Post-Variscan (end Carboniferous–Early Permian) basin evolution

Transcription

Post-Variscan (end Carboniferous –Early Permian) basin evolution in
Western and Central Europe
T. MC CANN1, C. PASCAL2,3, M. J. TIMMERMAN4, P. KRZYWIEC5, J. LÓPEZ-GÓMEZ6, A. WETZEL7,
C. M. KRAWCZYK8, H. RIEKE9 & J. LAMARCHE10
1
Geologisches Institut, Bonn University, Nußallee 8, 53115 Bonn, Germany (e-mail: [email protected])
2
Vrije Universiteit, De Boelelaan 1085, 1081 HV Amsterdam, Netherlands
3
Present address: NGU, Geological Survey of Norway, N-7491 Trondheim, Norway
4
Institut für Geowissenschaften, Universität Potsdam, Karl-Liebknecht-Strasse 24, 14476 Potsdam, Germany
5
Polish Geological Institute, ul. Rakowiecvka 4, 00-975 Warsaw, Poland
6
Instituto de Geologı́a Económica, CSIC-UCM Facultad de Geologı́a, 28040-Madrid, Spain
7
Geologisch-Paläontologisches Institut, University of Basel, Bernoullistrasse 32, CH-4046, Basel, Switzerland
8
Geoforschungszentrum, Telegrafenberg, 14473 Potsdam, Germany
9
PanTerra Geoconsultants B.V., Weversbaan 1 – 3, 2352 BZ Leiderdorp, Netherlands
10
Université de Provence, Unité CNRS Géologie des Carbonatés, Place Victor Hugo, Case 67, 13331 Marseille, France
Abstract: The Variscan orogeny, resulting from the collision of Laurussia with Gondwana to form the supercontinent of Pangaea, was
followed by a period of crustal instability and re-equilibration throughout Western and Central Europe. An extensive and significant
phase of Permo-Carboniferous magmatism led to the extrusion of thick volcanic successions across the region (e.g. NE German
Basin, NW part of the Polish Basin, Oslo Rift, northern Spain). Coeval transtensional activity led to the formation of more than 70 rift
basins across the region. The various basins differ in terms of their form and infill according to their position relative to the Variscan
orogen (i.e. internide or externide location) and to the controls that acted on basin development (e.g. basement structure configuration).
This paper provides an overview of a variety of basin types, to more fully explore the controls upon the tectonomagmatic –sedimentary
evolution of these important basins.
The Permo-Carboniferous collision of the continents of Gondwana
and Laurussia, termed the Variscan orogeny, led to their amalgamation and the formation of the late Palaeozoic supercontinent
of Pangaea. Oblique convergence resulted in collisional processes
in the Appalachians and the Urals, whereas sinistral wrench faulting caused widespread rifting of the Northern European crust
(Pegrum 1984a,b; Ziegler 1990). In addition, the collapse of the
thickened Variscan orogenic crust resulted in late orogenic
crustal extension. The post-Variscan period was one of intense
crustal re-equilibration and reorganization under an alternating
transtensional and transpressional tectonic regime, and the combined effect of re-equilibration and tectonic activity controlled
the kinematic patterns and subsidence of approximately
70 basins, all of which are characterized by a major strike-slip
component in their deformational history. Western and Central
Europe was already thermally weakened by the preceding
orogeny and most of the Carboniferous– Permian basins trace
long-lived Variscan fault systems (Henk 1993). Subsequent
basin evolution involved extensive, predominantly clastic sedimentation (e.g. Glennie 1990; Maynard et al. 1997), and some
of the newly formed rifts became the loci of extensive intraplate
magmatism (e.g. Neumann et al. 2004).
In recent times, there have been a number of volumes published
on the Variscan history of Europe (e.g. Dallmeyer et al. 1995;
Franke et al. 2000). Although these books provide much-needed
information concerning the nature and style of deformation during
the Variscan orogeny, they are not really concerned with the postVariscan evolution of the region. One of the main problems with
unravelling the post-Variscan deformational history of Western
Europe is the localized nature of deformation. Small isolated
grabens were gradually filled by predominantly locally derived
sediments and/or associated volcanic and volcaniclastic rocks.
The nature of the sedimentation (predominantly alluvial) constitutes
another problem, as correlation is much more difficult, both within
and, more importantly, between grabens. The nature of the
sedimentary record within the basins that formed as a result of
Permo-Carboniferous wrench-fault activity, therefore, frequently
precludes detailed investigation, as it is difficult to correlate the
sedimentary record from one basin to another. In those Early
Permian basins that are rich in biogenic remains it may be possible
to correlate within the basin (e.g. Schneider 1989, 1996). However,
the general level of biostratigraphic uncertainty leads to problems
with correlation in this stratigraphic interval. It is only when basinwide crustal subsidence gave rise to the widespread Northern and
Southern Permian basins, with the establishment of a unified depositional pattern across Northern and Central Europe, that areally more
extensive correlation becomes possible (e.g. McCann 1998a). There
are, however, many Permo-Carboniferous basins that are outside
the Southern and Northern Permian basins region, and such
basins provide much-needed understanding of the broader evolution
of the Permo-Carboniferous transition in terms of magmatic, tectonic and sedimentary activity outside the area of the former
Carboniferous foreland basin. The following summary, by nature
selective, reviews the evidence for basin formation at this time
and investigates a selection of these basins. Of marked interest is
the similarity in terms of basin infill (sedimentary and magmatic)
between the basins, despite differences in both timing and location.
The aim of this study, therefore, is to provide an overview of the
main areas of basin initiation that followed the cessation of Variscan
orogenic activity in Western Europe. The following sections will
outline the regional framework of Western Europe and provide an
introduction to the most recent research in these areas, including
sedimentology, tectonics, magmatic history and basin modelling.
Background geology and palaeogeographical setting
The Variscan belt is a broad (c. 1000 km) complex curvilinear
feature extending across Europe and marking the zones of
Variscan-age deformation (Fig. 1). Variscan orogenic activity was
From: GEE , D. G. & STEPHENSON , R. A. (eds) 2006. European Lithosphere Dynamics.
Geological Society, London, Memoirs, 32, 355–388. 0435-4052/06/$15.00 # The Geological Society of London 2006.
355
356
T. MC CANN ET AL.
Fig. 1. The main tectonic elements of
Central and Western Europe (modified
after Ziegler 1990; Berthelsen 1992). Inset
map shows the tectonic zones
corresponding to volcanic belts:
MZ, Moldanubian Zone; OMZ, Ossa–
Morena Zone; RHZ, Rheno-Hercynian
Zone; SPZ, South Portuguese Zone;
STZ, Saxo-Thuringian Zone; TTZ,
Tornquist–Teisseyre Zone. Massifs: a,
Armorican; b, Bohemian; bf, Black Forest;
c, Massif Central; h, Harz Mountains; ic,
Iberian Cordillera; rh, Rheinische
Schiefergebirge; v, Vosges (after Francis
1988).
associated with the convergence of the southern continent of Gondwana with the northern continent of Laurussia (i.e. Laurentia, Baltica
and Avalonia), to form the supercontinent of Pangaea and leaving a
relict Palaeotethys to the east (Scotese & Langford 1995; Fig. 2). The
reconstructed geometry of the Variscan orogen can be divided into a
number of distinct geotectonic zones, many of which are separated
by steeply dipping faults or shear zones, that exhibit a general
continuity around the belt. Some of the boundaries may represent
suture zones and these are marked by the occurrence of ophiolite
assemblages and calc-alkaline continental arc volcanic rocks of
Devonian or Carboniferous age. The high-grade core of the Variscan
orogen runs through central and NW Spain, France, Germany and the
Bohemian area of the Czech Republic (Figs 1 and 3). The Devonian
to Early Carboniferous evolution of the Variscan orogenic system
was governed by alternating tensional and compressional tectonic
cycles, reflecting the development of an essentially intra-continental
Pacific-type, back-arc system (Ziegler 1990). Four principal phases
of deformation and exhumation may be recognized, each of which
lasted between 20 and 39 Ma and was restricted in geographical
extent. These phases probably resulted from the successive
docking of continental lithospheric fragments along the southern
margin of Laurussia. The first of these periods, the Ligerian Phase
(Late Silurian–Early Devonian), was contemporaneous, but not
cogenetic, with the Acadian phase of the Caledonian orogeny. The
remaining three phases, Bretonian (Late Devonian–Early Carboniferous), Sudetian (Viséan–Early Namurian) and Asturian (Westphalian–Early Permian), are assigned to the Variscan orogeny
(Warr 2000).
