e r c y n i a n syncline (or northern Variscan trough) does not follow the I b e r o - A r m orican curvature (Gardiner & Sheridan 1981), but extends westward in the direction of the Grand Banks of Newfoundland. The scarcity of Variscan events in Newfoundland does not present any major problems if the thrust does really veer to the south and link with the Taylors Brook fault; the dextral shear of the fault would be related to the oblique thrust. Such a reactivation of major 0500-060 ~ trending Appalachian features at their junction with prominent trans-Atlantic E - W shear belts has previously b e e n observed in Nova Scotia (Lefort & Haworth 1978) and in New England (Lefort & Haworth 1981) (Fig. 4).

229

The discrepancy between the different behaviour of the magnetic ridges known in the southern Labrador Sea can be explained if the underformed magnetic ridge is either a Mesozoic intrusion related to the opening of the north Atlantic or was not involved in the thrusting as part of the Precambrian or Taconic basement, while the deformed ridges were involved in the thrusting and deformation at the same time as the cover (thin skin tectonics). In conclusion, we support the idea that equivalents of the Variscan front, as known in England and western Ireland, exist in Canada. This boundary is not the northern limit of Variscan deformation; the existence of folds in the Mississippian and Pennsylvanian trending N - S , north of the thrust and on the southern Labrador Shelf, reinforces the European conclusion that thrusting is only due to a shallow basement bounding an E - W synform that is in itself possibly a Carboniferous basin.

References BAILEY, R. J. 1975. Sub-Cenozoic geology of the British continental margin (lat 50~ to 57~ and the re-assembly of the North Atlantic late Paleozoic supercontinent. Geology, 3, 591-4. BAIRD, D. M. 1966. Carboniferous rocks of the Conche-Groais Island area, Newfoundland. Can. J. Earth Sci. 3, 247-57. BARD, J. P., CAPDEVILA,R. & MATTE,Ph. 1971. La structure de la chaine Hercynienne de la Meseta iberique: comparison avec les segments voisins. In: DEBYSER,J., LE PICHON, X. & MONTADERT, L. (eds) Histoire Structurale du Golfe de Gascogne. Publ. Inst Francais Petrole, 22, (1), 14.1-67. BELX, E. S. 1969. Newfoundland Carboniferous stratigraphy and its relation to the Maritimes and Ireland. In: KAY, M. (ed.) North Atlantic--Geology and Continental Drift. Mere. Am. Ass. Petrol. Geol. 12, 734-53. BLUNDELL, D. J. 1975. The geology of the Celtic Sea and southwestern approaches. In: YORATH,C. J., PARKER, E. R. & GLASS, D. J. (eds) Canada's Continental Margins and Offshore Petroleum Exploration. Mem. Can. Soc. Petrol. Geol. 4, 341-62. CASTON,V. N. D., DEARNLEY,R., HARRISON,R. K., RUNDLE, C. C. & STYLES, M. T. 1981. Olivine dolerite intrusions in the FastnetBasin. J. geol. Soc. London, 138, 31-46. CHERKIS, N. Z., FLEMING,H. S. & MASSINGILL,J. V. 1973. Is the Gibbs Fracture Zone a westward

projection of the Hercynian Front into North America? Nature, 245, 113-5. Cuqq, B. J. & LAVING,J. G. 1977-Tectonic elements and geological history of the South Labrador and Newfoundland continental shelf, eastern Canada. Bull. Can. Petrol. Geol. 25, 1037-58. DUNNING,F. W. 1966. Tectonic Map o f Great Britain and Northern Ireland. Institute of Geological Sciences, London. -1977. Caledonian-Variscan relations in northwest Europe. In: La Chaine Varisque d'Europe Moyenne et Occidentale. CoUoques int. Cent. nat. Rech. Scient. 243, 165-80. 1980. The geotectonic position of the British Isles in northwest Europe. In: Geology o f European Countries--Austria, Federal Republic o f Germany, Ireland, The Netherlands, Switzerland, United Kingdom, 331-4. Dunod, Paris. FRESHNEY,E. & TAYLOR,R. 1980. The Variscides of southwest Britain. In: Geology o f the European Countries--Austria, Federal Republic o f GermanY, Ireland, The Netherlands, Switzerland and United Kingdom, 379-87, Dunod, Paris. GARDINER, P. R. A. 1978. Is the Hercynian Front in Ireland a local feature? Nature, 271, 538-9. SHERIDAN, D. J. R. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan Front. J. struct. Geol. 3, 317-31. GRANT, A. C. 1972. The continental margin off Labrador and eastern Newfoundland--Morphology and Geology. Can. J. Earth Sci. 9, 1394-430.

-

-

230

J.-P. L e f o r t & R. T. H a w o r t h

1974. Structural trends of the western margin of the Labrador Sea. In: VAN DER LINDEN, W. J. M. & WADE, J. A. (eds) Offshore Geology of Eastern Canada. Pap. geol. Surv. Can. 74-30, 2, 217-31. HALL, J. 1978. Crustal structure of the eastern North Atlantic seaboard, In: BOWLS, D. R. & LEAKE, B. E. (eds) Crustal Evolution in Northwestern Britain and Adjacent Regions. Spec. Iss. geol. Soc. Lond. 10, 23-38. HANCOCK, P. L., DUNNE, W. M. & TRINGHAM, M. E. 1981. Variscan structures in southwest Wales. Geologie Mijnb. 6, 81-8. HAWORTH, R. T., GRANT, A. C. & FOLINSBEE, R. A. 1976a. Geology of the continental shelf off southeastern Labrador. Pap. geol. Surv. Can. 76-1C, 61-70. & JACOBI, R. D. 1983. Geophysical correlations between the geological zonation of Newfoundland and the British Isles. In: HATCHER, R. D. (Jr), ZIE'IZ, I. & WILLIAMS, H. (eds) Tectonics and Geophysics of Mountain Chains. Spec. Pap. geol. Soc. Am. 158, 25-32. & LEFORT, J.-P. 1979. Geophysical evidence for the extent of the Avalon zone in Atlantic Canada. Can. J. Earth Sci. 16, 552-67. ~, POOLE, W. H., GRANT, A. C. & SANFORD, B. V. 1976b. Marine Geoscience Survey northeast of Newfoundland. Pap. geol. Surv. Can. 76-1A, 7-15. HURLEY, P. M. 1968. The confirmation of continental drift. Scient. Am. 218, 51-69. Institute of Geological Sciences 1965. Aerornagnetic map of Great Britain Sheet 2. JANSA, L. F. & MAMET, B. 1984. Offshore Visean of Eastern Canada: paleogeographic and plate tectonic implications. In: Proc. 9th int. Carboniferous Congress, Urbana, Illinois, 1979. In press. KAY, M. 1969. Continental drift in the North Atlantic Ocean. In: KAY, M. (ed.) North Atlantic Geology and Continental Drift. Mere. Am. Ass. Petrol. Geol. 12, 965-74. KEEN, C. E. & HYNDMAN, R. D. 1979. Geophysical review of the continental margins of eastern and western Canada. Can. J. Earth Sci. 16, 712-47. LEFORT, J.-P. 1979. The Ibero-Armorican arc and the Hercynian orogency in western Europe. Geology, 7, 384-8. 1980. Un "fit" structural de l'Atlantique Nord: arguments geologiques pour correler les marqueurs geophysiques reconnues sur les deux marges. Mar. Geol. 37, 355-69. & HAWORTH, R. T. 1978. Geophysical study of basement features on the western European and eastern Canadian shelves: trans-Atlantic correlations and Late Hercynian movements. Can. J. Earth Sci. 15, 392-404. &~ 1981. Geophysical correlation between basement features in North Africa and Eastern New England: their control over North Atlantic structural evolution. Bull. Soc. gdol. mindr. Bretagne, 13, 103-16. -

-

& MAX, M. D. 1984. Development of the Porcupine Seabight: the direct relationship between early oceanic and continental structures. J. geol. Soc. London, in press. MATTHEWS, S. C. 1974. Exmoor thrust? Variscan front? Proc. Ussher Soc. 3, 82-94. MAX, M. D. 1979. Geotectonic map of Ireland. In: Atlas of Ireland, Sheet 5, Southwest. Dublin Institute of Advanced Studies. , INAMDAR,D. D. & MCINTYRE, T. 1982. Compilation magnetic map: The Irish continental shelf and adjacent areas. Rep. geol. Surv. Irel. 82/2. NAYLOR, D. & SEVASTOPULO, G. D. 1980. Ireland, stratigraphy and palaeogeography. In: Geology of European Countries: Austria, Federal Republic of Germany, Ireland, The Netherlands, Switzerland, United Kingdom, 338-43. Dunod, Paris. PARK, R. G. 1980. The Lewisian of NW Britain. In: Geology of European Countries, Austria, Federal Republic of Germany, Ireland, the Netherlands, Switzerland, United Kingdom, 338-43. Dunod, Paris. RAST, N. & GRANT, R. 1973. Transatlantic correlation of the Variscan-Appalachian orogeny. Am. J. Sci. 273, 572-9. RIDDIHOUGH, R. O. & MAX, M. D. 1976. A geological framework for the continental margin to the west of Ireland. Geol. J. 11, 109-20. RUFFMAN, A. & VAN HINTE, J. E. 1973. Orphan Knoll--a "chip" off the North American "plate". In: HOOD, P. J. (ed.) Earth Science Symposium on Offshore Eastern Canada. Pap. geol. Surv. Can. 71-23,407-49. SANDERSON, D. J. & DEARMAN, W. R. 1973. Structural zones of the Variscan foldbelt in SW England, their location and development. J. geol. Soc. London, 129, 527-36. SHERIDAN, R. E. & DRAKE, C. L. 1968. Seaward extension of the Canadian Appalachians. Can. J. Earth Sci. 5, 337-73. WADE, J. A., GRANT, A. C., SANFORD, B. V. & BARSS, M. S. 1975. Basement Structures: Eastern Canada and adjacent areas. Map Geol. Surv. Can. 1400A. WEBB, G. W. 1969. Palaeozoic wrench faults in Canadian Appalachians. In: KAY, M. (ed.) North Atlantic Geology and Continental Drift. Mern. Am. Ass. Petrol. Geol. 12, 754-86. WHITMARSH, R. B., LANGFORD, J. J., BUCKLEY, J. S., BAILEY, R. J. & BLUNDELL, D. J. 1974. The crustal structure beneath Porcupine Ridge as determined by explosion seismology. Earth planet. Sci. Lett. 12, 197-204. WILLIAMS, H. 1975. Structural succession, nomenclature and interpretation of transported rocks in western Newfoundland. Can. J. Earth Sci. 12, 1874-94. ZIEGLER, W. H. 1975. Outline of the geological history of the North Sea. In: WOODLAND, A. E. (ed.) Petroleum and the Continental Shelf of Northwest Europe, Volume 1: Geology, 165-90. Applied Science, London. -

-

The Canadian continental margin JEAN-PIERRE LEFORT, Laboratoire de G6ologie dynamique, CNRS--Centre Armoricain d'Etude Structurale des Socles, Institut de G6ologie, Universit6 de Rennes, Campus de Beaulieu, 35042 Rennes Cedex, France. RICHARDT. HAWORTH,Atlantic Geoscience Centre, Geological Survey of Canada, Bedford Institute of Oceanography, PO Box 1006, Dartmouth, Nova Scotia B2Y 4A2, Canada. Now at: Institute of Geological Sciences, Nicker Hill, Keyworth, Nottingham NG12 5GG.

231

The deformation and metamorphism of Carboniferous rocks in Maritime Canada and New England Sharon Mosher & Nicholas Rast SUMMARY: In the northern Appalachians, rocks showing polyphase Variscan/Alleghenian deformation are found in Nova Scotia, New Brunswick, Rhode Island and Massachusetts. In Canada the deformed succession ranges from Tournaisian to Permian, and in New England from Westphalian to Stephanian. The same polyphase deformation sequence (F1-F4) involving the same sequential changes in the geometry and vergence of structures can be recognized in New Brunswick and in Rhode Island. This implies that the Carboniferous rocks in both areas were affected by a similar set of stresses in late Carboniferous-Permian times. The folded succession in both areas is affected by the essentially post-deformational regional metamorphism which reaches garnet grade in New Brunswick and sillimanite grade in Rhode Island. Granites, probably of anatectic origin, cut through the metamorphic rocks in both areas, but are more prevalent in Rhode Island. In both areas overthrusting can be related to fold episodes, and late transcurrent and normal faults of a brittle-ductile nature cut earlier structures. Although there are several current plate tectonic interpretations of Variscan events in the northern Appalachians, all are speculative and further data are needed.

Late Palaeozoic, post-Acadian structures are developed locally in the Carboniferous rocks of New England (Mosher & Wood 1976; Skehan & Murray 1979; Mosher 1983) and eastern Canada (Gussow 1953; Rast & Grant 1973a, 1977) and are related to a Variscan-Alleghenian orogenic event (Rast & Grant 1973b; Skehan & Murray 1980 i Mosher 1981, 1983; Rast 1983). The deformation and associated regional metamorphism are best seen in southeastern New Brunswick (Rast et al. 1978) and Rhode Island (Quinn 1971; Mosher 1981, 1983) where multiple phases of folding and faulting have affected the rocks. Although local variations exist, the overall sequence, style, and orientation of structures is similar in nearly all respects, indicating that the deformational history in both areas is the result of analogous, and presumably related, stress systems. In contrast, many Carboniferous rocks in adjacent areas are undeformed or only mildly deformed (Fig. 1B). This paper compares VariscanAlleghenian structures and associated metamorphism and igneous activity in the two intensely deformed areas and contrasts these rocks to the less intensely deformed and undeformed Carboniferous rocks elsewhere in the Canadian Maritimes and New England. Although both authors have proposed possible plate tectonic models elsewhere (Mosher 1981, 1983; Rast 1983), none will be postulated in this paper. Instead, the constraints on any models will be discussed and the need for more data stressed.

Geological setting Carboniferous rocks in the northern Appalachians are restricted to local basins in the Canadian Maritimes (Belt 1968) and Rhode Island-Massachusetts (Quinn & Oliver 1962) (Fig. 1). The basins are interpreted as a series of pull-apart grabens related to major transcurrent faults (Belt 1968; Webb 1963, 1969) (Fig. 1). Sediments throughout the area are mainly non-marine, although marine Vise a n - N a m u r i a n carbonates and evaporites are well known from Nova Scotia (Schenk 1970) and Newfoundland (Baird 1959) and even parts of New Brunswick. During early rifting ( R u s t 1981; McMaster et al. 1980), alluvial fan and braided stream sedimentation dominated; slower subsidence followed with fluvial or lacustrine deposition (Belt 1968; Bradley 1982). In the Canadian basins sedimentation started in Late Devonian-Mississippian time and was concurrent with, and presumably related to, movements on NE-trending transcurrent faults, most of which appear to have been right lateral faults (Webb 1969). Consequently there are considerable variations in the D e v o n o - C a r boniferous deposits of various parts of the Maritimes. By Pennsylvanian time, most active strike-slip faulting had ceased. A n exception to this was the NE-trending Harvey-Hopewell fault in south-eastern New Brunswick which was left lateral during the Pennsylvanian but 233

234

S. M o s h e r & N. R a s t

FIG. 1. (A) Distribution of Carboniferous deposits and major faults (heavy lines) in New England and Canadian Maritimes (modified after Rast & Grant 1977). (St John's Basin is within the southern New Brunswick thrust complex; NB--Norfolk Basin.) (B) Location of major faults and intensity of deformation (after Keppie et al. 1982). Large solid arrows show primary direction of tectonic transport (LCF--Lake Char and Honey Hill fault system; BBF--Bloody Bluff fault; NF--Norumbega fault; LBF--Lubec-Belleisle faults; HHF--Harvey-Hopewell fault; HF--Hollow fault; LRF--Long Range fault; CCF--Cobequid-Chedabuto fault). was right lateral before this (Webb 1969) (Fig. 1B). Towards the end of the Mississippian there was a general cessation of horst-graben generating movements and slow subsidence of the basins had begun (Bradley 1982). Pennsylvanian rocks of the Maritimes are much more uniform continental and fluviatile sediments

and were deposited in broad flood basins under semi-arid conditions (Legun & Rust 1982). In contrast, the basins in south-eastern New England started later during the Pennsylvanian and we think they were related to left lateral motion on NE-trending transcurrent faults (McMaster et al. 1980; Mosher 1983). Most of these sedi-

Deformation of Carboniferous rocks in New England ments are fluviatile and were deposited in broad humid alluvial fans (Severson & Boothroyd 1981). The two basins which show the most intense deformation are the Saint John basin in New Brunswick and the southern portion of the composite Narragansett basin of Rhode Island (Figs 1 and 2). Some polyphase deformation has also been recognized in Nova Scotia but has only been partially studied (Fyson 1967; Currie 1977) (Fig. 1B). The orogenic effects in Nova Scotia are referred to by Keppie (1981) as Variscan. All areas affected by strong polyphase deformation are floored by Avalonian basement and are found adjacent to the E - W trending Cobequid-Chedabucto fault (Minas geofracture) in Cape Breton Island or south of this E - W latitude (Fig. 1B). The CobequidChedabucto fault shows major right lateral displacement in Pennsylvanian-Permian times (Webb 1969) (Fig. 1B). Permian deformation of pre-Carboniferous rocks has been documented in New England. A widespread Permian metamorphism was syntectonic to deformation along the Lake Char-

235

Honey fault system in eastern Connecticut (Wintsch & Lefort, this volume) (Fig. 1) and associated faults in western Rhode Island (O'Hara 1983; Hermes & Gromet 1983). In addition, post-Acadian granites intruding Avalonian rocks adjacent to the Narragansett basin are deformed with the older rocks indicating that much of the deformation affecting the basement is also Permian in age (Dr eier 1983).

