Variscan Tectonics of the North Atlantic Region
Variscan Tectonics of the North Atlantic Region
edited by D. H. W. H...
49 downloads
980 Views
39MB Size
Report
This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!
Report copyright / DMCA form
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 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
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 ._.. 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
;~.~