During the late Viséan the Variscan orogenic system entered the
Himalayan-type continent– continent collision stage. Continued
dextral-oblique convergence of Gondwana and Laurussia and
the progressive closure of the Proto-Atlantic and western
Proto-Tethys oceans was accompanied by the propagation of
their collision front into the Appalachian and central Mediterranean domains. At the same time, major crustal shortening was
achieved in the Variscan fold belt as indicated by the occurrence
of major nappe structures, in part involving basement in the Moldanubian area, the southern Massif Central and the Variscan externides (Behr et al. 1982; Burg et al. 1984; Cazes et al. 1986; Behr &
Heinrichs 1987). During the late Viséan and Namurian, the collisional front between Africa and the Southern European margin
propagated rapidly westwards and eastwards.
Following the Late Westphalian –Early Stephanian consolidation of the Variscan fold belt, convergence between Gondwana
and Laurussia apparently changed from an essentially north–
south-directed collision to an east– west convergence. During
Late Carboniferous times a subduction zone developed along the
Variscan Deformation Front, with oceanic crust being subducted
beneath the continental Variscan system (Gast 1988). Destruction
of the subducted crust led to stretching (dominantly east – west)
and associated block faulting. Coeval counter-clockwise rotation
of the southern African plate against the stable northern European
craton caused wrenching along NW– SE-trending faults (Ziegler
1990). Continental-scale dextral shears (e.g. the Tornquist–
Teisseyre fracture zone) were linked by secondary sinistral and
dextral shear systems (Ziegler et al. 2006). The westward translation of Gondwana relative to Laurussia along a 3500 km long
dextral megashear has been proposed, but recent work from northern Spain does not support the existence of such a major structure
(Weil et al. 2001). It is, however, possible that a series of smaller
shear structures were active (see the following section).
The development of the Variscan orogen involved major crustal
shortening and subduction of substantial amounts of supra-crustal
rocks, continental and oceanic crust, and mantle-lithosphere
POST-VARISCAN BASIN EVOLUTION, EUROPE
357
From late Early Carboniferous times, compression within Central
Europe was continuous for c. 45 Ma, with Variscan deformation
advancing northward at a rate of c. 0.5 cm a21 (Ahrendt et al.
1983). This is only slightly less than the value of 0.79 cm a21 that
characterized Alpine orogenic compression (Schmid et al. 1996).
In western Germany, for example, a series of major north- and southverging, basement-cored nappes were emplaced during Namurian
and Westphalian times. The widespread occurrence of low-pressure
metamorphic rocks and late to post-orogenic calc-alkaline intrusive
rocks in the internal parts of the Variscan orogen suggests that a significant amount of crustal shortening, accompanied by crustal delamination, subduction and anatectic remobilization of lower crustal
material, and partial melting of upper mantle material, occurred
(Ziegler 1984, 1986). Rough estimates of total crustal shortening
yield a value of c. 150–200 km (Ziegler 1990). This is less than
half of the crustal shortening recorded from the Alpine Realm
during the Tertiary, which was estimated at c. 500 km (Dewey
et al. 1989; Schmid et al. 1996). By latest Westphalian times, the
Variscan fold-and-thrust belt of Western and Central Europe was
consolidated and inactive (Ziegler 1990). However, localized tectonic activity is evidenced by diverse angular unconformities
between the Westphalian and Permian deposits (e.g. NE German
Basin; McCann 1999). In Germany, post-orogenic uplift of the
Rheno-Hercynian thrust belt, preceding deposition of the Permian
sediments, amounted along its northern margin to 2–3 km, increasing southwards to 6 km (Littke et al. 2000), with values of up to
10 km in the area of the Saar–Nahe Basin (Oncken et al. 2000).
The eastwards extent of Variscan deformation
Fig. 2. Early (a) and Late (b) Permian palaeogeography (after Scotese &
Langford 1995).
(Ziegler et al. 1995). The convergence rate between Gondwana
and Laurussia, however, was not constant during Devonian and
earliest Carboniferous times, as suggested by the Early Devonian
development of an extensive back-arc rift system in Western and
Central Europe, which remained intermittently active until the late
Viséan (Zieger 1990). Full-scale collision between Gondwana and
the southern extension of Laurussia occurred during the late
Viséan at a time when the back-arc extensional system that had
governed the early evolution of the Variscan orogen in Western
and Central Europe finally became replaced by a compressional
stress regime (Ziegler 1990). Gondwana and Laurussia became
increasingly coupled during the Early Carboniferous, as is
evident from their joint clockwise rotation and northward drift
(see Ziegler 1990, Fig. 8).
As noted above, the collision between Gondwana and Laurussia
continued to develop until late in the Carboniferous–Early
Permian, at which time the intercontinental collision began to
affect the northwestern part of Africa (e.g. the West African
orogens have a Late Carboniferous to Early Permian age; Lécorché
et al. 1989). As a result of the extensive collision episode, a
central Pangaean mountain range was formed, extending from
Mexico to Poland (Golanka & Ford 2000) and southwards to
Morocco (Pique & Michard 1989). Late Carboniferous events
were also marked in the Alps (e.g. Matter et al. 1987), the
Carpathians (Dallmeyer et al. 1995) and the Rhodope area
(Yanev 1992). Further eastwards, the situation was a more
complex one, as it involved the rifting (from Late Carboniferous
times onward) from Gondwana of several Cimmerian continents
(e.g. parts of Turkey, Iran and Afghanistan; Sengör et al. 1984;
Scotese & Langford 1995; Fig. 4). During the main phase of the
Variscan orogeny (late Viséan to Westphalian), the collision
front between Gondwana and Laurussia propagated eastward and
southwestward in conjunction with the progressive closure of the
Palaeotethys and Protoatlantic oceans. By Westphalian times,
Gondwana had collided with North America whereas to the east
Palaeotethys remained open. Therefore, the western parts of the
Variscan megasuture were characterized by a Himalayan-type
(continent– continent collision) setting, whereas its eastern parts
remained in an Andean-type (continent– ocean collision) setting
(Stampfli 2000; Ziegler & Stampfli 2001). The setting of the
West European segment of the Variscan orogen was, therefore,
transitional between a Himalayan- and an Andean-type setting
(see Ziegler et al. 2006, for details). During the Stephanian and
Autunian orogenic movements continued in the Appalachian –
Mauretanides (Rodgers 1970; Michard & Sougy 1977) and in the
Urals (Ivanov et al. 1977), whereas the Variscides region remained
largely inactive. These latter movements were accompanied by the
emplacement of a right-lateral transform fault system, which
linked the southern Uralides and northern Appalachians, and
which crossed Europe, where it caused the development of a
complex pattern of conjugate shear faults and related pull-apart
structures (Arthaud & Matte 1977; Ziegler 1978a).
358
T. MC CANN ET AL.
Fig. 3. Structural sketch map of the
European Variscan orogen. Three
external coal basins are shown in black.
C.C.S., Coimbra–Cordoba suture;
L.R.S. Lizard–Rhenish suture; N.V.F., North
Variscan Foreland; M.C.S., Massif Central
suture; M.T.S., Münchberg–Tepla suture;
O.M.S., Ossa– Morena suture. The Iberia
and Corsica–Sardinian blocks are
represented in their possible Permian
positions relative to Europe (after
Matte 1991).
An important problem is the continuation of the Variscan
orogen in SE Europe and Turkey. In Western Europe, subduction
ended with the collision of NW Gondwana (i.e. Morocco, Algeria)
with SW Europe (i.e. Iberia, Corso-Sardic block). Further to
the east (i.e. Tunisia, Libya, Egypt) there is no evidence of
subduction-related activity (e.g. intense folding, subductionrelated magmatism) at the northern margin of Gondwana during
the Late Palaeozoic (Dornsiepen et al. 2001). In contrast, orogenic
activity can be traced along the southern margin of Laurussia from
the Alps to the Caucasus, although it appears that the closure
of Palaeotethys was asymmetrical, with subduction occurring
on the northern (i.e Laurussian) margin (Vavassis et al. 2000).
It is probable that the active spreading ridge of Palaeotethys
arrived at the southern margin of Laurussia around the
Permo-Carboniferous boundary, leading to a cessation in active
subduction, and the convergent margin was then transformed
into a dextral strike-slip movement zone (Dornsiepen et al.