Stratigraphy The stratigraphic succession in the Saint John basin of south-eastern New Brunswick (Table 1) ranges from Tournaisian (Early Mississippian) to Westphalian C (Middle Pennsylvanian) (Rast et al. 1978). The oldest units exposed are a red-bed and volcanic sequence which interfingers with polymict conglomerates. Pillow basalts and felsic tufts and flows are interlayered with the red-beds. Paraconformably overlying these units is a sequence of early Pennsylvanian sandstones, siltstones, and conglomerates which vary in colour from red to

TABLE 1 Rhode I s l a n d Narragansett Basin

Time Deriod/Stage

Quinn and Oliver, 1962 Mutch, 1968

New Brunswick

St. John Basin

Skehan et al., 1979 Mosher, in press

Dighton Cg

IDighton

Rhode Island Fm

Rhode

al.,

1978

Parker, in press

Purgatory Cg

Cg

g

et

I South

North u

Rast

Island

Fm

i

> m

== 0a

l

,t

Pondville Arkose

.,~

Lancaster Fm !~amsutta/Purgatory

Fm (north)~ Pondville

Fm

Wamsutta Fm

(south) Cg

Lancaster Pondville Cg

Fm

Lorneville Volcanics Lorneville

Sediments

Balls Lake Fm

West Beach Fm

,7 i !

236

S. M o s h e r & N . R a s t

green. Quartz monzonites and quartz diorites intrude this sequence. A gentle unconformity, dated as sub-Westphalian A, separates these rocks from thickly bedded grey sandstones and interbedded red siltstones and sandstones of Westphalian A to C ages. An angular unconformity separates Westphalian C sediments from the Triassic units. Ages of Carboniferous and Triassic rocks are based on plant fossils. The stratigraphic succession in the Narragansett basin complex of Rhode Island and Massachusetts (Table 1) ranges from Westphalian A to Stephanian B (Lyons et al. 1976; Skehan et al. 1979). The basin complex can be divided into two separate grabens of different age based on palaeontological, stratigraphical, and structural relationships (Mosher 1983). In the older northern graben, the oldest exposed units are a red-bed and volcanic sequence which interfingets with coarse, polymict conglomerate and coarse-grained sandstone sequences (Mutch 1968). Basalt and rhyolite flows are interlayered with the red-beds (Quinn & Oliver 1962; Mutch 1968). Overlying and interfingering with these sediments are massive grey sandstones, carbonaceous shales, siltstones, coals and conglomerates (Mutch 1968). In the younger southern graben, the oldest units exposed are Westphalian D (Table 1) (Lyons et al. 1976; Skehan et al. 1979), and boulder to pebble quartzite conglomerates and coarsegrained grey sandstones predominate (Mosher 1983). In the centre of the southern graben, carbonaceous shales, coals, fine-grained grey to green sandstones, and thin calcareous units interfinger with the coarser-grained units (Mosher 1983). The total stratigraphical section in the younger southern graben is apparently less than that of the northern graben (Mosher 1983). Sedimentation in the Saint John basin and in the lower portion of the northern Narragansett Basin graben is similar and consistent with that of pull-apart grabens at continental margins. The Middle to Upper Pennsylvanian sediments in both Narragansett basin grabens are the result of continued rifting and fluvial deposition; no units of this age are exposed in the Saint John basin. Other pull-apart grabens in the Canadian Maritimes, however, contain Middle to Upper Pennsylvanian age sediments that are the result of slow basin subsidence and primarily fluviatile-lacustrine deposition. In New England the Pennsylvanian-age Norfolk basin of Massachusetts (Fig. 1A) has the same stratigraphic section as the lower portion of the northern Narragansett basin (Mutch 1968) and is most likely an eroded remnant of a once

larger Narragansett basin (Skehan et al. 1979). The other adjacent Boston, Woonsocket, and Situate basins are now considered Precambrian in age rather than Carboniferous (Kaye & Zartman 1980; Kaye 1981).

Deformation Deformation in south-eastern New Brunswick consists of three major folding episodes (Figs 2 & 3) and one late stage kinking event. The first deformation produced isoclinal folds (F1) with a well-developed axial planar schistosity ($1). Folds trend NE and are recumbent with a NW vergence. Thrusting with a NW direction of transport coincided with the folding, and resulted in a series of complex overthrust sheets which juxtapose slices of Precambrian, early Palaeozoic, and Carboniferous age rocks (Rast & Grant 1973b, 1977). The second, nearly coaxial, deformation refolded the earlier structures as well as the first schistosity; an axial planar crenulation cleavage is well developed locally ($2). Folds (F2) are asymmetric to overturned with vergence towards the SE. Both sets of structures were later folded by upright, conjugate N-S- to NW-SE-trending folds (F3) and with a poorly developed axial planar cleavage ($3). Interference of F 2 and F3 form type 1 (Ramsay 1967) outcrop patterns (Fig. 3). All S surfaces have been further deformed by later kinking (Rast et al. 1978). Deformation in Rhode Island consists of three major folding episodes (Figs 2 & 4), one late stage of kinking, and a later stage of boudinage (Mosher 1983). The first deformation produced tight to isoclinal folds (F1) with a well-developed axial planar schistosity ($1). Folds trend N10~ and are overturned to recumbent with a WNW vergence. Thrusting with a WNW direction of transport coincided with the folding. The second, nearly coaxial, deformation refolded the earlier structures as well as the first schistosity; an axial planar crenulation cleavage ($2) is pervasive throughout the area. Folds (F2) are asymmetric to overturned with vergence towards the east (Fig. 4). Both sets of structures were later folded by upright, conjugate, NE-ENE-trending folds (F3) with a poorly developed axial planar cleavage ($3). This third deformation is only locally developed along NNE- and NE-trending faults, and type I interference patterns are observable on an outcrop scale. Kinking followed folding and was for the most part progressive with the third deformation. Unlike south-eastern New Brunswick, major extension in a N - S direction

Deformation of Carboniferous rocks in New England

237

FIG. 2. Isograd and deformation map of multiply deformed areas in New Brunswick and Rhode Island; Carboniferous rocks shaded. (1 and 2 refer to fold generations F 1 and F 2 respectively; Isograds: B--biotite, Ctd.--chloritoid, G--garnet, S--staurolite; Si--sillimanite. Locations. DH--Dipper Harbor; M--Mispec; SJ--St Johns; N--Newport; W--Warwick.) Note structure has been simplified for clarity; complex nature of the deformation shown in Fig. 3 (from New Brunswick, map location, X) and Fig. 4 (from Rhode Island, cross-sections a and b). in Rhode Island produced meso- to mega-scale boudinage of all the above structures (Farrens 1982; Farrens & Mosher, in prep.). Northeastand ENE-trending, right lateral strike-slip faults also cut the early structures (Fig. 2); these faults cause many small-scale structures which are apparently synchronous with the boudinage. The three deformations reported previously by Skehan & Murray (1979) and Murray & Skehan (1979) (based on reconnaisance work) are similar but not equivalent to those discussed above (see Burks & Mosher 1981; Farrens 1982; Farrens & Mosher 1982). Detailed studies (Burks 1981; Farrens 1982; Berrybill & Mosher 1983) have shown most of those deformations

were part of a single progressive event and have documented other deformations. In Rhode Island the intensity of the first and second deformations varies with lithology and with location relative to basement faults which bound intrabasinal horsts and grabens. The features of the third deformation, and to some extent the boudinage, are localized along these pre-existing faults. Deformation has been grouped into four distinct phases based on overprinting relationships (Mosher 1983). F 1 folds and associated thrusting and F 2 folds are caused by basin closure; F3, kinking, and many small-scale structures (Burks 1981; Farrens & Mosher, in prep.; Mosher 1983) are the result of

238

S. Mosher & N. Rast

FiG. 3. New Brunswick: structural map showing trends of fold generations and thrusts; type I interference pattern. Location of map shown in Fig. 2 by X. left lateral motion on NNE- and NE-trending basement faults. The right lateral strike-slip faults and related structures and the boudinage are caused by N E - E N E - t r e n d i n g right lateral shear couples, In New Brunswick, the original basin configuration is obscured by the later deformation,

F~ and F 2 folds and the thrust slices are presumably caused by closure of the basin. F 3 folds and kinking are either related to continued closure or possibly to some later strike-slip motion. The polyphase deformation of Carboniferous rocks in south-eastern New Brunswick and in

FIG. 4. Rhode Island: Structural cross-sections showing interference of two fold generations and effect of lithologies on style of deformation. Upper section (A) through massive conglomerate unit, lower (B) through shales, sandstones, and minor conglomerates. Thrusts and high-angle faults common in conglomerate section (from Farrens 1982; Fattens & Mosher, in prep.). F 3 Fold axes are at a high angle to these sections. Locations of sections shown in Fig. 2(B).

Deformation o f Carboniferous rocks in New England Rhode Island is in marked contrast to the relative lack of deformation affecting the majority of Carboniferous rocks in the Canadian Maritimes and New England. In the central New Brunswick syncline (Fig. 1) dips rarely exceed 10 ~ All moderately deformed areas in the Canadian Maritimes are adjacent to faults (Fig. 1), and the single phase of folding can be attributed to strike-slip movement or minor thrusting. In New England the northern Massachusetts portion of the Narragansett basin and the adjacent Norfolk basin show a single phase of folding with ENE-trending axes. There, however, the deformation is apparently related to basin closure. The thicker stratigraphic section in the northern portion of the Narragansett basin and the oblique orientation of that portion of the basin and of the Norfolk basin relative to the approximately E - W stresses accounts for most of the difference in fold axis orientation and in deformation intensity (Mosher 1983). As the third and fourth deformations affecting the southern portion of the Narragansett basin are localized along faults, it is possible that these deformations did affect the northern portion but have not been recognized due to poor exposure.

Metamorphism and igneous activity In south-eastern New Brunswick two syntectonic to post-tectonic phases of metamorphism can be recognized. The phase which is syntectonic with D 1 and D 2 (M1) results in the formation of extensive slates and fine-grained phyllites and mylonites (Rast & Dickson 1982). The post-tectonic (M2) phase results in the generation of biotite, chloritoid and finally garnet in pelites (Rast et al. 1979). The grade of M 2 in the area is progressively higher to the SE (Fig. 2). Indeed to the NW Carboniferous strata rapidly lose their state of deformation and in central New Brunswick syncline (Fig. 1), there are no signs of regional metamorphism. The granitoids in south-eastern New Brunswick vary from those with fine hypidiomorphic texture to granophyres and the occasional dioritic bodies are also fine grained. Therefore all these bodies are high-level intrusives (J. S. D. Parker 1982, pets. comm.). In Rhode Island, Barrovian metamorphism (M1) post-dates the first deformation. In the south-western portion of the Narragansett basin, M 1 pre-dates the second deformation whereas in the south-eastern and south central portion of the basin, M 1 is synchronous with the second deformation. The metamorphism

239

increases to the SW reaching sillimanite grade (Fig. 2). All isograds are apparently faultbounded by late stage normal and strike-slip faults (Mosher 1983). The northern Massachusetts portion of the basin is affected by anchizone metamorphism (Skehan & Murray 1979). A later retrograde metamorphism (M2) affected the basin sediments after the third deformation. The metasediments in the southern portion of the basin are intruded by a two mica, S type granite, the Narragansett Pier granite, which truncates isograds (Skehan & Murray 1979). The intrusion post-dates D2; however, the relationship to the subsequent deformation is currently unknown. The granite has been dated at 272 ___4 Myr (Hermes et al. 1981). A basic difference between the Rhode Island and New Brunswick regional metamorphism is that in Rhode Island the M 1 metamorphism essentially increases to the W and SW and is progressive whereas in New Brunswick M 2 decreases to the SW. The more or less uniform M 1 of New Brunswick is very low-grade and is uniform throughout the deformed belt. Extension of the metamorphism to the west of the Narragansett basin beyond the traces of the Bloody Bluff, Lake Char, Honey Hill faults (Wintsch & Fout 1982) implies that these faults, traditionally taken as boundaries of Permo-Carboniferous deformation, in fact lie within the Variscan-Alleghanian orogenic belt.

Discussion The basins in the Canadian Maritimes and in south-eastern New England are similar sedimentologically and appear to be pull-apart grabens related to NE-trending transcurrent or normal faults. The sense of motion on the faults differs in the two areas, however, and suggested right lateral motion in the Canadian Maritimes occurred during the Mississippian, whereas left lateral motion in the Canadian Maritimes and in south-eastern New England occurred during the Pennsylvanian times. The formation of the Canadian Maritime basins and those of New England may be related to the interaction between an Avalonian terrane or terranes and North America but the details are most complicated. After basin formation and the termination of sedimentation, south-eastern New Brunswick and south-eastern New England have had a similar deformational history. The first two foldings (F 1 and F2) and all thrusting were caused by compression in approximately an

240

S. Mosher & N. Rast

~'

~ ......

Fro. 5. Block diagram showing overthrusting in south-eastern New Brunswick and associated right lateral transcurrent fault (Minas geofracture) which forms one boundary of the overthrust plate. Dotted line is Bay of Fundy shoreline (compare with Fig. 1). Other major faults shown; on Belleisle fault (1) is pre-Pennsylvanian movement, (2) is post-Pennsylvanian movement. E - W to E S E - W N W direction. The lack of compressional deformation in the rest of the Canadian Maritime basins such as the central New Brunswick syncline indicates that New England and south-eastern New Brunswick were parts of the orogenic belt while other Maritime basins were platformal. Because all areas affected by polyphase deformation are adjacent to, and south of, the E - W latitude parallel to the Cobequid-Chedabucto fault, which underwent as much as 130 km of right lateral displacement in the Permian (Webb 1969), this fault is interpreted as a fundamental boundary between two distinct tectonic terfanes. It is possible that the deformation involving the Pennsylvanian-Permian movement on the Cobequid-Chedabucto fault (Minas geofracture) was synkinematic with the overthrusting in south-eastern New Brunswick. Thus the right lateral transcurrent fault forms one boundary of the overthrust plate (Fig. 5). The compressional deformation in the two polyphase deformational areas (New Brunswick and Rhode Island) appears to be related to the closure of a Variscan ocean which resulted in the collision of rigid Africa and the Meguma terrane with North America. The increase in metamorphic grade across the Narragansett basin and into the basement to the west and the associated ductile deformation of basement rocks in that area suggest any Permian suture must be west of the Narragansett basin. In south-eastern New Brunswick the pattern of

metamorphism and deformation suggest collision occurred near that area, but somewhat to the east. Later deformation in south-eastern New Brunswick and in south-eastern New England is caused by shearing (Arthaud & Matte 1977). Local variations in the effects of shearing are observed within and between the two polyphase deformation areas. Most of the deformation affecting the other Maritime basins is the result of this later shearing. In Rhode Island the left lateral deformation (D3) is apparently related to Riedel shearing associated with a right lateral E - W - t r e n d i n g transcurrent fault zone. The latter strike-slip motion is the cause of the fourth deformation (Mosher 1983). More detailed mapping and geometric analysis is necessary before the complete stress history of these two areas can be accurately assessed. Although many geotectonic models have been proposed for the Variscan/Alleghenian orogeny (see for example Lefort & Van de Voo 1981), all are speculative. The deformational history of south-eastern New Brunswick and south-eastern New England puts constraints on such models, and these two areas must be considered together in reconstructing the North Atlantic tectonics during the Permian.

Conclusions Comparison of Variscan-Alleghenian structures from two multiply deformed areas in

Deformation of Carboniferous rocks in New England s o u t h - e a s t e r n N e w B r u n s w i c k and R h o d e Island shows that C a r b o n i f e r o u s rocks in the two sedimentologically entirely separate basins w e r e affected by a similar set of stresses in the late C a r b o n i f e r o u s - P e r m i a n times. E a r l y d e f o r m a t i o n (D a and D2) and associated m e t a m o r p h i s m resulted f r o m closure of a Variscan o c e a n w h e r e a s later shearing d e f o r m a t i o n (D 3 and D4) r e s u l t e d f r o m m o v e m e n t on preexisting faults at the e n d of collision. T h e senses of shearing on faults r e l a t e d to these late k i n e m a t i c e v e n t s are not necessarily an indication of the earlier m o v e m e n t s . In s o u t h - e a s t e r n N e w E n g l a n d the increase in m e t a m o r p h i s m w e s t w a r d across the N a r r a g a n s e t t basin and into the b a s e m e n t as well as ductile d e f o r m a tion of b a s e m e n t rocks to the west of the Narragansett basin leads us to c o n c l u d e that any P e r m i a n suture must be west of the N a r r a g a n sett basin. In the C a n a d i a n M a r i t i m e s the localization of m e t a m o r p h i s m in s o u t h - e a s t e r n N e w

241

B r u n s w i c k and its e a s t w a r d increase suggests that this area r e p r e s e n t s the w e s t e r n p o r t i o n of the collision zone. B o t h areas w e r e parts of the V a r i s c a n - A l l e g h e n i a n o r o g e n i c belt, and the similar d e f o r m a t i o n a l histories r e q u i r e that s o u t h - e a s t e r n N e w Brunswick and R h o d e Island be c o n s i d e r e d t o g e t h e r w h e n reconstructing N o r t h Atlantic tectonics during the Permian. ACKNOWLEDGMENTS: The authors wish to thank A. W. Berryhill, R. J. Burks, W. L. Dickson, R. B. Dreier, C. M. Farrens, M. C. Henderson, S. Parker, Hau Chong Teng, and K. J. Thomas for access to unpublished data and for many excellent discussions. R. H. Grant, O. D. Hermes, D. P. Murray, J. W. Skehan, and R. P. Winstsch are thanked for many stimulating and thought-provoking discussions. One author (S. M.) acknowledges the Donors of the Petroleum Research Fund, administered b~, the American Chemical Society, for support of part of this research.

References ARTHAUD F. & MATTE, P. 1977. Late Paleozoic strike-slip faulting in southern Europe and northern Africa; results of a right-lateral shear zone between the Appalachians and the Urals. Bull. geol. Soc. A m . 88, 1305-20. BAIRD, D. M. 1959. Development of gypsum deposits in southern Newfoundland. Trans. Can. Inst. Min. Metall. 62, 257-64. BELT, E. S. 1968. Post-Acadian rifts and related facies, Eastern Canada. In: E-AN ZEN et al. (eds) Studies o f Appalachian Geology: Northern and Maritime, 95-116. Wiley, New York. BERRYH[LL, A. N. & MOSHER, S. 1983. Fault-related polyphase deformation on Dutch Island. Abstr. Progr. geol. Soc. Am. 15, no. 3, 129. BRADLEY, D. C. 1982. Subsidence in Late Paleozoic Basins in the Northern Appalachians Tectonics, 1, 91-105. BURKS, R. J. 1981. Alleghenian deformation and metamorphism in southwestern Narragansett Basin, Rhode Island. M.A. Thesis. University of Texas at Austin. 93 pp. - - & MOSHER, S. 1981. Alleghenian deformation of the southwestern Narragansett Basin and surrounding basement. Abstr. Progr. geol. Soc. Am. 13, no. 7, 420. CURRm, K. U 1977. A note on Post-Missisippian thrust faulting in northwestern Cape Breton Island. Can. J. Earth Sci. 14, 2937-41. DREmR, R. B. 1983. The Blackstone Series: evidence for Avalonian tectonics in northern Rhode Island. Abstr. Progr. geol. Soc. Am. 15, no. 3,129. FARRENS, C. M. 1982. Styles of deformation in the Southeastern Narragansett Basin, Rhode Island and Massachusetts. M.A. Thesis. University of Texas at Austin, 66 pp. - & MOSHER, S. 1982. Alleghenian deformation

in southeastern Narragansett Basin, R. I. Abstr. Progr. geol. Soc. Am. 141 no. 1, 117. & 1984. Styles of deformation in the Southeastern Narragansett Basin, Rhode Island and Massachusetts. (in prep.). FYSON,W. K. 1967. Gravity sliding and cross-folding in Carboniferous rocks, Nova Scotia. Am. J. Sci. 265,

1-11.

Gussow, W. C. 1953. Carboniferous stratigraphy and structural geology of New Brunswick, Canada. Bull. Am. Ass. Petrol. Geol. 37, no. 7, 1713-816. HERMES, O. D., GROMET, L. P. & ZARTMAN,R. E. 1981. Zircon geochronology and petrology of plutonic rocks in Rhode Island. In: BOOTHROYD, J. C. & HERMES, O. D. (eds) Guidebook to Geologic Field Studies in Rhode Island and Adjacent Areas; New England Intercollegiate Geologic Conference 73rd Annual Meeting, 315-38. - & -1983. Recognition and comparison of the Late Precambrian and Paleozoic Plutonic terrains in Rhode Island. Abstr. Progr. geol. Soc. Am. 15, no. 3, 136. KAYE, C. A. 1981. Bedrock geology; Boston North, Boston South, and Newton Quadrangles, Massachusetts. U.S. geol. Surv. Map MF-1241. - & ZARTMAN, R. E. 1980. A late Proterozoic Z to Cambrian age for the stratified rocks of the Boston Basin, Massachusetts, USA. In: WONES, D. R. (ed.) The Caledonides in the USA. Mere. geol. Sci. Virginia Polytechnic Institute and State University Department. 2, 257-64. KEPPIE, J. D. 1981. Tectonics of Nova Scotia--a review. Abstr. Progr. geol. Soc. Am. 13, no. 3, 140. - - - , RU1TENBERG, A. A., FYFFE, L., MCCUTCHEON, S., ST JULIEN, P., SKIDMORE, B., BELAND, J., HUBERT, C., WILLIAMS, U. ~; BURSHALL, J.

242

S. M o s h e r & N. R a s t

1982. Structural Map of the Appalachian Orogen & -1973b. Transatlantic correlation of the in Canada. Map No. 4, scale 1:2,000,000. Varsican-Appalachian Orogeny. Am. J. Sci. 273, Memorial University of Newfoundland. 572-9. LEFORT, S. F. & VAN DER VOO, B. 1981. A kinematic & 1977. Variscan-Appalachian and model for the collision and complete suturing Alleghenian deformation in the northern between Gondwanaland and Laurussia in the Appalachians. La Chaine Varisque d'Europe Carboniferous. J. Geol. 89, 537-50. Moyenne et Occidentale, Editions du C.N.R.S. LEGUN, A. S. & RUST, B. R. 1982. The Upper Carno. 243, 583-6. C.N.R.S. Paris. boniferous Clifton Formation of northern New - - , STEPHEN, J., PARKER, D. & TENG, H. C. Brunswick: coal-bearing deposits of a semi-arid 978. The Carboniferous deformed rocks west of alluvial plain. Can. J. Earth Sci. 19, 1775-86. St. John, New Brunswick. In: LUDNAM, A. (ed.) LYONS, P. C., TIFFNEY, B. & CAMERON, B. W. 1976. Guidebook for Field Trips in Southeastern Maine Early Pennsylvanian age of the Norfolk Basin, and Southwestern New Brunswick, New England southeastern Massachusetts, based on plant Intercollegiate Geological Conference, 70th megafossils. In: LYONS, P. C. & BROWNLOW, Annual Meeting, 162-83. A. H. (eds) Studies in New England Geology." , --, PARKER, J. D. S. & TANG, H. C. 1979. Mere. geol. Soc. Am. 146, 181-97. The Carboniferous Succession in southern New MCMASTER, R. L., DEBOER, J. & COLLINS, B. P. Brunswick and its state of deformation. Ninth 1980. Tectonic development of southern NarInt. Congr. Carboniferous Stratigraphy and ragansett Bay and offshore Rhode Island. GeolGeology, 252. ogy, 8, 496-500. & DICKSON, W. L. 1982. The Pocologan MyloMOSHER, S. & WOOD, D. S. 1976. Mechanisms of nite Zone. ln: ST JULIEN, P. & BELAND, J. (eds) Alleghenian deformation in the Pennsylvanian of Major Structural Zones and Faults of the NorthRhode Island. In" CAMERON, B. (ed.) Geology o f ern Appalachians, Spec. Pap. geol. Ass. Can. 24, Southeastern New England; New England Inter249-61. collegiate Geological Conference, 68th Annual 1983. Variscan Orogeny. In: HANCOCK, P. L. Meeting, 472-90. (ed.) The Variscan Fold Belt in the British Isles, 1981. Late Paleozoic deformation of the Nar1-19. Hilger, Bristol. ragansett Basin, Rhode Island. Abstr. Progr. RUST, B. R. 1981. Alluvial deposits and tectonic geol. Soc. Am. 13, no. 7, 515. style: Devonian and Carboniferous successions 1983. Late Paleozioc deformation of the Narin eastern Gaspe. In: MIALL, A. D. (ed.) ragansett Basin, Rhode Island and MasSedimentation and Tectonics in Alluvial Basins. sachusetts. Tectonics, 2. Spec. geol. Ass. Can. Pap. 23, 49-76. MURRAY, D. P. & SKEHAN, S. J. J. W. 1979. A traverse SCHENK, P. E. 1970. Carbonate-sulfate-redbed facies across the eastern margin of the Appalachianand cyclic sedimentation of the Windsorian Stage Caledonide orogen, southeastern New England. (Middle Carboniferous), Maritime Provinces. In" SKEHAN, S. J. J. W. & OSBERG, P. H. (eds) Can. J. Earth Sci. 6, 1037-66. The Caledonides in the U.S.A., Geological SEVERSON, R. H. & BOOTHROYD, J. C. 1981. DeposiExcursions in the Northeast Appalachians. tional environments, facies associations, and coal I.G.C.P. Project 27, 1-21. Boston College. occurrence in Carboniferous sediments of the MUTCH, T. A. 1968. Pennsylvanian non-marine sediNarragansett Basin. Abstr. Progr. geol. Soc. Am. 13, no. 3, 176. ments of the Narragansett Basin of Massachusetts and Rhode Island. In: DE VRIES SKEHAN. S. J. J. W. & MURRAY, D. P. 1979. Geology KLEIN, G. (ed.) Late Paleozoic and Mesozoic of Narragansett Basin, southeastern MassachuContinental Sedimentation, Northeastern North setts and Rhode Island. In: CAMERON, B. (ed.) America. Spec. Pap. geol. Soc. Am. 106, Carboniferous Basins of Southeastern New 177-209. England, Guidebook for Fieldtrip Number 5; O'HARA, K. 1983. Ductile deformation in Avalonian Ninth Int. Congr. Carboniferous Stratigraphy and gneisses, NW Rhode Island/NE Connecticut and Geology, 7-35. its relationship to the Lake Char Faults. Abstr. 1980. A geologic profile across southProgr. geol. Soc. Am. 15, no. 3, 129. eastern New England. Tectonophys. 69,285-319. QUINN, A. W. 1971. Bedrock geology of Rhode - - , MURRAY, D. P., HEPBURN, J. C., BILLINGS, Island. Bull. U.S. geol. Surv. 1265, 68 pp. M. P., LYONS, P. C. & DOYLE, R. G. 1979a. & OLIVER, W. A. (Jr) 1962. Pennsylvanian The Mississippian and Pennsylvanian (Carbonrocks of New England. In: BRANSON, C. C. (ed.) iferous) systems in the United States--MassaPennsylvanian System in the United States, chusetts, Rhode Island, and Maine. Prof. Pap. 60-73. American Association of Petroleum U.S. geol. Surv. IlI0-A, 30 pp. Geologists. WEBB, G. W. 1963. Occurrence and exploration sigRAMSAY, J. G. 1967. Folding and Fracturing of nificance of strike-slip faults in southern New Rocks. McGraw-Hill, New York. 568 pp. Brunswick, Canada. Bull. Am. Ass. Petrol. Geol. RAST, N. & GRANT, R. H. 1973a. The Variscan Front 47, 1904-27. in southern New Brunswick. In: RAST, N. (ed.) 1969. Paleozoic wrench faults in Canadian N.E.I.G.C. Field Guide to Excursions, 4-11. Appalachians. In: KAY, M. (ed.) North Atlan-

-

-

-

-

-

-

-

&

-

-

D e f o r m a t i o n o f Carboniferous rocks in N e w E n g l a n d tic--Geology and Continental Drift. Mem. Am. Ass. Petrol. Geol. 12, 754-86. WINTSCH, R. P. & Foul, J. S. 1982. Structure and petrology of the Willimantic Dome and the Willimantie Fault, eastern Connecticut. Guide-

243

book for Fieldtrip h~ Connecticut and South Central Massachusetts," New England Intercollegiate Geological Conference, 74th Annual Meeting, 9-1-18.

SHARON MOSHER, Department of Geological Sciences, University of Texas at Austin, Austin, Texas 78712, U.S.A. NICHOLAS RAST, Department of Geology, University of Kentucky, Lexington, Kentucky 40506, U.S.A.

A clockwise rotation of Variscan strain orientation in SE New England and regional implications R. P. Wintsch & J.-P. Lefort SUMMARY: A temperature-time-strain path is described for the Honey Hill fault system in eastern Connecticut, U.S.A. The path is based on 4~ mineral age data, on metamorphic petrology and on the orientation of mineral lineations and other small-scale structures in ductile fault zones. Collectively, the data show that the thrusting direction rotated from ESE at 290 Ma to due south at 250 Ma. These results coincide in the sense of rotation and in the orientation of movement with predictions emerging from hypotheses concerning the collision of Gondwanaland with Laurussia. If southern New England can be shown not to have rotated during this collision, then the data from eastern Connecticut would suggest that the WNW approach of Gondwanaland toward Laurussia turned north only in mid-Permian times.

The uplift of eastern New England which ended the Alleghenian (Variscan) orogeny is now well documented in southern Maine (Dallmeyer & Van Breeman 1981), R h o d e Island (Dallmeyer 1982), and south central Connecticut (see below). However, the nature and causes of this uplift are not well understood. Mapping of the structure and petrology of several ductile fault zones in eastern Connecticut (Dixon & Lundgren 1968; Lundgren & Ebblin 1972; Wintsch 1979a,b; Hudson 1982) has demonstrated that these fault zones were repeatedly activated from high to low grade metamorphic conditions and a retrograde metamorphic path has now been established for the terrane (Wintsch & Fout 1982). With the cooling uplift history of eastern Connecticut now available (see below), the deformation in these fault zones can be recognized as Permian, or Alleghenian in age. Small-scale structures in these fault zones show that the transport direction was dominantly to the SE. Together these data allow construction of a t e m p e r a t u r e - t i m e - s t r a i n path for these faults which may have implications for largerscale Variscan tectonics. The faults to be described in eastern Connecticut are part of a larger fault system extending from south central Connecticut to Newbury, Massachusetts, a distance of 250 km. The system contains several discrete faults and fault zones, the most important of which a r e shown in Fig. 1. The faults dip 3 0 - 6 0 ~ to the north or NW, and movement is in general to the SE on all faults (Castle et al. 1976; Dixon & Lundgren 1968; Wintsch 1979b). Thrust and reverse fault motion cuts local stratigraphic units and juxtaposes terranes of contrasting metamorphic grade. Mylonites and less common breccia zones mark the location of the faults in the north (Castle et al. 1976) but the

fault rocks become increasingly schistose to the south where ductile deformation dominates in these more uniformly higher grade rocks. The rocks SE of this fault system belong to the late Precambrian Avalon terrane (e.g. Rast & Skehan 1981) and the rocks to the N W belong to the Lower and Middle Palaeozoic Merrimac group. Simpson, Shride & Bothner (1980) propose that this zone connected with the south Atlas shear zone in late Palaeozoic times. Skehan & Murray (1980) believe that this zone may be the site of both the Acadian and the Variscan suture between North America and Avalonia, and Lefort (1981) proposes that the rocks SE of the fault system were part of Gondwanaland in pre-Carboniferous times. Thus this fault system probably contains fundamental plate boundaries deformed several times during the Palaeozoic, and the structures in this zone may record events which have regional significance.

Honey Hill fault system In eastern Connecticut the H o n e y Hill, Tatnic and Willimantic faults mark the northern and western limit of exposure of largely gneissic and granofelsic rocks of late Precambrian (Avalonian) age. These rocks are dominated by metavolcanic and metaplutonic rocks, but also contain quartzites and pelitic schists (see Dixon & Lundgren 1968). North and west of these faults metapelitic rocks dominate, but metagreywackes, metasiltstones and rare quartzites and marbles are also present. The H o n e y Hill, Tatnic and Willimantic faults all occur at the base of this metasedimentary sequence, in the l o w e r part of the Tatnic Hill formation (ruled, Fig. 1). Because of their occurrence at this 245

R. P. Wintsch & J.-P. Lefort

246

FIG. 1. Map showing the distribution of selected thrust (teeth on upper plate) and high angle faults in south-eastern New England. The Tatnic Hill formation is ruled. Together the Honey Hill-Lake Char-Clinton-Newbury faults from a fault zone at least 250 km long. The inset (after Wintsch 1979a, Wintsch & Kodidek 1981) shows that the Rodgers Pond fault zone (RPFZ) with sinistral strike-slip motion cuts the Honey Hill fault (HHF). It in turn is cut by the Pattaconk Brook fault (PBF) in the north and the Falls River fault in the south, both verging to the SSW. common stratigraphic position, and because of their similarity in style of deformation, they are considered to be segments of a single thrust fault which underlies at least 1500 km 2 of eastern Connecticut (Wintsch 1979b). The Honey Hill and Tatnic segments of this fault system dip 20-50 ~ to the north and NW. The fault dips in all directions away from the Wilimantic dome which is a window through this fault plane into the Avalonian terrane below. The Lake Char fault is joined to the eastern end of the Honey Hill fault by a band of greenschist facies mylonites. Its activity overlapped in time with the Honey Hill fault system, but details are difficult to establish because this fault cuts rocks stratigraphically below the Tatnic Hill Formation.