2001; Fig. 4). Stampfli et al. (2001) noted that there is a general
pattern of rifting across the region of Tethys. The dynamic evolution of the region of Neo-Tethys (India, Arabia, Turkey)
reveals that the Early Carboniferous rifting phase was followed
by a second rifting phase in the Early Permian and the late Early
Permian– Late Triassic opening of Neotethys. A similar pattern
of rift activity is recorded in the East Mediterranean area, where
the initial Carboniferous rifting phase affected the southern
margin of the area between Syria and Tunisia. This was followed
by a Permian rifting phase and protracted Late Permian subsidence
(extending to the present day).
Following late-stage subduction of Palaeotethys, and in conjunction with slab roll-back (since Early Permian times)
and slab detachment beneath the Variscan orogenic belt, the
Cimmerian Terrane was detached from the Gondwana margin
and the Mediterranean– Neotethys Basin began to open along
the Eurussian margin. The post-collisional Permo-Carboniferous
rifting phase accompanied the transtensional collapse of the
overthickened Variscan lithosphere in a basin-and-range fashion
(Malavielle 1993). This analogy implies the presence of a
still-active or transform margin in areas where the Palaeotethys
was not yet closed during the Permian. The mid-ocean ridge of
the Palaeotethys was obliquely subducted beneath the active
Eurasian margin during the Permo-Carboniferous (Stampfli
1996). This was responsible for the widespread Late Carboniferous and Early Permian magmatism and volcanism characterizing
the southern Variscan domain.
Late to post-Variscan structures
During the main phase of the Variscan orogeny, lateral escape
tectonics was accompanied by the development of synorogenic
transtensional basins in the Armorica– Iberia– Avalonia domain.
During the latest Carboniferous and Early Permian the convergence of Gondwana and Laurussia changed from oblique collision
to dextral translation controlling the Alleghenian orogeny of the
Appalachians. At the same time the Variscan orogen was transected by conjugate wrench faults, partly terminating in pull-apart
basins (Ziegler 1990; Matte 1991). Along the northern margin
of Gondwana, rifting resumed during the Early Carboniferous,
reflecting an intraplate stress reorganization following the
docking of the composite Hun terrane to the Laurussian margin
(see Stampfli et al. 2001, for details).
A number of major late to post-Variscan tectonic structures have
been recognized. The Tornquist Zone, comprising a fan-shaped
series of fault and fault-zone splays was activated in the western
Baltic area (Berthelsen 1992). Indeed, right-lateral wrench movements along the Tornquist– Teisseyre lineament led to the formation of the Oslo – Skagerrak Graben system (Ziegler 1978b;
Neumann et al. 1992). According to Ziegler (1990), a major postVariscan wrench fault was active along the present Elbe Fault
System (northern Germany) as part of a complex megashear
zone that developed in North – Central Europe between the Appalachians and Uralides as a result of dextral translation of the
Armorican – European plate with respect to the African plate
(Arthaud & Matte 1977). Along the SE end of this wrench fault,
geological evidence for Late Carboniferous-age dextral ductile
shearing has been found in outcrops of the Elbe Zone (Mattern
1996). Scheck et al. (2002) have noted that maximum deformation
occurred along the Elbe Fault System during late Carboniferous
359
Teisseyre lineament. Displacements along the Bay of Biscay
fault system were in part taken up in the Arctic –North Atlantic
rift, in which crustal distension continued during the latest
Palaeozoic, as illustrated by the East Greenland graben system
(Haller 1971). Indeed, from the late Early Carboniferous
onwards, Laurussia was transected by the Arctic – North Atlantic
rift system, which was partially superimposed on the Caledonian
suture zone (Ziegler 1990).
Palaeogeography and palaeoenvironments
Fig. 4. Schematic plate configuration and plate boundary geometry
between Europe and Gondwana. (a) Early Permian. (b) Mid–Late Permian.
(c) Mid–Late Triassic. (After Dornsiepen et al. 2001.)
wrenching and volcanism (and subsequent to these events). Late
Palaeozoic dextral transtensional movements with estimated
displacements of 60– 120 km have been postulated. The major
Variscan faults in Germany and Poland (i.e. Elbe Zone, Main
Intra-Sudetic Fault, Odra Fault Zone) are considered to have
been dextral shears trending NW– SE during the late Devonian–
early Carboniferous with maximum offsets of 50 – 300 km
(Aleksandrowski 1995; Alexandrowski et al. 1997) and dextral
offsets of up to 400 km along the Teisseyre – Tornquist Zone
(Lewandowski 1993). In late Carboniferous–early Permian
times, many local and regional WNW –ESE- to NW –SE-oriented
fault or shear zones in the Western Sudetes were reactivated under
a semi-brittle regime.
The main elements of the Late Variscan fracture system of
northern Africa and Europe are the Agadir– Kelvyn, Gibraltar
and Bay of Biscay fault systems, as well as the Tornquist–
Lithologically, the Upper Carboniferous successions across
the region are relatively monotonous. This is partly due to the
similarity in terms of lithological and stratigraphical evolution
between the Variscan fold-and-thrust belt and that of its foreland
basins, related both to the infilling of the basins and the gradual
cessation of Variscan-age tectonic activity. Only the Namurian
A still shows the orogen –foreland differentiation. During Late
Carboniferous times, paralic coalfields extended across Laurussia
from North America across Britain and the Rheno-Hercynian
region of Germany to Silesia and the Donets region, Ukraine.
Coal-forming environments were present south of the Rheic
suture in the Saar – Lorraine Basin during late Westphalian times
(possibly earlier) and continued until early Stephanian times.
After a short hiatus, these conditions were re-established in
mid-Stephanian times (Cleal 1984). This extensive region
experienced several marine incursions and clearly most of the
Westphalian deposition (largely non-marine) occurred close to
sea level (Paproth 1991) on the northern (passive margin) side
of the Rheic suture. In contrast, to the south of the suture mainly
intermontane limnic basins (of Stephanian age) developed on
the various components of the Armorican Terrane Assemblage
(McKerrow et al. 2002).
During the Permian, continental size and aggradation were at a
maximum, as few continental fragments escaped incorporation
into Pangaea. The new continent, stretching almost from pole to
pole, had a significant amount of terrestrial area. A single world
ocean, Panthalassa, with a semi-enclosed Tethyan Sea, dominated
the marine environment. This singular continental configuration
led to the development of extreme climatic conditions. Furthermore, the Permian climate was not a stable one, with evidence
of profound changes throughout the duration of the period
(Veevers & Powell 1987; Barron & Fawcett 1995; Parrish 1995).
Climate was not the only unstable factor. Given the sheer size
of Pangaea, the supercontinent was unstable from the outset.
The timing of break-up was varied, beginning within the late
Carboniferous immediately north of the Variscan orogen. In
many places (e.g. Germany and Belgium), extension to the north
of the Variscides occurred coevally with thrusting and strikeslip faulting further south (Ziegler 1990). However, the timing
and extent of individual phases of extension and rifting
throughout the North Atlantic (and associated) rift systems are
still subject to some debate, as the dating of Late Carboniferous
and Early Permian red beds is imprecise. Following the
main phases of Variscan compression, thermal relaxation of the
crust occurred in Early Permian times, creating the rifts
and graben that allowed accumulation of the first phase of
sedimentation.
Stephanian– Autunian magmatic activity
During the Permo-Carboniferous, the Variscan foreland in Europe
was subjected to extensive rift-related tectonism and related
extensional magmatic activity (Fig. 5). The rare occurrence of
Autunian-age sediments suggests that much of the area later
occupied by the Southern Permian Basin was regionally uplifted
and subjected to profound erosion. This uplift, probably induced
360
T. MC CANN ET AL.
Fig. 5. Distribution of late Carboniferous and early Permian magmatic rocks in NW Europe (after Ziegler 1990). NGB, North German Basin; OS, Oslo Rift;
WSC, Whin Sill Complex; VF, Variscan Front; RFH, Ringkøbing–Fyn High; OR, Oslo Rift; MV, Midland Valley.
by a combination of wrench-related lithospheric deformation,
magmatic inflation of the lithosphere and thermal erosion of the
mantle lithosphere, was coupled with periods of significant
magmatic activity. Melt generation was probably related to localized divergent wrench-induced decompressional partial melting
of the uppermost aesthenosphere and the lithospheric thermal
boundary layer, possibly combined with the impingement of a
not very active mantle plume on the base of the lithosphere
(Ziegler 1996). Although the most voluminous and evolved
magmatism occurred in the North German Basin (Benek et al.
1996), dated at 297– 302 +3 Ma (Breitkreuz & Kennedy 1999),
the Oslo Rift contains the most extensive and best-preserved
sequences of basaltic lavas associated with this event. In the
latter, Rb –Sr age determinations indicate that the total period
of magmatic activity extended from c. 305 to 245 Ma (Sundvoll
et al. 1990), and the earliest basaltic magmatism appears to have
been restricted to a relatively short period (305– 290 Ma)
(Neumann et al. 2002). However, recent dating suggests that
these ages are slightly too young and that the duration of the
various magmatic periods in the Oslo Graben could be shorter
than proposed by Sundvoll et al. (1990) (see Neumann et al.
2004). The early basalts from the Oslo Graben contain the least
evolved magmatic products in the region and provide information
about the primary magmas and the magma sources involved in the
magmatic activity.
Volcanic activity, however, was widespread throughout Europe,
with thick sequences being deposited elsewhere; for example,
in Spain (e.g. Iberian Range, Lago & Pocovi 1984), dated at
282 + 12 Ma (Hernándo et al. 1980), and in the Italian Alps
(e.g. Collio Basin, Breitkreuz et al. 2001), dated at 283 + 1 Ma
and 281 + 2 Ma (Schaltegger & Brack 1999), Liguria and
Sardinia (e.g. Cortesogno et al. 1998). Autunian-age magmatic
activity also included the intrusion of the Whin Sill dolerite
(northern England), some dolerite dyke swarms in the Midland
Valley of Scotland and the Northumberland Basin (Francis
1978; Coward 1995), and in the Argyll area of Scotland
(Speight & Mitchell 1979), and episyenite dykes along tensional
structures in the Central Range of Iberia (González-Casado et al.
1996). In these areas volcanism followed Westphalian-age
inversion and predated much of the Permo-Triassic succession,
suggesting that the region was underlain by hot asthenosphere.
The widespread Stephanian– Early Permian (305– 285 Ma)
alkaline intrusive and extrusive magmatism of the Variscan area
and its northern foreland is mantle derived and shows evidence
of strong crustal contamination (e.g. Marx et al. 1995; Benek
et al. 1996; Breitkreuz & Kennedy 1999). Melt generation
by partial melting of the uppermost aesthenosphere and the
lithospheric thermal boundary layer was probably triggered by a
rise in the potential temperature of the aesthenosphere and localized divergent wrench-induced decompression. Aesthenospheric
upwelling was presumably triggered by detachment of subducted
lithospheric slabs, resulting in a reorganization of the mantle
convection system and the impingement of a not very active
system of mantle plumes on the base of the lithosphere.
There are at least two discrete tectonic settings for sedimentary
basins that formed simultaneously during the Late Palaeozoic
Variscan orogeny in Western Europe. Immediately adjacent to,
and to the north of, the Variscan deformation front, narrow foredeep basins, termed the external basins (Variscan externides),
are interpreted to have formed as a result of subsidence related
to thrust loading of the crust. The related internal basins (Variscan
internides) comprise those basins formed within the orogenic belt
and that were structurally controlled. The following discussion of
the magmatic activity within the Variscan-influenced regions
of Europe will examine both basin types.
Extrusive activity (Variscan foreland– externides)
In the foreland of the Variscan region, sedimentation and magmatism generally occurred in and close to the north- and
NW-trending grabens. Widespread magmatism occurred as
felsic to mafic volcanism, and emplacement of mafic dykes and
sills. In the Oslo Graben and Central North Sea, extension,
block faulting, tilting, uplift and erosion of structural highs
occurred simultaneously with, and subsequent to, the initial
phase of volcanic activity. As noted above, the most voluminous
magmatic activity occurred in the Oslo Rift, with more than
100 000 km3 of produced melts (Neumann et al. 2004), and the
NE German Basin, where more than 48 000 km3 of magmatic
material was extruded. Five stages of magmatism have been
recognized and dated in the Oslo Graben (Olaussen et al. 1994;
Neumann et al. 2004). The pre-rift stage (i.e. between 304
and 294 Ma) comprises felsic sill emplacement in underlying
Westphalian sediments. The second stage corresponds to rift
initiation and is characterized by widespread basaltic volcanism
(i.e. B1 basalts) representing the most primitive magmatic rocks
in the Oslo Graben. The main rift stage (i.e. Stage 3) resulted
in the fissure eruption of thick sequences of porphyritic trachyandesite flows (rhomb porphyries), which was accompanied by
the intrusion of large amounts of syenitic magmas that have
yielded ages of 292– 298 Ma according to recent U– Pb dating
(Neumann et al. 2004). During Stage 4 rifting abated and the
magmatic style changed to the development of central volcanoes
and caldera collapse. Rb – Sr dating of associated magmatic
rocks gives ages between 280 and 240 Ma (Sundvoll et al.
1990). The final magmatic stage was dominated by the
intrusion of composite batholiths of trachyandesite to rhyolitic
compositions.
In the subsurface of Northern Germany and Poland extensive
Late Carboniferous to Early Permian felsic to intermediate volcanic rocks have been cored by exploration wells (e.g. Hoth et al.
1993). In this region, centres of volcanic activity appear to
coincide with the intersection of fault systems (Plein 1978) and
are probably related to pull-apart structures at the termination of
subsidiary wrench faults that parallel the Tornquist– Teisseyre
lineament. In Northern Germany the magmatic succession is up
to 2000 m thick, and comprises predominantly calc-alkaline
SiO2-rich volcanic rocks (Kramer 1977; Marx et al. 1995;
Benek et al. 1996). The volcanic and pyroclastic succession of
Northern Poland is broadly similar to that of NE Germany and
comprises mainly acid volcanic rocks (rhyolites, rhyodacites)
and locally abundant trachytes and trachybasalts (Pokarski 1988,
1989). In the northern part of the Fore-Sudetic Monocline, there
is an Early Rotliegend-age volcanic succession composed of
lava flows, pyroclastic deposits, and hypabyssal and intrusive
rocks (Jakowicz 1994). The rocks, however, are not comagmatic,
as they were derived from two sources, the mantle and crust. The
parent magma was a high-aluminium basalt, but anatectic crustal
fusion produced acid melts that mixed with the derivatives
of the parent magma to produce poorly homogenized hybrids
with intermediate compositions. The distribution of these
rocks is parallel to the trend of regional tectonic units of Early
Rotliegend age, and their thickness varies from a few metres to
over 1500 m. In general, the volcanic rocks of the Polish Basin
are limited to its western margin and overall only attain a thickness
of c. 100– 200 m (with exceptions, see above), and are, thus, an
order of magnitude less than in the adjacent NE German Basin
(Karnkowski 1999).
In the offshore regions of Northern Europe it is more difficult to
determine the extent of magmatic activity. For the Central North
Sea, Denmark and Skagerrak Graben there is geophysical
evidence of mafic intrusions in the middle and lower crust. Furthermore, in the Central North Sea area, the thick basalts and
tuffs (,160 m) of the Inge Volcanics Formation (c. 299 Ma,
Heeremans et al. 2004) are interbedded with Rotliegend-age
mudstones and sandstones, whereas bimodal suites of basalt,
trachyandesite and rhyolite flows occur in the eastern North Sea.
However, in other parts of the Central and Northern North Sea
and along the Scottish– Irish Atlantic seaboard, the importance
of the Stephanian – Autunian tectonism is difficult to assess.
361
Many of these latter areas, however, were subjected to uplift
(e.g. Dunlap & Fossen 1998).