Temperature-time-strain path Most of the strain in the fault rocks of the Honey Hill fault system records ductile defor-

mation in amphibolite facies metamorphic conditions. Small-scale rootless intrafolial folds are present in many places along these faults, but the most conspicuous features are anastomozing shear zones which cut pre-existing foliation at small angles and isolate and rotate boudins and tectonic blocks (Wintsch 1979b, plate I). Most movement on the fault system brought the hanging wall in the SE, as demonstrated by: (1) The imbrication and thickening of the Tatnic Hill Formation along the Tatnic fault (Dixon 1968). (2) The thrusting of the high grade Tatnic Hill Formation along the Tatnic fault over the lower grade rocks between the Tatnic and Lake Char faults (Dixon 1968). (3) The NW-trending tear faults along the Tatnic fault (Dixon & Lundgren 1968; Wintsch 1979b). (4) The SE vergence of small-scale open folds (Wintsch 1979b; Hudson 1982). (5) The clockwise rotation (when viewed look-

R o t a g o n o f Variscan strain orientation ing north) of boudins and porphyroblasts (Dixon 1968; Wintsch 1979b). (6) Truncation of stratigraphic units against the Tatnic and Lake Char faults (Dixon 1968; Dixon & Lundgren, 1968). (7) The south-eastward displacement of pegmatite dykes by small-scale shear zones (Wintsch 1979b). The Tatnic Hill Formation containing these deformation features was metamorphosed to upper amphibolite facies metamorphic conditions in pre-Alleghenian times. KD'S calculated from microprobe analyses of co-existing biotite and garnet in the assemblage: garnet-biotitesillimanite-K-feldspar-plagioclase-quartz suggest temperatures between 650 and 700~ Mineral assemblages within the ductile shear zones and mylonite zones reflect P - T conditions less than those of their host rocks. In pelitic rocks they include andalusite, kyanite, muscovite and chlorite (Wintsch 1.980) and in calcareous rocks chlorite, epidote and actinolite. The coexistence of fabric forming K-feldspar + chlorite in some shear zones demonstrate lowest greenschist facies conditions prevailed during late strain events. The shear and mylonite zones cut the pre-existing higher grade schistosity and occasionally also older shear zones, but apparently do not inherit any aspects of the former fabric elements. This is so because the minerals defining the new foliation (except porphyroclasts) have largely recrystallized during the deformation, and are thus relicts of neither the former mineralogy nor the former fabric. The wide range of metamorphic conditions recorded in the fabric forming assemblages allows reconstruction in P - T space of the retrograde metamorphic path of this terrane. It passes from upper amphibolite facies conditions (-700~ 6 kB) through the upper pressure region of the andalusite field to lower greenschist facies conditions (Wintsch 1980). Collectively these assmblages which occur exclusively in cross-cutting shear zones demonstrate that the Honey Hill fault system was activated at a variety of metamorphic grades, and possibly continuously during the uplift of this terrane, with the implication that it was such faults which were responsible for this uplift. The minerals defining these retrograde assemblages always define a new mylonitic foliation and commonly define a mineral lineation in the plane of that foliation. The most abundant and conspicuous lineation is defined by sillimanite needles, but hornblende needles also define a lineation in some metadacites immediately below the fault surface. A lineation defined by streaks or trails of biotite flakes is

247

present in quartzofeldspathic schists. Feldspar rods are present in some mica-poor rocks, and always occur with quartz rods. In other micapoor rocks quartz rods occur in a matrix of equant feldspar grains. Streaks of retrograde muscovite and chlorite can occur with the quartz rods, and sometimes form haloes or trails around garnet or feldspar porphyroclasts. The sillimanite a n d hornblende needles reflect deformation in the upper amphibolite facies (-600~ Temperatures calculated from the composition of coexisting rod-shaped alkali and plagioclase feldspar grains are 300-400~ which may reflect the temperature of their formation (Wintsch & Fout 1982). Biotite streaks occur with both sillimanite and quartz-feldspar rods, and record strain at temperatures between 400 and 600~ Quartz rods and muscovitechlorite streaks record strain in middle to lower greenschist facies conditions ( - 4 0 0 ~ The range of temperatures present during the development of these lineations is shown on the vertical axis of Fig. 2. Many of these lineations plunge gently N60~ reflecting the SE transport and extension direction on the fault system (see above). By inference, then, they are stretching linea800 MINERAL

'IB E A R I N G OF'LINE# ' LINEATION~

'LINEATIONS'

\

Q HORNBLEND? -- ~ ~ ~

oow

600

' QUARTZ

T~

FEL DSPAR RO DS

N 50~ - 0~

\

400 i/ QUARTZ RODS; i CHI ORITE, MUSCOVITE STREAKS

N I0OE ~

-- "

200

I

o~o D

350 I M

I

300 ~

I

~.I

250 I

p

I

]~

200 I

150

d

time (re.y)

FIG. 2. The cooling/uplift curve (from Sutter et al. 1984) for south central Connecticut based on mineral and whole rock isotopic data (see text). The range of temperatures prevailing during the syntectonic development of mineral lineations is given on the temperature axis. The bearing of these lineations is given on the right side of the figure. The cooling/uplift curve establishes the age of the mineral assemblages in which the lineations (left side of the figure) occur, as well as the time interval during which compression direction rotated from ESE to south (right side of the figure).

248

R. P. Wintsch & J.-P. Lefort

tions. The following petrographic evidence suggests that the lineations are parallel to the long axis of the strain elipsoid: (1) they are parallel to the x-axis of microboudinage of porphyroclasts; (2) they are parallel to the quartz pressure shadows adjacent to megacrysts; (3) they are parallel to the direction of displacement of flakes of chlorite +_ biotite and muscovite alteration products of adjacent garnet and feldspar porphyroclasts; (4) they are perpendicular to the axis of rotation of garnet and feldspar porphyroclasts. On the basis of both megasopic and microscopic structures, then, these lineations can be shown to be parallel to the local transport direction in the shear zone or mylonite zone in which they occur.

The retrograde metamorphic path defined by these lineations can be used to ascertain their relative age. The sillimanite- and hornblendebearing assemblages are ubiquitous in these rocks. The lower grade lineations are restricted to narrow shear zones which cut the higher grade rocks (Wintsch 1979b), and therefore, post-date the strain recorded in the high grade lineations. Even in small outcrops where crosscutting relations may not be obvious, the lower grade assemblage defines a shallower, and thus later, increment of deformation. Thus the relative chronology of the lineations within the different mineral assemblages can be determined by assessing their metamorphic grades (Fig. 2) as well as by cross-cutting relationships. The orientation of these lineations is not constant. In the NE corner of Connecticut hornblende and sillimanite lineations trend N60~ biotite streaks N25 ~ and quartz-feldspar rods trend N5~ (Hudson 1982). A lesser clockwise rotation of lineation orientation is present along the Honey Hill fault at its west end (Wintsch 1979a). In the Willimantic area sillimanite needles, biotite streaks and quartz-feldspar rods share a constant N55~ trend (Wintsch 1979b), but the trend of the muscovite and chlorite streaks in less common middle and lower greenschist facies mylonite zones is north-south (Wintsch & Fout 1982). Thus in all areas a clockwise rotation of thrust direction is recorded in the progressively lower grade conditions of deformation. The above ranges in bearing of the lineations are summarized in Fig. 2. Support for this clockwise rotation of compression direction is present in the structures in the Chester area (Fig. 1, inset). Motion direction on the Honey Hill fault (as discussed above) was - S 5 0 ~ This fault is cut by faults of the

Rogers Pond fault zone (named here for reference). Lineations associated with this fault zone plunge gently N30~ (Wintsch 1979a). This fault zone is cut by the Falls River fault in the south (Wintsch & Kodidek 1981) and the Pataconk Brook fault to the north. Lineations along these ductile faults trend N0~176 Thus a clockwise rotation of at least 60 ~ is recorded in clearly cross-cutting relationships among several faults in the Chester area, as well as in the mineral lineations in the single Honey Hill-Tatnic-Willimantic fault system. Because the thrusting on the Honey Hill fault which produced the mineral lineations occurred during the uplift of the terrane, the cooling/uplift curve reported by Sutter, Wintsch & Grant (1984) may be used to date absolutely the retrograde mineral assemblages and thus the mineral lineations. This uplift curve (Fig. 2) is based on mineral and whole rock ages of intrusive granites 2 0 k m north of Chester, on 4~ ages of hornblende and biotite from the Chester area, and on the late Triassic sediments in the Hartford basin 20 km west of Chester. This uplift curve shows that upper amphibolite facies conditions ceased in late Pennsylvanian times (280 Myr), and greenschist facies began in late Permian times ( - 2 6 0 Myr). Note that the orientation of mineral lineations through this sequence shows that N W - S E compression persisted until - 2 7 0 Myr (early Permian) and then rotated clockwise to NNE through the Permian (Fig. 2).

Tectonic setting In spite of their highly speculative nature, an examination of Variscan tectonic models is useful as they provide a large-scale tectonic context for southern New England in the late Palaeozoic. The models of Arthaud & Matte (1977), Scotese et al. (1979), Lefort & Van der Voo (1981), Badham (1982) and Dewey (1982) describe the closing of oceans separating plates and microplates through middle Carboniferous times, leading to the creation of Pangea by Westphalian times ( - 2 9 0 Ma). The relative positions of the land masses at this time were probably close to those in the Permotriassic reconstruction of Fig. 3. The position of the southern boundary of Laurussia proposed by Lefort (1981, 1983) suggests that southern New England was unrelated to Laurussia. This is in contrast to the common assignment of this area to the Avalon arc (Rast 1980) or Avalon composite terrane (Keppie 1984). Regardless of the pre-Carboniferous history or correlation

Rotation o f Variscan strain orientation

Q

~EUROPE

/ ..--" NORT••H

......"-...~/.~I - ~ I

AFRICA

249

b

~ ! ~

EUROPE //

system / I.~.~

,

/~

AFRICA

:oo

.oo,,.

FIG. 3. A Permotriassic reconstruction (Mollweide projection) of the north Atlantic region, compiled primarily from Lefort & Van der Voo (1981, fig. 1). East-west textral strike-slip faults from north to south are the Armorican fault, the Cobequid-Chedabucto fault, the mid-Moroccan fault and the South Atlas fault (Lefort & Van der Voo 1981). The NE-trending set are (north to south) the Great Glen fault (Van der Voo & Scotese 1981), the Dover-Hermitage fault (Hammer 1981), the Fredericton-Norumbega fault (Keppie 1984), the New York-Alabama lineament (King & Zeitz 1978) and in Africa, the Zemmour fault (Rod 1962). The hatchered line marks the southern boundary of Laurussia (Lefort 1981). The possible dextral, counter-clockwise motion of Gondwanaland relative Laurussia in the Westphalian (-290 Ma) indicated by the large curved arrow and the ESE transport direction of thrusts in southern New England (small arrow). The possible northward migration of Gondwanaland in the Middle Permian and the southward transport direction of thrusts in southern New England.

of southern New England, there is complete agreement that the late Palaeozoic deformation recorded in its rocks is post-collisional. Other late Palaeozoic structures potentially relevant to the tectonic development of southern New England include east-west and NEtrending strike-slip faults. Most of the evidence along the east-west faults suggests that primarily dextral motion occurred during the Carboniferous (see Lefort & Van der Voo 1981, Badham 1982 and Lefort 1983 for details). In contrast to this motion, movement on the NEtrending faults may be either dextral or sinistral, and their activity probably spanned a greater period of time (Lefort & Van der Voo 1981). The approach of Gondwanaland towards Laurussia took place through the Devonian and

Carboniferous, and by Westphalian times ( - 2 9 5 Ma) all oceans between North America and Gondwanaland had probably closed. Throughout the Carboniferous the motion of Gondwanaland was WNW (Badham 1982, Dewey 1982), and palaeomagnetic evidence exists which suggests that Gondwanaland may also have rotated in a counter-clockwise sense during this northward migration (Lefort 1983). Fig. 3 shows the possible distribution of plates and microplates in the Westphalian. The curved arrow in Africa represents the WNW motion of Africa with its possible rotation. Such motion is in keeping with the strike-slip orogen hypothesis for the Hercynides as proposed by Arthaud & Matte (1977) and Badham (1982). The palaeomagnetic reconstructions of Scotese etal. (1979) and Scotese & Bamback (1979)

250

R. P. Wintsch & J.-P. Lefort

indicate that Gondwanaland continued to move with a strong northward component relative to Laurussia in the late Carboniferous and the Permian. Lefort & Van der Voo (1981) speculate that after the late Carboniferous convergence of Gondwanaland with Laurussia this northward migration produced (or reactivated) among other things the NE-trending strike-slip faults on the eastern margin of North America. Fig. 3 shows schematically this interpretation, and includes many features given by Lefort & Van der Voo (1981, fig. 2). A summary and concensus of these models and ideas concerning the collision of Gondwanaland with Laurussia leads to the proposal that Gondwanaland approached North America from the ESE throughout the Carboniferous, producing a dextral shear margin in the North Atlantic region. In late Carboniferous and probably also Permian times, Gondwanaland moved with a strong northward component into Laurussia activating the NE-trending faults. Of particular relevance to this study is the rotation of compression direction in the northern Appalachians from W N W to due north implied by the above scenario. Fig. 3 indicates the position of the H o n e y Hill fault system as well as the direction of thrusting as discussed above. The coincidence of the sense of rotation and the orientation of compression direction interpreted quite independently is remarkable. However, the data from the H o n e y Hill fault system cannot be taken as compelling evidence for the large-scale model, because southern New England could have rotated during the collision. If these structures do record large-scale movements of Gondwanaland, then the data of Fig. 2 suggest that the turn of Gondwanaland motion to the north did not

occur until Middle Permian times ( - 2 6 0 Ma). Much more field and palaeomagnetic data on late Palaeozoic tectonics are necessary for further exploration of the large-scale model of Fig. 3.

Conclusions Petrological, structural and isotopic data from the H o n e y Hill fault system lead to the conclusion that thrusting direction rotated from E S E to due south between approximately 300 and 250 Ma. This is remarkably consistent with plate tectonic models which speculate that Gondwanaland approached Laurussia from the E S E and then from the south during the Carboniferous. This cannot be taken as compelling evidence of the large-scale model, however, because of uncertainties in the orientation of southern New England during the collision. Much additional field and palaeomagnetic data are required before the large-scale model can be proposed with confidence. Finally, the field and analytical data from southern New England do indicate that the H o n e y Hill fault system was active in the late Palaeozoic, but they do not bear on its pre-Carboniferous activity or on its status as a potential cryptic suture.

ACKNOWLEDGMENTS: We thank H. Stuenitz and especially D. Hutton for helpful discussions and comments on earlier drafts, and T. Brown, W. Moran and R. Hill for help in manuscript preparation. Fieldwork by one of us (R. W.) over the last few years has been supported by Indiana University, the Connecticut Geological and Natural History Survey, the U.S. Geological Survey, and the Nuclear Regulatory Commission.