Intrusive activity (Variscan foreland– externides)
Stephanian to Early Permian-age intrusive activity in the Variscan
foreland occurred largely in the form of dolerite sills and
dyke swarms. For example, the Whin Sill Complex in northern
England comprises a series of sills (,90 m thick) and four
major ENE-trending dyke echelons of high-Fe, sub-alkaline
basalt (Johnson & Dunham 2001). The total volume of the
dyke complex has been estimated at 120–215 km3; however, as
it extends eastwards under the North Sea, this figure is a
minimum. In the Midland Valley of Scotland, earth –
west-trending dykes (up to 50 m wide and 130 km long) and a
sill (,180 m thick) are composed of sub-alkaline to transitional
basalt, with a total volume of almost 500 km3. This dyke
swarm also extends c. 200 km eastwards into the North Sea
(Smythe 1994). The Midland Valley of Scotland also contains
Westphalian– Early Permian-age pyroclastic volcanic rocks,
vents, necks, alkaline dolerite sills, dykes and plugs of basaltic
to trachytic and phonolitic composition (Francis 1992). Many of
the vents contain mantle and crustal xenoliths of high-grade
mafic and felsic gneisses. The mantle xenoliths probably reflect
derivation from variously metasomatized mantle sources, but
biotite, amphibole and/or alkali feldspar-bearing ultramafic
xenoliths and some of the megacrysts may be fragments of coarsegrained intrusions that fractionated under high-pressure conditions
in the upper mantle. The mafic crustal xenoliths may represent
metamorphosed cumulate or underplated mafic mantle melts,
whereas the felsic gneisses were derived from Precambrian to
Palaeozoic crust. In the western Scottish Highlands at least 3000
Early Permian, SW- to NW-trending camptonite and monchiquite
dykes, many containing mantle xenoliths, intruded the Caledonian
basement (although some of these dykes may be of Tertiary age;
Rock 1983).
In southern Sweden a c. 70 km wide swarm of NW-trending
sub-alkaline basalt and basaltic andesite dykes occurs. Individual
dykes can be up to 100 m wide but the majority vary between 1
and 50 m; their total volume has been estimated at c. 4000 km3
(Obst et al. 2004). Of similar age may be the isolated, NNWtrending dolerite dykes on the SW coast of Sweden and the dolerite sills that intruded Cambrian alum shales in south –central
Sweden (Västergötland). These dykes and sills may be related to
volcanic rocks obtained from drill core in the Kattegat and offshore eastern Denmark and interpreted from seismic data as volcanic edifices (Mogensen 1994; Marek 2000). Drill core from the
island of Rügen (NE Germany) contains a few basalt sills that
intrude lower Palaeozoic sediments, and that are also interpreted
as having an age of late Carboniferous– early Permian (Korich
& Kramer 1994).
Magmatic activity in the Variscan internides
Compared with the foreland and internal Variscides, post-tectonic
magmatism in the Rheno-Hercynian foreland fold-and-thrust belt
was largely restricted to the intrustion of granites (c. 295–293 Ma)
in SW England and granitoids (c. 290– 307 Ma), gabbros, and
rhyolitic to basaltic sills and dyke swarms in the Mid-German
Crystalline Rise and Harz area in Germany. Stephanian–
Autunian-age magmatism in the internal parts of the Variscan
orogen is dominated by granitoid intrusions exposed in midcrustal metamorphic terranes. Volcanic activity, on the other
hand, consisted of widespread ash-fall tuffs of distal provenance,
with evidence of more pronounced activity being locally
restricted. For example, the Ilfeld Basin (southern Harz region)
is a small dextral pull-apart basin that contains up to c. 400 m of
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T. MC CANN ET AL.
Autunian-age tuffs, ignimbrites and latitic, trachytic and rhyolitic
volcanic rocks that erupted from a large number of small centres to
overlie Stephanian-age sediments. To the west, the Saar – Nahe
Basin contains Autunian-age sub-alkaline basalts and andesite
flows that are associated with sub-volcanic dacite and rhyolite
domes and pyroclastic deposits, and that are interbedded with
fluvio-lacustrine sediments (Stollhofen 1994, 1998; Schmidberger
& Hegner 1999).
In the Western Mediterranean (Iberia, Pyrenees, Balearic
Islands, Sardinia, Corsica, Provence), Stephanian– Autunian-age
sedimentation occurred in deep basins. Here, volcanic activity
took place in at least two stages, separated by a period of strikeslip activity and granite intrusion (270– 290 Ma). The initial
magmatic stages are of predominantly calc-alkaline character,
whereas later, mid-Permian-age, magmatic rocks are alkaline. In
the Iberian Peninsula and the Pyrenees, deep basins were formed
in which sequences of late Westphalian-C to Stephanian-age
clastic sediments are overlain by Autunian-age volcanic rocks.
In some half-grabens in the Pyrenees, several cycles of
calc-alkaline rhyolitic to andesitic and alkaline basaltic volcanic
rocks have been recognized, ranging in age from Stephanian to
Early Permian. Late Carboniferous magmatic activity is evidenced
by thick sequences of lavas and ignimbrites, but Permian-age
volcanicity has also been established (Martı́ & Barrachina 1987;
Martı́ 1996). The Permo-Carboniferous volcanic rocks are represented by a suite of basaltic andesite to rhyolite of calc-alkaline
orogenic type, with a predominance of rhyolitic and rhyodacitic
volcaniclastic deposits. Volcanic activity was essentially continuous, and without any significant change in terms of its character,
from the deposition of the first Mid-Stephanian sediments to
the beginning of the Late Permian sedimentation (c. 18 Ma;
Martı́ 1996). During the Late Carboniferous, volcanism is represented by lavas and pyroclastic deposits mainly associated
with caldera-forming events. Indeed, at this time it would appear
that each basin had an independent volcanic history, with
several volcanic centres located around the basin margins. Most
of these deposits appear to be remnants of intra-caldera fill.
During the Early Permian only rhyodacitic and rhyolitic
magmas were extruded, and volcanic activity was concentrated
in one basin (i.e. Castellar de N’Hug Basin), although it is
possible that other volcanic centres were also active in the
Central Pyrenees at this time (Martı́ 1996; Fig. 3). Eruptions
were explosive, and generated widespread pyroclastic flows and
associated pyroclastic surge and fall deposits. In the northern
and central parts of the Iberian Peninsula, post-collisional
magmas of mafic composition are rare.
The magmatic rocks of the Massif Central area of France are
limited to centimetre- to decimetre-thick layers of strongly
altered air-fall tuffs (tonstein) in both the Stephanian and Autunian
sequences, and minor amounts of felsic to intermediate volcanic
rocks. In the Rodez Basin, c. 5 m of andesite lavas and c. 30 m
of trachytic tuffs occur at the base of Autunian sandstones and
are probably of early Autunian age. The Variscan basement
south of the Decazeville Basin is intruded by a topaz-bearing rhyolite dome and a later leucogranite, both of crustal origin (Badia &
Fuchs 1989). However, U– Pb zircon dating of a tuff
(332 + 4 Ma) and a rhyolite flow (333 + 2 Ma) at the base of,
respectively, the Bosmoreau and Decazeville basins showed
that some of the sediments and volcanic rocks, previously
thought to be of Stephanian age, were actually deposited in the
late Viséan (Bruguier et al. 1998).
and the external Variscides (i.e. the area of the northern foreland
basin, to the north of the Variscan orogenic belt) took place
within a relatively short time span (c. 300– 290 Ma) in a dextral
wrenching setting. In the internal parts of the orogen the basins
tend to be small, deep and isolated. Basin infill was a relatively
rapid process with the accumulation of thick sequences of
Rotliegend-age aeolian and fluviatile – lacustrine sediments.
Tectonic activity coincided with an overall change in climate to
semi-arid (Stephanian– Early Permian). As noted above, many
of the basins contain volcanic rocks, and the presence of these
magmatic successions aids in distinguishing the Permian
sequences from the underlying Stephanian (see Plein 1995). The
following section will examine a number of basins and regions
across Europe to provide an overview both of the main basin
types that formed at this time, and of the differences between
the basins that formed in front of and within the Variscan
orogen. The section is subdivided into those basins (Northern
and Southern Permian basins) that were located to the north of
the Variscan orogen, and those that were located within or to the
south of the orogenic belt (Fig. 6).
Basins occurring within the Variscan foreland
Permian basins in Europe
Southern Permian Basin (SPB). The Southern Permian Basin comprises a series of connected basins extending across Northern
Europe from England to Poland (Figs 1 and 6). The SPB, with a
north– south extension of 300–600 km and an east – west extension of c. 1700 km, developed between the northern foreland of
the Variscan mountain belt and the Ringkøbing– Fyn High. The
south– central parts of the basin are superimposed on the Variscan
fold-and-thrust belts and intervening Gondwana-derived blocks
(e.g. East Silesian Massif, SW part of the Malopolska Massif )
and on the Precambrian of the East European Craton. Despite
this, the geometry of the SPB shows no direct relationship with
either the different crustal domains or the suture zones on which
it subsided.
The SPB is of great economic importance, containing a number
of significant hydrocarbon finds (e.g. Groningen, Salzwedel–
Peckensen). By the early 1990s almost 180 wells had been
drilled into the Rotliegend of Western Germany, and in Eastern
Germany more than 1500 wells were drilled (see McCann et al.
2000, for details). The succession is subdivided into two distinct
units, separated by the Saalian unconformity (Schneider et al.
1995). The Lower Rotliegend is characterized by acid and intermediate volcanic rocks with only minor sediments, whereas the
Upper Rotliegend is essentially sedimentary and contains only
rare volcanic rocks, of a more basaltic composition (Fig. 7).
In the area of the Southern Permian Basin, up to 800 m of
Stephanian continental red beds, containing correlative marine
bands, were deposited in a broad successor basin to the
Namurian – Westphalian Variscan foreland basin. During the late
Stephanian –Early Permian, this basin was disrupted by predominantly transpressive wrench tectonics, as evidenced by the deep
truncation of Late Carboniferous series and the conspicuous
absence of deep Early Permian basins (Ziegler 1990; McCann
1999). These wrench tectonics were associated with the development of extensive volcanic fields (see above). Ziegler et al. (2006)
noted that crustal thinning within the North German part of the
SPB may be interpreted in terms of a ‘stretching’ factor of 1.45.
This thinning is attributed to late Stephanian– Early Permian magmatic destabilization of the crust – mantle boundary that was
paralleled by major thermal attenuation of the mantle-lithosphere
(Van Wees et al. 2000).
Following the end of Variscan contraction in latest Westphalian
times, subsequent Stephanian– Autunian magmatic activity and
basin formation in both the internal Variscides (i.e. the area
within the Variscan fold-and-thrust belt and to the south of it)
Southern Permian Basin: North German Basin. The North German
region, forming part of the later SPB, comprises crystalline basement of varying Precambrian ages (see McCann 1998b, for
details), which is covered by a thick (.12 km) sedimentary
363
Fig. 6. Outline of the Northern and
Southern Permian basins of NW Europe
(after Stemmerik et al. 2000).
succession of early and late Palaeozoic-, Mesozoic-, and
Cenozoic-age strata. Basin evolution commenced with the
destruction of the subducted oceanic crust of the Rhenohercynian
Ocean, which led to lithospheric stretching (dominantly east –
west) and associated block faulting (Late Carboniferous– Early
Rotliegend, and Early– Late Rotliegend; e.g. Gast 1991;
Fig. 7. Permian time scale (after Menning 1995). In addition, Menning (2001)
has provided an integrated Permian time scale based on absolute dates and
field indicators for the duration of permian stages. More recently, the
International Stratigraphic Chart (Gradstein et al. 2004) set the Carboniferous –
Permian boundary at 299 + 0.8 Ma (see also Stratigraphische Tabelle von
Deutschland, STD 2002; Menning et al. 2005).
Kiersnowski et al. 1995; Stemmerik et al. 2000; Kockel 2002),
accompanied by calc-alkaline magmatic activity (which was particularly pronounced in NE Germany), which resulted in the formation of a series of north– south-striking grabens. These acted
as a feeder rift system for sediment transport from the Variscan
hinterland to the areas to the north. The Lower Saxony rift
system and associated basins (e.g. Hessen Basin) were part of a
major continent-separating suture parallel to the subsequent MidAtlantic Rift (Gast 1988, 1991), which can be extended
northwards via the Glückstädter Trough (Schleswig– Holstein),
Horn Graben, and the Skagerrak Graben into the Oslo Graben
(Fig. 1). Analysis of the form and sedimentary infill of a series
of half-grabens along an east– west transect in Northern
Germany reveals variations in both, suggesting that there were
differences in both the rates and amounts of stretching.
In the NE German Basin (NEGB) the angular unconformities
between the the Westphalian– Stephanian and the overlying
Permian sequences reflect a period of tectonic activity (McCann
1998a). This was followed by the major Permo-Carboniferous
volcanic event that heralded the onset of Permian basin evolution.
Initial sedimentation (i.e. Autunian) was restricted and mostly
confined to isolated basins. At the onset of the latest Rotliegend
(i.e. Rotliegend II), there was a clear change in terms of the
basin geometry (e.g. Rieke et al. 2001, 2003). Initially, the basin
comprised two distinct sub-basins (Havel –Müritz and West
Mecklenburg basins), although it is not clear to what extent
these were isolated from one another (Fig. 8). However, c. 2 Ma
later there was a clearly unified depositional area across the
entire NEGB, with sediment being sourced from basin margin
highs, and adjacent orogenic piles, and transported towards the
basin centre (McCann 1998a). The distribution pattern within
the basin itself broadly resembles the models of closed-basin sedimentation as outlined by Leeder & Gawthorpe (1987) albeit with
the significant difference that the NEGB was not a half-graben
structure. Strongly increasing thermal subsidence modified the
facies architecture as well as the basin geometry through the
remainder of the Rotliegend II period, resulting in a broadly
smooth topography, a decrease in sediment supply and the expansion of a playa lake environment across almost the entire basin.
The NEGB shows no evidence of significant synsedimentary tectonism (e.g. Kossow et al. 2000), as described for the NW German
Basin, the Dutch Basin (e.g. Verdier 1996; Geluk 2005) and the
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T. MC CANN ET AL.
Fig. 8. Isopach maps of the
Rotliegend-age (a) Parchim, (b) Mirow,
(c) Dethlingen and (d) Hannover
formations, NE German Basin (after
McCann et al. 2000).
English Basin (e.g. George & Berry 1997). This absence presumably reflects the trend of gradually increasing tectonic activity in
these latter areas, which may be related to the initiation of the
break-up of Pangaea in Late Permian times.
Southern Permian Basin: Polish Basin. The Mid-Polish Trough
(MPT) was continuously connected with the North East German
Basin to the NW (Ziegler 1990), but not with the Tethyan
domain to the south (Dadlez et al. 1998; Figs 9 and 10). Rotliegend
development of the main depocentre of the MPT followed a period
of Westphalian wrench tectonics and Early Permian volcanic
activity (within the NW MPT and the NE Germany Basin)
associated with regional wrenching and generalized crustal destabilization. Immediately overlying the volcanic rocks are late
Rotliegend continental sediments deposited under arid conditions
(Marek 1988). The sediments are broadly similar to those deposited in the NE German Basin to the west, comprising mainly
coarser-grained clastic deposits (confined to the basin margins)
with finer-grained clastic deposits occupying the central parts of
the basin. Depositional environments were mainly fluvial,
aeolian and lacustrine (Karnkowski 1999).
The extent of the Variscan orogen in Poland is still uncertain
(Pozaryski et al. 1992; Dadlez et al. 1994), mainly because of
the lack of reliable data documenting the regional extent and
styles of orogenic deformation. However, recent studies of well
data have provided new clues to the styles of deformation within
the external Variscides and the possible location of the front,
including the possibly significant role of strike-slip movements
within the frontal part of the orogenic zone and source area for
foredeep basin infill (Aleksandrowski et al. 2003; Jaworowski
2002; Mazur et al. 2003). Extension, subsidence and basin development, subsequent to Variscan orogenic activity, was parallel to
the SW margin of the East European Craton and the Tornquist–
Teisseyre Zone (Figs 7 and 8).
The Polish Basin was formed in the Late Palaeozoic to the
north of the Variscan orogen (Dadlez 1997, 2006) and belonged
to a series of sedimentary basins that developed around the
margins of the East European Craton (Nikishin et al. 1996). The
main depocentre of the basin is termed the Mid-Polish Trough
(MPT), and this was active during the basin’s evolution, beginning
in the Permian and extending to the Mesozoic, as a zone of
maximum subsidence with almost uninterupted sedimentation
(Dadlez 1997; Kutek 2001). Indeed, the axial part of the MPT
was also the locus of the Late Cretaceous inversion of the Polish
Basin (see Krywiec 2002a and Lamarche et al. 1999 for details).
The present base of the Permian in the MPT ranges from 3 to
8 km (Dadlez 1998; Znosko 1999).
The Polish Basin developed on highly heterogeneous basement.
The MPT developed along the Tornquist– Teisseyre Zone, which
marks the boundary between the East European Craton and
the Palaeozoic Platform (Guterch et al. 1983; Grad et al. 1999).
Integration of industry seismic reflection data with gravity and
magnetic data has shown that the location of the NE MPT
margin was strongly and directly controlled by the southwestern
margin of the East European Craton (Krzywiec & Wybraniec
2003; Krzywiec 2004; Lamarche et al. 2003a). Indeed, the form
of the MPT is considered to have partly resulted from rheological
boundaries within the Trans-European Suture Zone (Stephenson
et al. 2003). Furthermore, the MPT is segmented along strike
into (from NW to SE) the peri-Baltic, Pomeranian, Kuiavian and
365
lithospheric cooling and contraction following the cessation of the
Autunian wrench faulting and magmatism.