References ARTHAUD, F. & MATTE, P. 1977. Late Paleozoic strike-slip faulting in southern Europe and northern Africa: result of a right-lateral shear zone between the Appalachians and the Urals. Bull. geol. Soc. Am. 88, 1305-20. BADHAM, J. P. N. 1982. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 493-504. CASTLE, R. O., DIXON, H. R., GREW, E. S., GRISCOM, A. & ZIETZ, I. 1976. Structural dislocations in Eastern Massachusetts. Bull. U.S. geol. Surv. 1410, 39 pp. DEWEY,J. F. 1982. Plate tectonics and the evolution of the British Isles. J. geol. Soc. London, 139, 371-412. DALLMEYER, R. D. 1982. 4~ ages from the

Narragansett basin and southern Rhode Island basement terrane: their bearing on the extent and timing of Alleghenian tectonothermal events in New England. Bull. geol. Soc. Am. 93, 1118-30. & VAN BREEMAN, O. 1981. Rb-Sr whole rock and 4~ mineral ages of the Togus and Hallowell quartz monzonite and Three Mile Pond granodiorite plutons, south-central Maine: their bearing on post-Acadian cooling history. Contr. Miner. Petrol. 78, 61-73. DIXON, H. R. 1968. Bedrock geology of the Plainfield area, Connecticut. Open-File Rep., U.S. geol. Surv. 308 pp. - - & LUNDGREN,L. W. 1968. Structure of eastern Connecticut. In: ZEN, E-AN, WHITE, W. S.,

Rotation o f Variscan strain orientation HADLEY, J. B., THOMPSON,]. B. (JR) (eds) Studies of Appalachian geology, northern and maritime, 219-29. Wiley, New York. HAMMER, S. 1981. Tectonic significance of the northeastern Gander Zone, Newfoundland: an Acadian ductile shear zone. Can. J. Earth Sci. 18, 120-35. HUDSON, M. R. 1982. Mineralogy, petrology and structural geology of the Tatnic Hill Formation, Putnam, Connecticut. Unpublished MA Thesis, Indiana University. 301 pp. KEPPIE, J. D. 1984. The Appalachian collage. In: GEE, D. G. & STURT, B. (eds) The Caledonide Orogen, Scandinavia and Related Areas. Wiley, New York, in press. KING, E. R. & ZEITZ, J. 1978. The New York-Alabama lineament: geophysical evidence for a major crustal break in the basement beneath the Appalachian Basin. Geology, 6, 312-8. LEFORT, J.-P. 1981. La limite meridionale de la Laurussia eutre la Floride et le Bassin d'Aquitaine. Bull. Soc. gdol. Fr. 7, XXIII, 6, 565-70. 1983. A new geophysical criterion to correlate the Acadian and Hercynian orogenies of western Europe and eastern America. In: HATCHER, R. D., WILLIAMS,M. & ZEITZ, J. (eds) Tectonics and Geophysics of Mountain Chains. Mere. geol. Soc. Am. 158, 3-18. d~ VAN DER V00~ R. i981. A kinematic model for the collision and complete suturing between Gondwanaland and Laurussia in the Carboniferous. J. Geol. 89, 537-50. LUNDGREN, L. W. & EBBLIN, C. 1972. Honey Hill fault in eastern Connecticut: regional relations. Bull. geol. Soc. Am. 83, 2773-94. RAST, N. 1980. The Avalonian plate in the northern Appalachians and Caledonides. In: WONES, D. R. (ed.) The Caledonides in the U.S.A. Mere. Virginia Polytechnic Inst. State Univ. 2, 63-6. Blacksburg. & SKEHAN, J. W. 1981. Possible correlation of Precambrian rocks of Newport, Rhode Island, with those of Anglesey, Wales. Geology, 9, 596-601. ROD, E. 1962. Fault pattern, northwest corner of the Sahara Shield. Bull. Am. Ass. Petrol. Geol. 81, 815-30. -

-

-

-

251

SCOTESE, C. R. & BAMBACK, R. K. 1979. Phanerozoic continental drift base maps. In: Paleogeographic reconstruction: State of the Art. Geol. Soc. Am., SE section, short course notes, 40 pp. --, BARTON, E., VAN DER VOO • ZIEGLER, )k. M., 1979. Paleozoic base maps. J. Geol. 87, 217-77. SIMPSON, D. W., SHRIDE, A. F. & BOTHNER, W. A. 1980. Offshore extension of the Clinton-Newbury and Bloody Bluff fault s~fstems of northern Massachusetts. In: WONES, D. R. (ed.) The Caledonides in the U.S.A. Mem. Virginia Polytechnic Inst. State Univ. 2, 229-33. Blacksburg. SKEHAN,J. W. & MURRAY, D. P. 1980. A model for the evolution of eastern margin (EM) of the Northern Appalachians. In: WONES, D. R. (ed.)The Caledonides in the U.S.A. Mere. Virginia Polytechnic Inst. State Univ., Dept o f Geol. Sci. 2, 67-92. Blackburg, Virginia. SUTTER, J. F., WINTSCH, R. P. & GRANT, N. K. 1984. Isotopic study of plagioclase gneiss, south-central Connecticut: implications for Hercynian deformation. (in prep.). VAN DER Voo, R. & SCOTESE, D. 1981. Paleomagnetic evidence for a large (2000 km) sinistral offset along the Great Glen fault during Carboniferous times. Geology, 9, 583-9. WINTSCr~, R. P. 1979a. Recent mapping in the Chester area, Connecticut and its bearing on the Chester Syncline (abstr.) Abstr. Progr. geol. Soc. Am. 11, 60. 1979b. The Willimantic fault: a ductile fault in eastern Connecticut. Am. J. Sci. 279, 367-93. 1980. Retrograde aluminosilicates and low AH20 in ductile shear zones, eastern Connecticut (abstr.) Abstr. Progr. geol. Soc. Am. 13, 184. & FOUT, J. S. 1982. Structure and petrology of the Willimantic dome and the Willimantic fault. In: JOESTEN, R. & QUARRIER, S. S. (eds) Guidebook for Field Trips in Connecticut and Southcentral Massachusetts. Guidebk Conn. Geol. Nat. Hist. Surv. 5, 465-82. & KODIDEK, R. L. 1981. Local and regional implications of recent mapping in the Essex area, Conn. Abstr. Progr. geol. Soc. Am. 13, 184.

-

-

-

-

R. P. WINTSCH,Department of Geology, Indiana University, Bloomington, Indiana 47405, U.S.A. J.-P. LEFORT, Institut de G6ologie, Universit6 de Rennes, 35042 Rennes C6dex, France.

Clues to the deep structure of the European Variscides from crustal seismic profiling in North America J. A. Brewer SUMMARY: COCORP data from two areas of North America where late Palaeozoic deformation occurred provide clues to the deep structure of the Variscides of Europe. These areas are: (1) the southern Appalachians, where deformation culminated in the Alleghenian orogeny, and (2) the Ouachita Mountains, formed during the Ouachita orogeny. The COCORP data show deep structures which were probably formed during collision and overriding of the early Palaeozoic edge of the North American Grenville basement. Two types of reflector sequences are characteristic: (1) subhorizontal or gently dipping reflectors defining the base of crystalline or sedimentary nappes, which in the interior of the orogen pass into: (2) more steeply dipping (25-40 ~ zones of reflectors, 10-20 km thick, that are possibly offshelf, basinal facies metasedimentary rocks, and slices of basement, stacked and imbricated against, and hence marking the edge of, the North American continent. Similar reflector sequences are observed in seismic data recorded from the European Variscides, suggesting that here similar processes occurred at the southern edge of the Old Red Sandstone continent. Changes in the character of the Variscan front can perhaps be explained in terms of the extent of overthrusting of the continental edge (i.e. the extent of nappe formation).

The Variscides of Europe are part of a mountain belt extending from eastern Europe to the Gulf of Mexico. In North America much of this belt formed as a late expression of the Appalachian orogeny whereas in Europe it was a separate entity to the Caledonides which lie further north. It is most instructive to study the North American portions of this belt because: (1) geological exposures are much more continuous than in Europe, and (2) geophysical data, especially deep seismic reflection data, are more abundant. This paper briefly describes C O C O R P (Consortium for Continental Reflection Profiling) deep reflection data collected from two areas of the United States where Variscan deformation occurred. These data are then compared with as yet limited reflection data recorded in the E u r o p e a n Variscides (Fig.

1). Although many more profiles are n e e d e d to confirm findings and comparisons, it is apparent that two basic types of reflection sequence are recorded in the upper and middle part of the crust in these areas, suggesting that orogenic processes occurred in Europe similar to those in North America. Thus, although emphasis has been placed on variations in the surface geology of the E u r o p e a n Variscides, such as circular or arcuate structures which are not seen in, for example, the Appalachians (e.g. Badham 1982), seismic reflection data do indicate somewhat similar deep reflection sequences. The seismic data are c.ompared as line drawings taken without significant modification from the various sources quoted. Reference should be made to these sources for the original data. Variations in the line drawings, e.g. continuity

FIG. 1. Pre-drift reconstruction of North Atlantic region (after Le Pichon et al. 1977) showing location of seismic profiles discussed in text that cross areas of late Palaeozoic deformation. 253

254

J. A . B r e w e r

of particular reflecting horizons, partly reflect variations in data quality as well as variations in the way different authors interpret their data.

COCORP southern Appalachian data These data provide a continuous traverse across the southern Appalachian orogen, from the Valley and Ridge province in the interior of the craton to the Coastal Plain that overlaps the orogen on the east coast (Fig. 2). The Valley and Ridge province consists of a shelf sequence, deposited on the eastern edge of the early Palaeozoic North American continent (Rodgers 1968), that was folded and thrusted westwards (continentwards) during the P e r m o Carboniferous Alleghenian orogeny. To the SE lie the Blue Ridge and Inner Piedmont provinces, which consist of: (1) Lower Palaeozoic sequences that once lay basinward from the continental edge, (2) late Precambrian, possibly rift-related sequences, and (3) slivers of Grenville basement and other continental basement which are caught up in westward-directed thrusts and nappes (Hatcher 1978).

COCORP

OUACHITA

The two basic types of reflection sequences considered in this paper are best demonstrated in this area (Fig. 3). Flat or subhorizontal reflections which are continuous with reflections from Valley and Ridge strata can be traced eastwards, under the high-grade crystalline rocks of the Blue Ridge and Inner Piedmont. The layered reflections are typically 0.5-1.0 s (1.5-3.0 kin) thick, lie between 6 and 10 km deep and indicate at least 200 km of westward-directed transport of the overlying Blue Ridge and Inner Piedmont. Near the Inner Piedmont-Charlotte belt transition (Fig. 2) the flat-lying reflections pass into a zone, between about 8 and 18 km depth, of easterly-dipping reflectors (Cook et al. 1979 and Fig. 3). Further to the east subhorizontal, discontinuous reflections between 12 and 18 km depth possible indicate that the detachment under the Appalachians extends even further east, although other interpretations of the data are possible (Cook et al. 1981). Here we are concerned with the easterlydipping reflectors. They lie basinward of the approximate position of the early Palaeozoic shelf edge obtained by palinspastically restoring folded and thrusted Valley and Ridge rocks

PROFILE

APPALACHIANS

PROFILE

'4,

i l l l H ~ f l / I

/

%

J' ~ + ~ ' . %2%+ ..

_

c"

/ ./ ~+ ,~,'"'" "++

I \

COASTAL PLAIN

/

"~'r~'/Jc'PC'/P~',/J'~' ~

G,~

..~,j

*'v~4r~ Co 4 "9"r4,[ ~,,(4/4,'

BFZ - BREVARD FAULT ZONE o

KM

300

FIG. 2. Location of COCORP lines in the southern Appalachians and the Ouachita Mountains discussed in text. Note that locations are somewhat schematic, and only those parts of lines with data illustrated in Fig. 3 are shown. Several cross-lines have been recorded in the three areas which are not illustrated. Base maps after Bayley & Muehlberger (1968).

255

9 C~

zO

~5

~"

~

o

,C"%,

>,.

J J

9~ = Z

.

, ,,~j, y,p,, .~ " ~ '"~ 00

;~.~

<~

g

, p:,,,,

Z I--'

-~

"

~

_

N "r

o~

D

-o

~.,,

': :

w

o ~ =1 ,,, ' "~e m ..~ ,,,~ &... ~ _o ,.,I ,',,:,,,, ,,',,C,,~' ,._ o ,"

*,3

~.~

I ~"~'A;'f'&

o

(,..)

, \ \', C~',% '.'h

\\ \".

!, ~ll.o

I..g U b_ 0 nr"

',,,, ,,, ~\, ~ '~{'<',~%~~%

z

-<,

~o

x*

., ,.,,.

o- I

,

>,

t~ 9,,.~ ~ m ~

,,,, ~.~

o

o_'~&~ (';)

s163

"6

.~ ~

t!

~

~ ,,,~ "1"

z _z

=o/11\

o,J

{/ttl, ~

~.~ " -

~'~

o

';,

JL,

h

fl

/11

~, #

,=,,

'~..

9o~

r,.~,,~r II,Ii

~

~,

o1~ I I

z

I

,.,4

a'

'

'

~

~

~. "'1 , ~' -.,I I ~El

,

~ I,, ~ ~

~

,;'~,

; ''n

9 ~

~

~.-

I'1"" I

I

~ =< I=.~

~,

j @

~ "1, ,:' .~1

'.

,,

,'

t

,:.4 ~ o "

89':18

(/)

0

I ~;,z

SO:IS

-

,

~J

256

J. A . B r e w e r

(Harris & Bayer 1979). Also in this region, gravity and magnetic data suggest a transition from less dense, less mafic crust on the west to denser, more mafic crust under, and east of, the easterly-dipping reflections. The data as a whole suggest a tectonically buried continental edge (Cook & Oliver 1981). It should be emphasized that in the southern Appalachians the easterly-dipping reflectors have not yet been identified at the surface or by drilling, but this interpretation is supported by C O C O R P data from the northern Appalachians. Here possibly analogous reflecting sequences can be traced to the surface, to the region where the lower Palaeozoic shelf-offshelf transition forms the cover to basement massifs (e.g. the Green Mountains) thrusted westwards during Taconic or Acadian times (Ando et al. 1983). Alternative models for the easterly-dipping reflectors include deep crustal root zones to the Blue Ridge and Piedmont nappes (Hatcher 1981; lverson & Smithson 1982). Such interpretations do not fully account for the crustal reflection character further east, which suggests the detachment under the Blue Ridge and Inner Piedmont might continue under the Coastal Plain. However, until the dipping reflectors in the southern Appalachians have actually been traced to the surface or drilled, such interpretations cannot be ruled out. No direct evidence exists for the timing of the events which caused the dipping reflectors. Deformation in the southern Appalachians started in the Middle-late Ordovician (I aconic phase) and culminated in the Alleghanian orogeny when over 130 km of stratal shortening took place in the Valley and Ridge foreland (Harris & Bayer 1979), in front of the advancing crystalline thrust sheet of the Blue Ridge and Inner Piedmont (Cook et al. 1981). The dipping reflectors could have been formed at this time, or they could have been formed earlier in the Palaeozoic (Cook et al. 1981).

COCORP data in the Ouachita Mountains, Arkansas The Ouachita mountains are the south-westerly continuation of the Appalachians although the exact nature of the junction between the two systems is not known. These mountains are also thought to contain basin-derived metasedimentary rocks thrust over the Palaeozoic North American shelf edge. However, several important differences with the Appalachians exist: no basement rocks are exposed (and hence the

nature of the basement under and south of the Ouachitas is unknown), and much greater regional Mesozoic subsidence has occurred (causing burial of much of the system). C O C O R P profiles run from the middle of the foreland Arkoma basin (which lies on Grenville basement), southwards through Carboniferous thrust-bounded flysch sequences, into the middle of the Ouachitas where Lower-Middle Ordovician deep sea turbidites and fan deposits, greatly disrupted by continentwarddirected thrusts, are exposed in the antiformal Benton uplift (Fig. 2). South of this uplift Carboniferous flysch (which is stratigraphically continuous with the sequence to the north) is again traversed before structures disappear under thick Cretaceous deposits. Some Triassic basins underlie the Cretaceous coastal plain overlap. The seismic data show that early Palaeozoic foreland sequences of the A r k o m a basin can be traced to the south, under the Carboniferous flysch sequences, to the northern edge of the Benton uplift (Fig. 3; Nelson et al. 1982; Lillie et al. 1983). The total sedimentary thickness here is about 14 km. Under the Benton uplift a large antiformal structure exists whose top lies at about 15 km depth. Along strike to the S W i n Texas two similar antiformal seismic structures (the Devil's River and Waco uplifts, Nicholas & Rozendal 1975) have been profiled and drilled. Late Precambrian (possibly Grenville) basement has been drilled in the Devil's River uplift, and early Palaeozoic carbonate sequences similar to those of the foreland were drilled in both, suggesting that the antiformal structures represent extensions of Grenville basement under the Ouachita orogen. By analogy therefore, an anticline of Grenville basement exists under the Benton uplift. This implies that the Lower--Middle Palaeozoic deep marine sequences in the core of the uplift have been thrusted over the basement from the south. South of the anticline a very thick sequence of south-dipping layered reflectors occurs. They project to the surface in a zone of south-dipping thrusts, or :thrust slices, in the Lower-Middle Palaeozoic deep-water sequence or in the Carboniferous tlysch sequence on the south side of the Benton uplift. The reflectors continue to at least 14 km depth and their base may be as deep as 22 km, where gently north-dipping layering is seen. If the reflections at depth are caused by the same rock types as at the surface, they then define a thick prism of metasedimentary rocks which originally lay offshelf from Grenville basement (whose southerly termination is tentatively identified as the southern part

257

D e e p structure o f the E u r o p e a n Variscides

of the antiform under the Benton uplift). This offshelf sequence was imbricated against, and thrust over the edge of, the basement during the Ouachita orogeny. Gravity data are consistent with a model of a transition from thick continental-type crust under the Arkoma basin and Benton uplift to thinner (perhaps oceanic) crust to the south under this thick layered reflector sequence (Lillie et al. 1983). Thus in the Ouachitas extensive nappes of Lower-Middle Ordovician deep marine strata appear to have overthrusted North American basement. This basement has also been deformed into an antiform, perhaps above deep-seated crustally penetrating thrusts that were active during the final stages of the Ouachita orogeny (Nelson et al. 1982; Lillie et al. 1983). Dipping reflector sequences also occur which lie outboard t'rom a continental edge suspected to lie under the southern margin of the Benton uplift. Whether the differences with the southern Appalachians in reflector geometry, and hence structure, represent fundamental differences in mountain building processes, or whether they simply represent differences in factors such as intensity of deformation, thicknesses or basinal sediments, nature of impinging crust, degree of normal faulting in the overridden continental margin, or time span of deformation is not clear at present. However the presence in these two datasets of relatively flat-lying, and more steeply dipping, reflectors in the same relative position, suggests that fundamental structures in the two areas are similar, and perhaps that similar processes formed them (Ando et al. 1983).

Seismic

2~

, . . i

L="

~-'~ /

[

from

the

-.7c ,-~

I

data

Variscides

Few deep reflection data have been recorded in the European Variscides as yet, but those that exist show striking similarities to parts of the COCORP profiles. For example data recorded across the northern margin of the Rhenohercynian zone (Meissner et al. 1980, this volume) show that subhorizontal reflecting horizons at about 1 s define the base of a nappe complex. Around the Lizard Peninsula seismic data contain zones of moderately dipping (20-30 ~) reflectors thought to be thrusts (Leveridge et al., this volume). The Rhenohercynian zone forms the northern external zone of the Hercynian orogen in Germany, and is bordered to the north by the flysch basin of the sub-Variscan foredeep (Fig. 4). To the south lies the Saxothuringian zone. in SW England the major part of the Devonian and Carboniferous outcrop can be considered as the westward extension of the Rhenohercynian zone (e.g. Gardiner & Sheridan 1981) with a small area around Lizard Head as the northern margin of the Saxothuringian zone (Stile 1951), although Matthews (1978, this volume) considers the extrapolation invalid since the intervening Belgium/northern France region has a different character. Although it is generally considered that the Variscan orogeny in Europe developed on the southern margin of the Old Red Sandstone continent (see Anderton et al. 1979, for review), there is little agreement about the location of this southern margin nor whether the basin that

<;"

I

reflection

European

~

"~

.

.

"

v

"~ ...._ .

.

.

. ~ ~

"

.-

,o+

g~o+ ~;~c~t~ ~ .

--.-

.. ........ [. ..... . . . . . . . =. . . . . . . . . . . . . . .

.=,O.\IF'~

~t-0"~'6''" .

-

Lz - LIZAII) COMPLEX .

.

.

.

0

KM

300

FIG. 4. Locations of seismic reflection profiles in the Rhenohercynian and Saxothuringian zones. Lines are: (A) Meissner et al. (1981); (B) Graulich (1980); (C) C16ment (1963); (D) Kenolty et al. (1981); (E) Leveridge et al. (this volume); (F) Shell UK profile.

J . A . Brewer

258

developed outboard of it, and collapsed during the orogeny, was underlain by oceanic crust (possibly of a marginal basin) or thinned continental crust (which would imply an ensialic orogen). Basin development began in midDevonian times, with deformation and uplift starting towards the end of the Devonian (the Bretonic phase), and culminated in northward (i.e. continentward-directed) thrusting

and nappe formation in the late Carboniferous Asturic phase (Ziegler 1978). Sequences deposited during basin development and collapse constitute the Rhenohercynian zone. Sequences deposited further north, comprising a continental shelf succession of predominantly clastic sediments which developed into a shallow-water carbonate shelf succession in the early Carboniferous, and which was then flooded by Midi fault system extension

Rhenohercynian

3~ MEISSNER ET AL

I

9

1981

i

/

3~ Overthrust GRAULICH

~ ~

[

1980

Midi fault system C) S ~ -~-.. -Thrust'A_~J~ N , J/S..-~.- -~'-~\

S

o

d)_

CLEMENT 1963 -;'.-:;---

,

Variscan front /i

_

Thrusts (9)

/

o

N

--~

o

15-[i

I

CHADWICK ET AL

lOkm$

i

~

__

~

~-15

198:5

SSE o

NNW

6,1

LEVERIDGE ET AL

1983

f)WSW z "

61_

ENE

~

SHELL

~

.-.U.K.

.-

.... .

FIG. 5. Line drawings of seismic data from Rhenohercynian and Saxothuringian zones. Letters alongside profiles correspond to locations shown in Fig. 4. These profiles are unmigrated except for (d), and (also except for d) can be crudely converted to depth in kilometres by multiplying two-way travel time in seconds by 3. Horizontal scale in (d) applies to all the profiles. Note that thrusts in (a) to (c) are rather gently dipping while those in (d) to (f) dip more steeply to greater depths.

D e e p structure o f the E u r o p e a n Variscides southerly-derived flysch in the midCarboniferous Sudetic phase, constitutes the Variscan foreland. Large-scale over-thrusting carried the northern margin of the Saxothuringian zone, with basement exposed in the midGerman crystalline rise, northward over the Rhenohercynian zone (Weber 1981), although the exact timing of deformation in these crystalline rocks is not entirely clear (Walliser 1981). The mid-German crystalline rise is considered to correspond in SW England to the Lizard complex since this marks the northern limit of shallow basement (Krebs 1977, fig. 1; Ziegler 1978, fig. 5). Several reflection seismic lines exist in the Rhenohercynian zone, and its structure in the top 6 km is well known. In the northern part of the zone, in north Germany (Fig. 4), a strong reflector at 3 - 4 km depth, dipping slightly to the SSE, has been traced for over 20 km downdip (Meissner et al. 1981; Fig. 5), and can be extrapolated to the outcrop of a prominent fault system marking the boundary between the Rhenohercynian zone and the sub-Variscan foredeep. Reflectors in similar positions have been found 100 km to the SW (Bless et al. 1980; Graulich 1980; Fig. 5) and 175 km to the west (Cl6ment 1963; Fig. 5), where they correspond to the Midi fault system and its extensions (Geukens 1981). These flat reflectors are thought to mark a large thrust along which rocks of the Rhenohercynian zone have been displaced continentward an unknown distance. Above the thrust the reflection character is complex with many discontinuous reflections and diffractions probably caused by relatively small-scale folding and faulting, whereas below the thrust reflectors are more continuous and indicate more regular, flat-lying sedimentary rocks (C16ment 1963; Graulich 1980). Seismic reflection data have also been recorded across the northern margin of the Rhenohercynian zone in southern England where structures are rather poorly defined due to lack of exposure. These data, in contrast to those further east, show that reflectors dip more steeply. In southern England a series of reflections dipping between 20 ~ and 30 ~ to the south are thought to be two major thrust zones subcropping beneath Mesozoic cover that affect upper Carboniferous and earlier rocks (Kenolty et al. 1981; Chadwick et al. 1983; Figs 4 & 5). The northernmost zone, the Variscan front, can be traced down to 7 km depth with no significant dip change, but the more southerly zone can be traced down to between 10 and 11 km depth where it appears to level out (Chadwick et al. 1983).

(g) B

259 HUNSRUCK BORDER FAULT

$AXOTHURINGIAN ZONE

~

RHENOz~I~RCYNIA N

o

N 0

-

20

i ~~ 30

MEtSSNER

ET

AL 19B0

FIG. 6. Line drawing of seismic profile crossing the boundary between the Rhenohercynian and Saxothuringian zones in Germany (here overlain by the SaarNahe trough). Data are migrated and depth-converted The Saxothuringian zone has so far been little studied with seismic profiles. In Germany it is thrust over the Rhenohercynian zone (Weber 1981), but the boundary is largely buried beneath an intramontane basin, the Saar-Nahe trough. One deep (over 12 s) seismic line crosses the north-western margin of this trough, the Hunsruck border fault zone (Meissner et al. 1980; Fig. 6). The fault is a listric zone of disturbance defined by reflected refractions at the surface and at depth by scattered reflections and reflector terminations. It dips steeply to depths of over 30 km, separating a crustal block to the NW with essentially horizontal layering, from the crust to the SE characterized by strong, north-dipping layering. This geometry therefore suggests normal offset at depth, with downthrow to the SE, presumably associated with formation of the Saar-Nahe trough. Thrusts are not recognized; however the profile is short and extensions might enable thrusts to be identified at the surface and thus followed to depth. Some of the most interesting results of seismic profiling in the Saxothuringian zone (if the Lizard complex is taken to be the northern margin of this zone) are the finding of pronounced and continuous south-dipping reflectors, traceable to over 5 s (about 12-15 km depth) with dips of about 30 ~, in the English Channel south of the Cornubian platform (Edwards 1984) and in the region of Lizard Head and Start Point (Day 1982; Leveridge et al., this volume; Fig. 5). These dipping events are thought to be a series of imbricate thrusts since they can be extrapolated to thrusts around the

260

J. A . B r e w e r

ophiolite complex of the Lizard and inferred in the Devonian Gramscatho beds further north (Leveridge et al., this volume). Short, n e a r horizontal reflections occur near the base of some of the seismic data (i.e. at about 15 km depth) which possibly indicate that the thrusts flatten at these depths (Leveridge et al., this volume). Surface rocks onshore are interpreted to be predominantly offshore shelf or offshelf, deposits of Devonian age deposited close to the southern margin of the Old Red Sandstone continent (see review by Anderton et al. 1979), and it seems likely that these rocks comprise much of the thrust belt at depth. Dipping reflector sequences that are somewhat similar to those of SW England also occur on data recorded by Shell UK in the eastern English Channel (Fig. 5). The Rhenohercynian zone is conventionally extrapolated through this area between northern France and southern England although the correlation is very poorly understood (Wallace 1983). These reflectors lie below subhorizontal Cretaceous and later beds and their age and composition is essentially unknown. They also differ somewhat from the dipping reflectors of SW England in that they continue below about 15 km depth without appreciably flattening out, and no horizontal reflectors are seen below them. However they more closely resemble the dipping reflectors to the west than those in the Rhenohercynian zone to the east. To summarize, seismic data recorded in southern England and the English Channel contain reflectors, probably thrusts, that can be traced from the near-surface into the middle crust with moderate dips. Thin-skinned thrust sheets similar to those identified under the Rhenohercynian zone in France and Germany are not seen.

Discussionma model for continental overthrusting? Similar reflecting sequences to those recorded around the edges of the North American continent are found in parts of the European Variscides. Two sorts of reflectors are found: (1) relatively high-level (upper 5 - 8 km) sequences that are subhorizontal or only very gently dipping, and (2) more steeply dipping (about 30 ~) sequences that penetrate to deeper levels (about 15 km) and may flatten at depth. It is not being suggested that the sequences can be directly correlated, but it is possible that similar processes occurred in the Variscan orogeny in

Europe to those deduced for the Appalachian mountain chain. In the Appalachians closure of basins outboard from the North American continent caused imbricated thrust-slices of offshelf rocks and slivers of basement to be stacked against the basement edge. As shortening proceeded material was pushed continentwards to form nappes underlain by subhorizontal thrusts. The amount of actual overthrusting depended on the presence of promontories or embayments in the continental margin as well as on the intensity of shortening. There is thus necessarily no direct correlation between major geological boundaries at the surface and the position of the basement edge of depth. By analogy with the North American data, the European reflectors can possibly be interpreted (Fig. 7) in terms of large subhorizontal overthrusts (roughly corresponding to the Rhenohercynian zone) that southwards and westwards give way to more steeply dipping thrusts. These more steeply dipping structures probably formed at the southern edge of Old Red Sandstone continental basement and consist of thrust slices of (a) offshelf or basinalfacies metasedimentary rocks and (b) basement. Thus it is possible for the 'Variscan front' to be marked by subhorizontal nappe structures in one area (e.g. the Rhenish massif, Meissner et al. 1981), but marked by more steeply dipping thrusts, flattening at about 6 s, in other areas (e.g. in southern England, Kenolty et al. 1981; Chadwick et al. 1983) because of variations in the amount of overthrusting and the morphology of the continental margin. No single structural style should be considered diagnostic of the 'Variscan front'. In Germany the continental margin must lie somewhere south of the existing reflection profiles that show nappe structures, and a continuous traverse extending Meissner et at.'s (1981) profile to the south is required to locate the edge. This should be marked by sets of reflectors dipping at about 30 ~ to the south. If the zones of dipping reflectors do indeed occur at the southern edge of the Old Red Sandstone continent, few clues exist to the nature of the basement south of this edge. Evidence from Germany suggests that the Variscan orogeny is entirely ensialic (Weber 1981), but the Lizard ophiolite lies in the region of the southerly-dipping reflectors which suggests that crust of more oceanic affinities might once have existed under the English Channel area. Gravity and magnetic data in the southern Appalachians is consistent with oceanic, or thinned continental, crust lying outboard of the easterly-dipping reflectors there (Cook &

261

D e e p structure o f the E u r o p e a n Variscides A L L O C H T H O N CAN COMPRISE: 1) DEEP WATER EQUIVALENTS OF SHELF SEQUENCE 2) BASEMENT AND COVER ORIGINALLY OUTBOARD FROM CONTINENTAL EDGE 3) BASEMENT STRIPPED OFF CONTINENTAL EDGE

|

50

KMS

I

0 (~+,+++.++,++++++++++++

210

+ + + + §

20

+ * * * § 'FORELAND

"

.....

**********

......... ***+* OVERRIDEN CONTINENTAL

....

**

BASEMENT

SHELF

SEQUENCE

EDGE

ZONE OF DIPPING REFLECTORS. SEEN IN: SOUTHERN A P P A L A C H I A N S OUACHITAS SOUTH ENGLAND ENGLISH CHANNEL

ZONE OF SUBHORIZONTAL REFLECTORS. SEEN IN: SOUTHERN APPALACHIANS RHENOHERCYNIAN ZONES OF GERMANY & FRANCE

FIG. 7. Schematic representation of possible connection, highly simplified, between subhorizontal reflectors under an allochthon and more steeply dipping reflectors at the edge of the overridden continental margin. Bold lines indicate main reflecting surfaces and lighter lines are subsidiary reflecting surfaces. Note that subhorizontal reflecting surfaces are also found in the Canadian Rockies (e.g. Bally et al. 1966) and more steeply dipping reflectors found in COCORP data in the northern Appalachians (Ando et al. 1983) and in the Scottish Caledonides (Brewer & Smythe 1983).

Oliver 1981) and similarly in the Ouachitas (Lillie et al. 1983). Palaeomagnetic data, which at the moment are equivocal (Irving 1977) are the only way of determining whether a wide ocean existed south of the Old Red Sandstone continent. The two seismic characters described in these late Palaeozoic orogenic belts do occur elsewhere. In the New England Appalachians 30-40~ reflectors can be traced from close to the surface, where they coincide in part with the westward-transported early Palaeozoic shelf-offshelf transition, t o 18-21 km depth (Ando et al. 1983). Subhorizontal reflections are also observed from the top of shelf carbonates that underlie offshelf rocks of the Taconic nappes. These structures were formed in Taconic or Acadian times. In north Scotland 20-30~ thrusts occur along the offshore extension of the Caledonian Moine thrust zone, whereas onshore the Moine thrust is thought to be a subhorizontal structure (Watson & Dunning 1979; Brewer & Smythe 1984). In the Canadian Rockies subhorizontal nappes can be traced westward 100-200 km (Bally et al. 1966) but as yet the 'root zone' which should be defined by about 30 ~ dipping reflectors at the basement edge has not yet been crossed by reflection profiles.

Conclusions Despite apparent differences in the surface geology it appears that crustal seismic reflector geometries are somewhat similar along the northern margin of the North American and the European Variscides. Subhorizontal reflectors (thrusts) under the exterior zones of the Appalachians pass into more steeply dipping reflectors (also thought to be thrusts) under the more interior zones, which are thought to correspond approximately to the overthrust continental margins. The spatial relationships between these two reflector sequences are not precisely known in the European Variscides but it appears that they lie in equivalent positions with respect to the margins of the Old Red Sandstone continent. The data discussed here will soon be supplemented by extensive deep reflection profiles recorded in the English Channel by the British Institutions Reflection Profiling Syndicate, and in the Paris basin by the French Programme ECORS. These two new datasets should resolve the problem of the relationship between the two reflector sequences in those areas. However we already know that at depth the character of the Variscan front is quite variable along strike, and these reflector sequences must be accounted for in all models for the formation of the Variscan orogenic belt.

262

J. A. Brewer

ACKNOWLEDGMENTS: Shell U.K. kindly allowed publication of the line drawing of their data in the eastern English Channel. The C O C O R P data were collected by Geosource, Inc. J. Edwards, G. Day and A. Chadwick kindly sent preprints of forthcoming papers and gave permission for use of diagrams in these pap-

ers. Many of the ideas expressed in this paper were developed by, and through discussion with, C O C O R P personnel at Cornell University. The author acknowledges support by a N E R C research fellowship held at Cambridge University. Cambridge contribution to Earth Sciences No. 400.