Northern Permian Basin: Denmark. The Rotliegend sediments of
Fig. 9. Structure on the base Upper Rotliegend (below sea level) with
the geographical segments of the Polish Basin. Contour lines are at
500 m interval. The position of the cross-section shown in Figure 10
is indicated (after Stephenson et al. 2003).
Malopolska segments (Dadlez 1998), which are related to
crustal-scale fracture zones (Królikowski & Petecki 1995;
Grad et al. 1999; Lamarche et al. 2003b; Lamarche & SchechWenderoth, M. 2003). Along the axial zone of the basin the sedimentary succession is 3 –7 km thick (Marek & Pajchlowa 1997)
including about 1500 m of Zechstein salt (Tarka 1994). Deposition
of Zechstein evaporites during the Permian was limited to the SE
by the Grójec Fault, which was also active during the Carboniferous (Zelichowski et al. 1983).
Northern Permian Basin. The Northern Permian Basin (NPB) is an
elongate east–west-oriented basin extending across the Central
North Sea from Scotland to Northern Denmark (Fig. 6). The
basin is bounded to the south by the fragmented Mid-North Sea–
Ringkøbing–Fyn High. To the east and NE the Sorgenfrei–
Tornquist Zone separates it from the stable Fennoscandian Shield.
As a result of Mesozoic overprinting, the outlines, geometry and
facies patterns of the NPB are less well understood that those of
the SPB (Ziegler 1990), largely because of the lack of outcrops and
limited drilling (Stemmerik et al. 2000). To the west the basin rests
on Devonian clastic deposits whereas to the east it overlies
Caledonian basement, on Lower Palaeozoic sediments preserved in
the Caledonian foreland, and on Precambrian crystalline rocks of
the Fennoscandian Shield (Hospers et al. 1985). Ziegler (1990) has
noted that it is assumed that the NPB also evolved in response to
Denmark were deposited in the Danish Central Graben and the
Danish– Norwegian Basin, located north of the fragmented
Mid-North Sea – Ringkøbing–Fyn High, and form part of the succession deposited within the Northern Permian Basin (Fig. 11;
Stemmerik et al. 2000). Modelling suggests that early Permian
extension resulted in crustal thinning of up to c. 1.35 and was probably related to a major Late Carboniferous –Early Permian heating
and extension phase (Frederiksen et al. 2001).
In the Danish North Sea the Rotliegend is subdivided into two
formations. The lowermost Karl Formation is areally widespread
and is defined to include the syntectonic volcanic, volcaniclastic
and sedimentary fill of the Permian half-grabens. The succession
is dominated by volcanic rocks of alkaline basaltic composition
and volcaniclasti sediments. With the exception of some isolated
rhyolite flows (e.g. Horn Graben Rhyolite Member), the succession differs in composition from the more acid volcanic rocks
that characterize the Lower Rotliegend of the SPB (Aghabawa
1993). The rhyolitic flows were extruded as silicic lava, which
solidified as glass and was subsequently devitrified (Aghabawa
1993). The majority of the volcanic rocks are basic in composition,
described as alkaline basalts, hawaiites and mugearites (Aghabawa
1993). Associated sediments are mainly volcaniclastic sandstones
and conglomerates that were probably deposited in alluvial and
fluvial settings. The overlying Auk Formation represents the postrift succession, comprising sandstones, pebbly sandstones and
conglomerates with thin interbeds of silty mudstone, which was
probably deposited in aeolian, fluvial and sabkha –lacustrine
settings.
Oslo Rift. The Permo-Carboniferous Oslo Rift extends northwards
from the Sorgenfrei– Tornquist Zone and comprises three opposing half-graben segments: the southern Skagerrak Graben,
the central Vestfold Graben and the northern Akershus Graben
(Figs 12 and 13). The basin formed as a result of dextral
transtension, which led to the development of a purely extensional
regime in the Oslo Rift system and the culmination of rifting
and peak magmatic activity in the Oslo Graben (Olaussen et al.
1994; Heeremans et al. 1996; Torsvik et al. 1998). The onshore
part of the present-day Oslo Rift consists of a c. 400 km long
and c. 35– 65 km wide graben containing large volumes of
rift-related extrusive and intrusive rocks and minor amounts of
rift-related sedimentary rocks (Neumann et al. 1995, 2002).
The driving force behind the initiation of rifting in the Oslo
region is difficult to ascertain in terms of active (plume-related)
and/or passive (lithospheric stretching) end-members (Kirstein
et al. 2002). There is no evidence of crustal doming prior to
magmatism (Olaussen 1981). Late Carboniferous pre-rift sediments (i.e. Asker Group) were deposited close to sea level and
contain evidence of episodic marine incursions (Olaussen 1981).
Rifting in the Oslo region appears to have begun in Late Carboniferous time with the formation of a shallow depression close to
Fig. 10. Geological cross-section based on
well and seismic data through the Polish
Trough (location shown in Fig. 9). The
approximate location of the Trans-European
Suture Zone (TESZ) is indicated (after
Stephenson et al. 2003).
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T. MC CANN ET AL.
Fig. 11. Simplified SE –NW cross-section from the
eastern North Sea–Ringkøbing –Fyn High to the
Danish–Norwegian Basin showing the distribution
of Rotliegend sediments. (Note the onlap of
the Auk Formation towards the Ringkøbing–Fyn
High (after Stemmerik et al. 2000).
Fig. 12. Regional structural map of the
Oslo Rift and adjacent areas based on
maps by Ramsberg et al. (1977), Falkum
& Petersen (1980), Buer (1990) and
Ro et al. (1990b). FZ, fracture zone;
GS, graben. The location of the
seismic profile OG-7 (shown in Fig. 13) is
indicated.
367
Fig. 13. Seismic line drawing of profile
OG-7 across the Skagerrak Graben (after Ro
et al. 1990b). (See Fig. 12 for location.)
TWT, two-way travel time.
sea level, in which deposited sediments are sealed by the firsterupted lava flows (i.e. B1 basalts). This suggests that rifting
began with a period of lithospheric stretching and thinning prior
to the onset of the main magmatic phase. Coeval intrusion of
sills of intermediate to felsic character is interpreted in terms
of compressional activity prior to the main phase of extension
(Sundvoll et al. 1992). The initial magmatic phase (involving
basalt extrusion) and vertical movement along NNW– SSE- to
north– south-oriented faults appears to have been contemporaneous in response to ENE – WSW- to east– west-oriented crustal
extension (Heeremans et al. 1996). The main phase of rifting
and volcanism involved large displacements (up to 3 km) along
some of the basin-bounding faults (i.e. the Oslofjorden fault;
Neumann et al. 2004, and references therein). This phase also
coincided with the development of the Oslo Graben. Subsequent
caldera formation was accompanied by changes in the magma
chemistry and the increasing dominance of intrusive activity
(Neumann et al. 1995), associated with a change in the orientation
of structural deformation to more NE– SW, NW – SE and
west – east.
The Permian Oslo Rift was located within a broad intracratonic
basin during early Palaeozoic time (Ramberg 1976) containing an
up to c. 4 km thick sedimentary sequence of Cambrian to late
Silurian age. This succession is locally unconformably overlain
by Upper Carboniferous and Lower Permian sedimentary rocks
which underlie the Permian lavas (Olaussen 1981). Sedimentary
rocks also occur as thin layers (,10 m) between lava flows.
Significant thicknesses of fanglomerates are preserved close to
the Oslofjorden fault zone in the SE part of the Oslo Graben.
Here the clasts mainly comprise lavas and these deposits are
indicative of synvolcanic tectonic activity (see Larsen et al.
1978, for more details). The thickness of Permian sediments,
however, is secondary to that of the magmatic succession,
attaining (in the Skagerrak Graben) a maximum thickness of
only c. 1 km (Ro et al. 1990a).
The North Swiss Permocarboniferous Basin (NSPB) is found
within the subsurface of northern Switzerland and was first discovered early in the 1980s (e.g. Matter et al. 1987; Fig. 15).