References ANDERTON, R., BRIDGES, P. H., LEEDER, M. R. & SELLWOOD,B. W. 1979. A Dynamic Stratigraphy o f the British Isles. Allen & Unwin, London. 301 pp. ANDO, C. J., COOK, F. A., OLIVER, J. E., BROWN, L. D. & KAUFMAN, S. 1983. Crustal geometry of the Appalachian orogen from seismic reflection studies. In: HATCHER, R. D. & ZIETZ, I. (eds) Tecton&s and Geophysics o f Mountain Chains. Mem. geol. Soc. Am. 158, 83-101. BADHAM, J. P. N. 1982. Strike-slip orogens--an explanation for the Hercynides. J. geol. Soc. London, 139, 495-506. BALLY, A. W., GORDY, P. L. & STEWART, G. A. 1966. Structure seismic data and orogenic evolution of southern Canadian Rocky Mountains. Bull. Can. Petrol. Geol. 14, 337-81. BAYLEY, R. W. & MUEHLBERGER, W. R. 1968. Basement Rock Map o f the United States 1:2,500,000. U.S. Geological Survey, Washington, D.C. BLESS, M. J. M., BOUCKAERT, J. & PAPROTH, E. 1980. Meded. Rijks geol. Dienst. 32, 3-13. BREWER, J. A. & SMYTHE, D. K. 1984. MOIST and the continuity of crustal reflector geometry along the Caledonian-Appalachian orogen. J. geol. Soc. London, 141, 105-20. CHADWICK, R. A., KENOLTY, N. & WHITTAKER, A. 1983. Crustal structure beneath southern England from deep seismic reflection profiles. J. geol. Soc. London, 140, 893-911. CLEMENT, J. 1963. Resultats preliminaire des campagnes geophysiques de reconnaissance dans les permis de recherches "Arras et Aresnes" de l'Association Shell Francaise -P.C.R.B.-SAFREP. Objectifs du forage profond JeumontMarpent No. 1. Annls Soc. geoL N. 83, 237-41. COOK, F. A., ALBAUGH, D. S., BROWN, L. D., KAUFMAN, J., OLIVER, J. E. & HATCHER, R. D. 1979. Thin-skinned tectonics in the crystalline southern Appalachians: C O C O R P profiling of the Blue Ridge and Piedmont. Geology, 7, 563-7. , BROWN, L. D., KAUFMAN, S., OLIVER, J. E . & PETERSEN, T. A. 1981. C O C O R P seismic profiling of the Appalachian orogen beneath the Coastal Plain of Georgia. Bull. geol. Soc. Am. 92, 738-48. & OLIVER,J. E. 1981. The Late PrecambrianEarly Palaeozoic continental edge in the Appalachian orogen. Am. J. Sci. 281, 993-1008. DAY, G. A. 1982. Gravity and seismic interpretation in the Western English Channel (abstract). Geophys. J. R. astr. Soc. 69, 278.

GARDINER, P. R. R. & SHERIDAN, D. J. R. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan Front. J. struct, geol. 3, 317-31. GEUKENS, F. 1981. Cross-sections through the Belgian Variscan Massif. Geologie Mijnb. 60, 45-48. GRAULICH, J. M. 1980. Les structures profondes en Famenne, Constituent-elles des reservoirs de gaz? Revue gdn. Gaz. 4, 165-72. HARRIS, L. D. & BAYER, K. 1979. Sequential development of the Appalachian orogen above a master decollement - - a hypothesis. Geology, 7, 568-72. ~, HARRIS, A. S., DE WITI', W. & BAYER, K. C. 1981. Evaluation of southern eastern Overthrust Belt beneath Blue Ridge--Piedmont thrust. Bull. Am. Ass. Petrol. Geol. 65, 2497-505. HATCHER, R. D. (Jr) 1978. Tectonics of western Piedmont and Blue Ridge, southern Appalachians: review and speculation. Am. J. Sci. 278, 276-304. -1981. Thrusts and nappes in the North American Appalachian orogen. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics., Spec. Publ. geol. Soc. Lond. 8, 491-500. Blackwell Scientific Publications, Oxford. IRVING, E. A. 1977. Drift of the major continental blocks in the Devonian. Nature, 270, 304-9. IVERSON, W. P. & SMITHSON, S. B. 1982. Master Decollement root zone beneath the southern Appalachians and crustal balance. Geology, 10, 241-5. KENOLTY, N., CHADWICK, R. A., BLUNDELL, D. J. & BACON, M. 1981. Deep seismic reflection survey across the Variscan Front of southern England. Nature, 293, 451-3. K~U~BS, W. 1977. The tectonic evolution of Variscan Meso. Europa. In: AGER, D. V. & BROOKES, M. (eds) Europe from Crust to Core, 119-39. Wiley, New York. LE PICHON, X., SIBUET, J. C. & FRANCHETEAU, J. 1977. The fit of the continents around the North Atlantic Ocean. Tectonophys. 38, 169-209. LILLIE, R. J., NELSON, K. D., DE VOOGD, B., BREWER, J. A., OLIVER, J. E., BROWN, L. D., KAUFMAN, S. & VIELE, G. W. 1983. Crustal structure of the Ouachita Mountains, Arkansas: a model based on the integration of C O C O R P reflection profiles and regional geophysical data. Bull. Am. Ass. Petrol. Geol. 67, 907-31. MATTHEWS, S. C. 1978. Caledonian connexions of Variscan tectonism. Z. dr. geol. Ges. 129, 423-6.

D e e p structure o f the E u r o p e a n Variscides MEISSNER, R., BARTELSEN, H. & MURAWSKI, H. 1980. Seismic reflection and refraction studies for investigating fault zones along the geotraverse Rhenoherzynikum. Tectonophys. 64, 59-84. &~ 1981. Thin-skinned tectonics in the northern Rhenish Massif, Germany. Nature, 290, 399-401. NELSON, K. D., LILLIE, R. J., DE VOOGD, B., BREWER, J. A., OLIVER, J. E., KAUFMAN, S. & BROWN, L. D. 1982. COCORP seismic reflection profiling in the Ouachita mountains of western Arkansas: Geometry and geologic interpretation. Tectonics, 1, 413-30. NICHOLAS, R. L. & ROZENDAL, R. A. 1975. Subsurface positive elements within Ouachita foldbelt in Texas and their relationship to Paleozoic cratonic margin. Bull. Am. Ass. Petrol. Geol. 59, 193-216. RODGERS, J. 1968. The eastern edge of the North American continent during the Cambrian and Early Ordovician. In: ZEN, E-AN, WHITE, W., HADLEY, J. B. & THOMPSON, J. B. (Jr) (eds) -

-

Studies of Appalachian Geology: Northern and Maritime, 141-50. Wiley, New York.

263

STILLE, H. 1951. Das mitteleuropaische variscische Grundegebinge in Bilde des gesamteuropaischen. Beih. geol. Jb. 2, 1-138. WALLACE, P. 1983. The subsurface Variscides of southern England and their continuation into continental Europe. In: HANCOCK, P. L. (ed.)

The Variscan Fold Belt in the British Isles, 198-208. Hilger, Bristol. WALLISER, O. H. 1981. The geosynclinal development of the Rheinische Schiefergebirge. Geologic Mijnb. 60, 89-96. WATSON, J. • DUNNING, F. W. 1979. Basementcover relations in the British Caledonides. In: HARRIS, A. W., HOLLAND,C. H. & LEAKE,B. E. (eds) The Caledonides of the British Isles: reviewed. Spec. Publ. geol. Soc. Lond. 8, 67-92. Blackwell Scientific Publications, Oxford. WEBER, K. 1981. The structural development of the Rheinische Schiefergebirge. Geologie Mi]nb. 60, 149-60. ZIEGLER, P. A. 1978. North-western Europe: tectonics and basin development. Geologic Mijnb. 57, 589-626. 1981. Geological Atlas of Western and central Europe. Shell int. Petrol. Maatsch.

J. A. BREWER, Department of Earth Sciences, Bullard Laboratories, Madingley Road, Cambridge, England.

Subject Index Many local place names and relatively small structures, if mentioned only in one paper, have been omitted from this index or have been referred to under a more general entry. References to the Devonian and Carboniferous are only indexed where more detailed stratigraphical subdivision is made in the text. Aachan, 23-9 thrust, 23, 27 Acadian, 15-6, 65, 69, 197, 205-6, 209-11, 233, 245,256, 261 9 Accretionary prism, 158 Acritarchs, 49 Actinolite, 247 Aigurande nappe, 51, 57 Alabama, 199, 202, 207-8 Albite, 52, 54, 66-7 Algeria, 206 Alleghenian, 78, 181, 197-212, 233-41,245, 247, 254 Allegheny front, 208-10 platform, 202-3 synclinorium, 203 Allochthon, 208-9 Alluvial fan, 233,235 Alpine belt, 3, 5, 14, 29 Alston block, 139-46, 157 Altenburen fault, 75-6 Alto Alentejo, 4, Amphibolite, 219 Anatexis, 4-5, 7-8, 47, 50-1, 67-8 Andalusite, 52-6, 66, 247 Andrews-Sleep cell, 15-18 Andurn fault, 116-7 Anguille Group, 204, 221 Anti-Atlas orogeny, 182 Apatite, 67 Appalachian axis, 181 basin, 202-3 mountains, 24, 29-30, 73, 77-80, 197-210, 229, 253-6, 260-1 platform, 200-2 Ardmore syncline, 168 Arkansas, 210, 256-7 Arkoma basin, 210, 256-7 Armorican massif, 3, 161 Arnsbergian, 134 Arundian, 133, 135 Asbian, 133, 136 Askrigg block, 131-4, 157 Asthenosphere, 15 Asturic, 257 Austria, 5-6 Avalonian belt, 79-80,206-7,210-1,235,239,245, 248 Back-arc basin, 14, 50, 99-100, 158 Bajocian, 227 Bala fault, 126-7 Barachois Group, 221

Barrandian, 79 Basalt, 5, 34, 37, 106, 200, 219, 235-6 alkali, 35, 158-9 tholeiite, 7, 29, 35, 49-50, 158 Bavaria, 11, 99 Bavarian facies, 35-44, 74-5 Beara, 151-5 Belgium, 81, 99 Belle Isle, 207, 220, 223, 228, 240 Benton fault, 92 Benton uplift, 210, 256-7 Berga anticline, 36, 41 Biotite, 52, 54-6, 66-8, 237, 239, 247-8 BIRPS, 261 Birrimian, 128 Biscay, Bay of, 76 Blackdene mine, 141-2 Black Forest, 17 Black Warrier basin, 119, 202 Bloody Bluff fault, 205, 234, 239, 246 Blue Ridge, 79, 207-8, 254-6 Bodmin Moor granite, 94-5, 113 Bohemian massif, 3, 8, 73 Bolt Tail, 113 Bonaventure Formation, 201 Boston basin, 204, 236 Boudinage, 40, 47, 49, 51, 153, 237-8, 246--8 Bouguer anomaly, 119-20, 150, 156, 180 Bovey basin, 96 Bow and arrow rule, 96 Braided stream, 233 Branch line, 95-6 Bretonic, 161,258 Brevard fault, 79, 208 Brigantian, 135-6 Brioverian, 104, 110 Bristol Channel, 76, 81, 98, 177, 181,224 Brittany, 69, 73-4, 99-100, 161-2 Brittle-ductile transition, 153 Burtreeford disturbance, 141-2 Buttevant, 168, 170-2 Cabot fault, 201,204, 207 Cadomian, 3, 17, 76, 80, 100, 128, 161 Caledonian, 3-7, 17, 76-7, 80, 126, 180-1, 185-6, 191, 197, 253,261 basement, 156, 224 cleavage, 143, 190 granites, 139 Cambrian, 5, 25, 34, 47, 69, 203, 219 Canadian shield, 203 Rocky Mountains, 167, 173,261 Cannes de Roches Formation, 201 Cannington thrust, 181 Cap Rouge peninsula, 223 Cape Rouge Formation, 221 Cardigan Bay, 180-1 Carrick nappe, 105, 108-10 thrust, 104-9 Celtic Sea, 72, 119, 123, 125-7, 177, 180, 224-7 Cephalopod, 35

266

Subject Index

Chadian, 133, 135-7, 189 Chambon thrust, 63-9 Channel Islands, 71, 73 Charlotte belt, 254-5 Chateaulin basin, 74, 161-2 Ch6niers, 67-8 Chert, 35, 103, 210 Chester, 248 Cheviot block, 144 Chlorite, 52, 54-6, 67, 152, 168, 247-8 Chloritoid, 52, 54, 237, 239 Church Stretton fault, 126-7 Cincinnatti dome, 202 Clare, 189-90 Cleavage, 40-2, 68, 76, 98, 152-3, 167, 190 crenulation, 51-2 pressure solution, 95 slaty, 42, 93 Clinton-Newbury fault, 205-6, 246 Coal, 202, 204 COCORP, 23-4, 73,208,210, 253-61 Cobequid-Chedabucto fault, 181,203, 225, 234-5, 240, 249 Codroy Group, 204, 221 Collision, 99, 110, 163, 202, 211-2, 241 Comeragh Mountains, 150, 168, 170 Conch peninsula, 220, 223 Condroz thrust, 23-4 Connecticut, 205, 235, 245-8 Conodonts, 103 Coral, 35 Cordoba-Abrantes shear zone, 4 Cordierite, 52, 54-6, 67 Cork, 150-1,154-6, 167-74, 179 -Kenmare line, 150, 173 Syncline, 168, 171-3 Cornubian batholith, 121-3, 160 basin, 224, 227 platform, 119-23 Cornwall, 71, 73, 93-98,103-11, 119, 128,157-61, 205,228 Courtmacsherry Formation, 168 Crack-seal, 152 Craven basin, 131-7 fault system, 132-6 Cretaceous, 119, 121,228 Creuse River, 65-8 Crevant massif, 66 Crouse Harbour Formation, 221 Crozant massif, 66 granite, 66 Crustal thinning, 161 Culmination, 97 Culm Basin, 159-60 Synclinorium, 93, 98, 160 Czechoslovakia, 6 Damara orogen, 80 Dartmoor, 93, 160 granite, 94, 113, 115 Dartmouth, 93, 95 antiform, 114, 116-7 Beds, 114, 117, 159 Decazeville, 57 D6collement, 121,125-8,156, 167,169, 173,208-9

Devils Bit Mountains, 187 Devil's River uplift, 256 Devon, 89, 93-8, 113-7, 159-60, 181 Dextral Shear, 76-7, 79, 156, 160-2, 177-82, 185, 221,229--see also Wrench fault Diapir, 33, 228 Dinantian, 25, 131-6, 139, 156 Dingle-Dungarvan line, 150, 155, 167, 172, 177-81 Diorite, 9-10, 49, 236, 239 Dodman Point, 106 nappe, 109-10 thrust, 107-10 DSDP, 227 Dungarvan, 177-81 syncline, 168, 170, 172-3 Durham coalfield, 140 Dyke, 157, 189 ECORS, 261 Eguzon unit, 63-9 Eifelian, 103 Elbe, 8 Emsian, 25 England North, 131-7, 139-46, 157, 189, 224 Southwest, 71, 73, 93-8, 103-11,113-7, 119-23, 125-8, 157-62, 181,205, 219, 223-6, 228, 257-60 English Channel, 73, 98-9, 106-8, 128, 259 Entropy, 142-3 Epidote, 247 Erbendorf line, 10, 37, 44, 75 Eulengebirge, 33 Evaporite, 233 Exmoor thrust, 181 Extension direction, 95 Facing direction, 90 Faille du Midi, 23-4, 258-9 Fall River fault, 246, 248 Famennian, 25, 37-8 Faulting, 154 en echelon, 132, 135-7, 189 extension, 98 listric, 98, 116 normal, 95, 98, 116, 141,144-5, 154, 187, 189, 193,202, 239, 257 oblique, 189 reverse, 155, 178--see also Thrust strike-slip--see Wrench tear, 95-6, 100, 208, 246--see also Wrench thrust--see Thrust Fergus shear zone, 185-93 Fichtelgebirge, 33, 35-6, 41, 44, 75 Falmouth, 108 Florida, 79 Fluviatile, 233,236 Flysch, 4, 14, 89, 98, 103-4, 157-8, 203, 210, 256, 258 Folding, 37, 40-1, 151-2, 178, 190-2, 221,236 back, 95, 116 box, 170 chevron, 51, 93, 135 drag, 38 en echelon, 132-7, 193 intrafolial, 246

Subject Index isoclinal, 6, 51 oblique, 108 pericline, 135 sheath, 96, 108 step, 152 Foredeep, 98 Foug~res unit, 63-4, 67-8 Foynes, 185, 189-90 Fracture trace analysis, 142-3 France, 6-8, 47-58, 63-9, 73-6, 99-100, 161-2, 260 Frankenberg, 33 Frankenwald, 35-6, 39, 42 Frasnian, 104 Fredericton-Norumbega fault, 201-2, 249 Fundy, 203 Gabbro, 9, 34, 49, 208,219 Gaily Head, 171 Galty Mountains, 178-9, 187 Galway Bay, 173 granite, 190 Gargilesse unit, 63-9 thrust, 63-4, 66 Garnet, 6, 49, 52-6, 67-8, 237, 239, 247-8 Gaspe, 209 Gedinnian, 25 Georgia, 79, 207, 211 Geothermal gradient, 16, 30, 160 German Crystalline Rise, 259 Germany, 4-6, 9-14, 23-9, 33-44, 75-6, 81, 99, 257, 259-60 Gf6hler orthogneiss, 6 Gibbs fracture zone, 180 Givetian, 98, 104 Glandore high, 171, 174 Gondwanaland, 15, 17, 245, 249-50 Gortdrum fault, 192 Gower, 96-7 Graben, 7, 76, 135, 150, 197,201-3,206,228,233, 236-7, 239 Gramscatho Group, 97, 103-6, 108-10, 260 Basin, 157-61 Grand Banks, 219, 228-9 Granite, 9, 30, 42, 47, 49, 66, 208, 219, 235, 239 S-type, 160, 239 Granodiorite, 9-10 Gravity data, 119-23, 132-3, 156, 180-1,187, 190-1,211,224, 227, 256-7, 260 Great Glen fault, 206, 249 Great Valley, 207 Green Mountains, 256 Grenville, 208, 254, 256 Gressen nappe, 11 Groais Island, 221 Gulf of Mexico, 211 Gulf of St Lawrence, 200-1 Hadrynian, 219, 223 Haig Fras granite, 119-21, 123, 127-8 Hare Bay 219-21,223 Harvey-Hopewell fault, 233-4 Harz Mountains, 11, 16, 73 Hartzburgite, 50 Haut Allier nappe, 51-3, 55, 57-8

267

Haut Limousin nappe, 51, 57-8 Heat flow, 14, 18 Helikian, 219 Hercynite, 56 Hesperian massif, 5, 7-8 Hesse, 11 Holes Venn, 23-5 Honey Hi!l fault, 206, 234-5,239, 245-50 Hornblende, 6, 9, 67, 247-8 Horst, 135,203, 237 Horton, 203 Hunsruck, 28-30, 73,259 Iapetus ocean, 3, 15, 76-7, 79-81 suture, 80, 157, 185, 189 Iberian peninsula, 72, 74, 76 Ibero-Armorican arc, 73,219 Icartian, 100, 110, 128 Ignimbrite, 50, 200 lie de Groix, 161-2 Illinois basin, 202 Illite crystallinity, 42-3 Imbrication, 246 Ireland SW, 81, 96, 126, 149-57, 167-74, 177-81, 224-7 Central, 185-93 Irish Sea, 126, 179-80, 226-7 land mass, 180 Isograd, 52, 56, 209, 237 Isopach, 150 Isotopic age, 30, 34, 205, 239 4~ 248 K-Ar, 8, 10, 14, 68, 97, 106, 158 Rb-Sr, 4, 6, 8, 10-1, 49, 57, 66 87Sr/86Sr initial ratios, 3, 66 U-Pb, 3-5 Iveragh, 154 Jacksboro fault, 203 Jennycliff slates, 116-7 Johnston thrust, 91-3, 97 Joints, 141 en echelon, 141 extension, 141, 145 shear, 141, 144 Jurassic, 119 K-feldspar, 7, 55, 66-8, 247 Kansas-Nebraska basin, 202 Kelvin fracture zone, 78 Kenmere, 150 Kentucky, 203 Keratophyre, 34-5, 50 Kerry, 154, 156 Kildare, 179 Killarney, 152, 172-3, 179 Kiltorcan Formation, 168-9 Kink band, 41, 154, 236-8 Kinsale, 170 Formation, 167-9, 172 Klippe, 117 Knockmealdown Mountains, 170, 173 Kossmat's zones, 71-6 Kyanite, 5, 52-6, 67-8, 247

268

Subject Index

Labrador sea/shelf, 219-23, 228-9 Lacustrine, 233, 236 Lake Char fault, 234-5, 239, 246-7 La Marche shear zone, 63-5, 69 Land's End, 119-21,226 Laurasia (Laurussia), 17, 248-50 Laval basin, 74 Lead isotopes, 7 Leptyno-amphibolitic group, 47-50, 57, 67 Ligerian suture, 14, 18 Limerick, 189 Limousin, 55, 57, 69 Lithosphere, 44, 150 thinning, 150 Lizard complex, 71, 74, 93, 95-6, 99, 103, 108, 157-8, 162, 257, 259-60 nappe, 110-11 ophiolite, 93, 98-9, 260 thrust, 107-10, 160 Llanoria, 210-2 Lleyn peninsula, 180 Logan line, 210 London-Brabant massif, 26, 30 Louisiana, 197 Lyonnais, 55, 57 Lyons, 8 Macroom, 171 Madura body, 121 Magnetic anomaly, 177, 179-80, 187, 190-1,207, 221,224, 227, 229, 260 Maine, 199, 245 Mallow, 167, 170-3, 177, 179 Malvern, 76, 127 Malvernian, 126 Mantle, 28 Marathon belt, 198 Margeride granite, 52-3, 56 Marginal sea basin--see Back-arc Basin Maritime Canada, 198,205, 211,233-41 Maryland, 79, 208 Marzan-Trap Mountains, 210 Massachusetts, 204, 233-6, 245 Massif Central, 4, 6, 8, 47-58, 63-9, 76, 99 Mauretanides, 72 McKenzie model, 99, 150-1, 156, 161 Meadfoot Group, 106, 114, 117 Meguma, 240 M61ange, 38-9, 103-4, 110, 219 Menai Straits fault, 126 Mendips, 98 Meneage breccia, 103-6, 108-9 Merrimac Group, 245 Metamorphism Abukuma-type, 17 amphibolite facies, 5-8, 33, 43, 49, 52, 67,246-7 Barrovian, 55, 205-6, 239 eclogite facies, 5, 7-8, 34-5, 49, 67-9, 74, 99 granulite facies, 5-8, 16, 47, 50, 55, 74, 99 greenschist facies, 10, 33, 35, 43, 89, 248 low pressure/high temperature, 18, 30, 72, 77, 161 low temperature/high pressure, 161 retrograde, 56, 247-8 Michelbach thrust, 10, 16 Michigan basin, 202

Mid-German high (crystalline rise), 8 - i 1, 29, 74-5 Midi fault--see Faille du Midi Midland Valley of Scotland, 149, 157 Migmatite, 5-6, 8, 67 Migmatitic Unit, 63-4, 66-9 thrust, 63-4, 66 Millstone Grit, 131 Minas geofracture, 240 Mineralization, 141, 145-6, 154, 185, 187-9 Mississippi, 197, 199, 202 embayment, 197-8 Moho, 15, 29, 75, 121-2, 151, 156, 160-1 Moine thrust, 261 Molasse, 202, 211 Moldanubian zone, 10, 16, 18, 28-9, 37, 73, 75, 78, 99 granite, 6 Montagne Noire, 47 Monzonite, 49, 236 Monts de Lyonnais, 6 Morbihan, 8 Morocco, 79, 181,206 Mfinchberg, 5, 9, 33-44, 74-5 Munster Basin, 150-6, 172-4, 178,226-7 Muscovite, 52, 54-6, 66-8, 247-8 Mylonite, 4-6, 10-11, 38-9, 47, 50-1, 99, 162,205, 239, 245, 247-8 Mylor slate, 97, 104-6, 109-10 Namurian, 25, 91-2, 98,131-3,135-6, 139, 156-7, 161,185,189-90, 200,202-3,224-5,228,233 Nappe, 7, 11, 14, 17, 33, 37, 42-4, 50-4, 63-4, 75, 104-11,260 Narragansett basin, 78, 204-5, 234-6, 239-41 Pier granite, 205-6, 239 Navan, 156, 185, 189 Nenagh block, 185, 191 New Brunswick, 198-207, 210-1,233-41 syncline, 201-2 New England, 181,204-5, 210, 229, 233-41, 245-50 New York, 199-200, 202-3,208 Newfoundland, 78, 181,199, 204,206, 209, 228-9, 233 Newquay, 114 Norfolk basin, 201,236, 239 Normandy, 73 North Carolina, 197, 211 North Curlew Mountains, 187 North Devon basin, 160 North Pennine ore field, 139-46 Northumberland basin, 131,144 Nova Scotia, 198-200, 203-6, 211,229, 234-5 Obduction, 158 Odenwald, 9-10, 14, 17 Old Red Sandstone, 89, 91,168, 185, 187, 257, 260-1 Oligocene, 119 Olistolith, 89, 104, 133 Olistostrome, 89, 97, 104 Olivine, 15 Ophiolite, 12, 50, 93, 98,103,110, 158,127-8,211, 260

Subject Index Ordovician, 4-7, 11, 16, 25, 33-5, 38, 40-1, 49, 74, 79, 89, 103-4, 185, 197,206, 219, 221 Orphan Knoll, 225,227-8 Orsennes massif, 66 Ouachita belt, 197-8,200, 208-12, 255-7, 261 Ougarta orogeny, 182 Ozark dome, 202 Palaeocene, 119 Palaeomagfietism, 97, 163,206, 209, 249-50 Pan-African, 80 Pangea, 248 Pattaconk Brook fault, 246, 248 Pembrokeshire, 89-93, 96-8, 177, 179, 181-2 Pennine block, 131 fault system, 139-40 Pennsylvania, 199-200, 203,207-8 Pentleian, 134-7 Peridotite, 15, 49 Permian, 11, 16, 28, 77, 80, 98, 106-8, 116, 119, 131,144-5, 180-1,185, 201-2, 204, 235, 240-1,245, 248, 250 Perranporth-Pentewan line, 106 Phengite, 12 Picton Group, 204-5 Piedmont, 79, 207-8, 211,254-6 Pin Hill, 204-5 Pine Mountain, 203 Pinite, 54 Plagioclase, 52, 54, 247 Plateau d'Aigurande, 63-9 Plymouth, 93, 113-5, 160 limestone, 114-5 Poland, 33 Polzeath, 95-6, 98, 159 Porcupine Bank, 180, 219, 224-8 Seabright, 224-5, 227 Porphyroblast, 56, 247 Porphyroclast, 247-8 Portugal, 11, 74 Prasinite, 34 Prague syncline, 79 Precambrian, 33-4, 69, 75, 80, 92, 167, 207, 229, 236, 245 Pressure solution, 93, 95 Pressure fringes/shadows, 153, 248 Prince Edward Island, 198,201 Proterozoic, 80 Pseudotachylite, 51 Pull-apart basin, 162, 206, 233, 236, 239 Pyriclastite, 5 Pyroxene, 15 Quebec, 199, 209 Quin shear zone, 188, 191-3 Radiometric ages--see Isotopic ages Randschiefer, 35, 40-1 Red beds, 202, 204, 235-6 Rheic ocean, 72 Rheinisches Schiefergebirge, 11-4 Rhenish Massif, 27-9, 99, 260 Rhenohercynian zone, 9, 11-4, 16, 18, 73-5, 78, 99, 110, 125, 128, 177, 181, 228, 257-60 Rhine embayment, 24, 26

269

Rhode Island, 198, 203-5,233-41 Rhyolite, 236 Ribband Group, 178 Ribblesdale fold belt, 131-7, 157, 189 Riebeckite, 11 Rifting, 7, 14, 17, 30 Riedel shears, 141-2, 145, 189, 191,240 Ritec thrust, 91-3 Rogers Pond fault, 246, 248 Rome trough, 203 Roseland, 103-6, 110 breccia, 104-6, 108-10 volcanics, 104 Rossendale block, 131 Rosslare, 180 Rotgneiss, 4, 7, 9 Rouergue, 51-2 Rusey fault, 95 Rutile, 52, 55 Saar, 8-9, 11, 17, 74 Saar-Nahe trough, 17, 28-9, 259 Salt, 200, 204 Saxony, 5, 8, 75 Saxothuringian zone, 4, 7-9, 14, 16-18, 33-44, 73-80, 99, 257-9 Schwarzwald, 4, 7-8, 10 Scilly, Isles of, 119-21 Scituate basin, 204-5,236 Seismic refraction, 108, 116-7, 121,160, 226, 228 reflection, 23-9, 106-8, 121, 123,219, 221-8, 253-61 Sericite, 54, 56 Serpentinite, 44, 49, 67, 73,210 Settle, 132-3 Shannon River, 185, 190, 192 Shear zone, 10-11, 64-5, 93, 96-7, 153, 161-2, 179-80, 185-93, 229, 246-8 Siegenian, 25 Silesian basin, 226 Sillimanite, 49, 52, 54-7, 67-8,247-8 Silurian, 4-8, 33, 35, 39, 69, 77, 80, 89, 92, 104, 157, 173, 185, 208-9, 221 Silvermines, 156, 185, 187-9 -Navan fault, 149, 157, 185, 187-8, 191 Simple shear, 178, 193 Sioule nappe, 51, 55, 57 Skarn, 49 Skipton anticline, 134-5 Rock fault, 134-5 Slickensides, 38, 95-6, 141,153, 168, 189 Slieve Aughty, 186-8, 190, 192-3 SOQUIP, 209 South Armorican Shear Zone, 64, 149, 161-2 South Atlas fault, 72, 206, 245, 249 South Carolina, 211 South Western Approaches basin, 119, 123 Spain, 11 Spessart, 9-10 Sphene, 67 Spilite, 29, 37, 106 St Anthony basin, 204, 228 St John's basin, 234-6 Start complex, 116-7, 259 Staddon Grits, 115-7

270

Subject I n d e x

Staurolite, 5, 52-6, 67 Stavelot, 12 Stephanan, 139, 144-5, 162, 202, 211,236 Strain, 5, 95-7, 153, 155 oblate, 91, 96, 160 shear, 190, 193 Stretching lineation, 248 Strike-slip orogen, 68, 89, 249 fault--see Wrench fault Strokestown, 190 Stromatoporoid, 35 Subduction, 12-18, 29-30, 73, 75, 158, 163 Subfluence, 14-15, 75 Sudetic, 206, 259 Sucessor basin, 203

Transport direction, 38, 51,245 Transpression, 99, 135-7, 154-7, 160-3, 185, 191, 193 Trevone basin, 159-60 Triassic, 106-8, 119, 131, 180-1, 197, 236, 248 Trondjheimite, 47, 49-51, 57 Truro antiform, 104 Tulle, 52 Turbidite, 11, 35, 103-4, 256 Tynagh, 156, 191 fault, 187 Tynehead granite, 140, 143

Taconic, 78, 197, 209, 229, 256, 261 Tatnic fault, 245-8 Hill Formation, 245-7 Taunus, 30 Taylors Brook fault, 219-20, 223,229 Tectonic slide, 209 Teesdale dome, 140, 145 Tennessee, 199, 203, 207 Tertiary, 96, 123 Texas, 256 Theic ocean, 72, 79 Thin-skinned tectonics, 89-90, 99, 125-8, 155-6--see also Thrust Thick-skinned tectonics, 155-6 Thuringian facies, 35-44, 74 Thrust, 12-14, 17, 26, 30, 36-43, 49-52, 56-7, 63-9, 73-4, 80, 121, 181,203,205, 208-10, 223, 236-7, 240, 256-61 back, 90, 92, 95, 98, 160 blind, 117, 125, 156 culmination, 97 fault, 144, 154 footwall ramp, 108-10 hangingwall ramp, 109-10, 173 imbricate, 91-2, 117 lateral ramp, 92, 99, 108-10, 168, 173 lateral tip, 96-7, 99-100 oblique, 229 pin-line, 169 piggy-back, 91-3, 98 ramp, 26, 93, 117, 172, 208 sole, 156, 172-3 tectonics, 89-90, 95-7, 103-11,113-7, 125-8, 161-2, 167-74 tip line, 125 Tintagel, 95, 98, 159 Toe Head Formation, 169 Torbay, 93, 98, 159 thrust, 114-5 Tornquist's line, 72 Torquay, 96, 113, 117 Tourmaline, 67 Tournasian, 38, 133,135,185,187-8,200, 203,235 Transform fault, 210

Valley and Ridge, 207-8, 254-6 Variscan front, 24, 26, 30, 72, 78, 81, 96, 121, 125-8, 131,150, 154, 160, 162, 169, 172-3, 177-82, 203, 208, 219, 223-9, 258-9, 258-61 Veins, 140-3, 152-4, 186 en echelon, 142, 190 extension, 153-4, 189-93 Vend6e, 8 Venn, 12, 23-6 Veryan limestone, 103 Virginia, 197-9, 203, 208 Visean, 50, 110, 133, 156, 161, 186, 200, 202, 204, 206, 228, 233 Voges, 8, 10

Unconformity, 50, 189 Urach, 28-9

Waco uplift, 256 Wales, south, 90-3, 96-8, 169, 177, 179, 181-2, 205, 224, 227 Weardale granite, 140, 143-5 Welsh massif, 26, 30 Wessex basin, 76 West African craton, 182 West Virginia, 203 Westphalian, 65-6, 68-9, 91-2, 98, 139, 185, 204-6, 211,224-5, 235-6, 248-9 Whin skill & dykes, 139-45 Wildenfels, 33, 75 Wildflysch, 35 Willimantic fault, 245, 248 Wiltshire 26, 125 Windsor, 203-4, 206, 221 Woonsocket basin, 204-5, 236 Wrench (strike-slip) fault, 77, 93, 99, 141, 144, 154, 169, 178, 189, 201-2, 206, 223,233-4, 239 fault (dextral), 133-7, 141,154, 158, 167, 181, 187-8, 193, 202, 206-7, 221,224, 228, 237-8, 240, 249 fault (sinistral), 141,145, 189, 206-7 tectonics, 26, 76-7, 133-8, 187-93, 212 Youghal, 168-9 Zemmour fault, 181,249