Initially, it was called the Constance– Frick Trough
(Konstanz– Frick Trog, e.g. Laubscher 1987) and later, as its
continuation further to the west was recognized, the NSPB
(i.e. Nordschweizer Permokarbon Trog). Formed within crystalline basement, the basin is 10– 12 km wide and filled with
.1500 m of continental clastic deposits (see Blüm 1989, for
overview). Supply was local (Matter 1987; Blüm 1989). Carboniferous deposits have been drilled only at Weiach and appear to
be restricted to a narrow, graben-like structure. The NSPB is
disrupted by a series of NW –SE-trending faults (Fig. 15). At
Weiach 572 m of Carboniferous sediments were drilled, but
the basement was not reached (Fig. 16). Microfloral remains
are dated as Stephanian (Hochuli 1985). Later radiometric
dating of zircons from ash layers confirmed the age (303 Ma
in the middle of the drilled Carboniferous deposits and
298 Ma at the top; Schaltegger 1997a).
The palaeogeographical position within the Variscides, the
varying sediment thickness, and the dominance of crystalline
basement clasts derived from local sources suggest a pull-apart
origin of the NSPB (e.g. Matter 1987; Blüm 1989). The basin is
interpreted as a series of en echelon pull-apart basins (e.g.
Diebold et al. 1992) that formed in relation to late Variscan
wrench tectonics.
Areally, the Permian deposits occupy a significantly wider
area than those of the Carboniferous. It would appear that
only Early Permian deposits were accumulated, but biostratigraphical dating using palynomorphs (Hochuli 1985) may be
subject to some uncertainty because of palaeotopographic
effects (e.g. Becq-Giraudon 1993). Except for Weiach, the
Permian deposits rest on crystalline basement. Permian deposits
were mainly correlated according to lithology (e.g. Blüm 1989)
and are all continental, ranging from lacustrine to alluvial fan
units.
Basins occurring within the Variscan orogen
Massif Central. The Variscan Massif Central shows the characteristic tectonometamorphic evolution of classic collisional belts,
with significant horizontal thrusting and progressive crustal
thickening (Echtler & Malavielle 1990). The Montagne Noire
forms the external, southernmost segment of the Massif Central,
which may be subdivided into a crystalline metamorphic core
(Axial Zone) and a tectonically overlying upper unit of strongly
deformed sedimentary rocks (Figs 17 and 18). Partly interfolded
sediments of late Viséan and Namurian age occur along the southeastern border of the thrust belt (Engel et al. 1982). Up succession,
the unit is increasingly dominated by synorogenic coarse-grained
turbidites and a progressively chaotic set of kilometre-sized olistolithic slabs, which are syntectonic (Engel et al. 1982). Deposition
of these units was followed by a period of Westphalian –
Stephanian extension.
Permo-Carboniferous Basin, Switzerland. The Late Palaeozoic
basins in northern Switzerland formed after the main Variscan
orogeny (dated in Switzerland as pre-late Westphalian), but in
the French Alps, post-orogenic sedimentation commenced in the
Late Namurian (Trümpy 1980). Because all these basins formed
within the same time span and in close proximity to one
another, they exhibit striking similarities, in that the basins, predominantly grabens or graben-like structures, tend to be small,
elongate and mainly trending SW – NE (Fig. 14). Furthermore,
Carboniferous deposits are restricted to a narrow trough whereas
the basins are filled with continental clastic deposits, mainly
eroded from the crystalline basement, and containing volcanic
and/or volcaniclastic material.
368
T. MC CANN ET AL.
Fig. 14. Permo-Carboniferous basins in
Switzerland north of the Penninic front.
Basins close to the Penninic front (BF,
Bifertengrätli; GV, ‘Glarner Verrucano
Basin’; SD, Salval–Dorénez Basin) are
exposed (after Trümpy & Dössegger 1972).
Basins in the subsurface (NSPB) after Boigk
& Schöneich (1974), Bachmann et al.
(1987), Meier (1994), Thury et al. (1994) and
Wetzel & Allia (2003).
Autunian
areas, from
France (e.g.
Paris Basin
deposits in France are found in a variety of
the classic region of the Massif Central, to NW
Carentan Basin, Chateauneuf & Farjanel 1989), the
(Bouas 1987), the northern part of the Pyrenees
(Bixel & Lucas 1983) and SE France (Chateauneuf & Farjanel
1989) (Figs 17 and 19). Stephanian– Autunian basin formation
in the Massif Central is attributed to the collapse of thickened
crust, aided by a NE– SW-oriented extensional stress field.
Fig. 15. North Swiss Permo-Carboniferous Basin in map view. Compiled from data by Boigk & Schöneich (1974), Diebold et al. (1992), Meier (1994), Thury et al.
(1994) and Wetzel & Allia (2003). In addition, data from boreholes have been used (see McCann et al. 2007, for details). URG, Upper Rhine Graben; RF, Rhine Fault
(Rhenish Lineanent).
Fig. 16. North Swiss Permocarboniferous Basin cross-section close to
the well at Weiach (see Fig. 15), and based on interpreted seismic records
(Diebold et al. 1992; after Schaltegger 1997a,b). The boundary between the
lower and upper basin fill units roughly corresponds to the
Carboniferous–Permian boundary.
In the Massif Central the most notable event marking the
Stephanian was the appearance of numerous small limnic coalbearing basins, closely related to a network of regional faults
(related to Variscan tectonics) that appeared or were reactivated
at this time (Arthaud & Matte 1975). From the end of the
Westphalian to Stephanian B north– south compression continued
in the Massif Central, but the Westphalian regional ductile
shear faults progressively gave way to fracture deformation,
responsible for the formation and evolution of the Stephanian
limnic basins. The anti-clockwise rotation of north– south-
Fig. 17. Permian basins of France (after Cassinis et al. 1995).
369
Fig. 18. (a) Schematic cross-section of the Montagne Noire metamorphic core complex from the southern part of the Massif Central, France. The lower plate (L.P.) comprises high- to medium-grade metamorphic,
gneissose and mica schist formations whereas the upper plate (U.P.) consists of undifferentiated low-grade or non-metamorphic Cambrian to Early Carboniferous sediments. (b) The main Permo-Carboniferous
basin in the French Massif Central (after Echtler & Malavielle 1990).
370
T. MC CANN ET AL.
oriented compression to an east– west direction was responsible
for the evolution of these regional strike-slip faults and consequently of the associated basins. At the end of the Stephanian,
the final stage of east – west-oriented compression interrupted sedimentation and caused intense deformation of the basins located on
the approximately north –south-striking faults. In the Pyrenees,
the North Pyrenean Fault, which borders the axial zone,
probably underwent dextral strike-slip with horizontal displacement estimated at 150 km (Arthaud & Matte 1975). Subsequent
to the period of Stephanian compression at the close of the
Variscan orogeny, a new period was initiated during that
intracontinental basins formed, the thickness and extent of
whose deposits differed from those that accumulated during the
Stephanian. The Stephanian compressive basins of the Massif
Central are long and narrow with varying thicknesses of sediment,
whereas the extensional Permian basins contain successions up to
several thousands of metres thick and are much more extensive,
occupying the sites of the future large Mesozoic basins.
In France the Permian includes only continental deposits. The
succession is generally subdivided into Autunian and ‘Saxonian’
strata (both of Early Permian age) and Thuringian deposits (considered to be more or less equivalent to the Late Permian)
(Cassinis et al. 1995). The typical Autunian was defined in the
Autun Basin, located to the north of the Massif Central. These
units comprise lacustrine calcareous and bituminous shales,
coarse fluvial deposits and some volcanic ash, which overlie
folded Stephanian beds (Cassinis et al. 1995). The many
Stephanian– Autunian basins in the Massif Central contain coal
measures and alluvial, fluviatile and lacustrine clastic sediments
deposited in an active tectonic environment; magmatic rocks are
present in only small volumes (Legrand et al. 1994; Djarar et al.
1996; Allemand et al. 1997). Saxonian deposits include those
that overlie the Autunian, and these units are generally separated
by an unconformity attributed to a period of intra-Permian
(i.e. ‘Saalian’) deformation. The expression of this varies from
changes in sedimentation to an actual angular unconformity
(Cassinis et al. 1995).
Permian-age sediments were deposited on a variety of strata that
were deformed and metamorphosed during the Variscan orogeny,
and in many basins sedimentation continued into late Permian
times (Châteauneuf & Farjanel 1989). Sedimentation was controlled by synsedimentary basin margin faults, and faulting
caused considerable palaeorelief and the formation of basement
horsts that were important sources of sediment, especially alluvial
cones. Some of the larger basins began as smaller sub-basins with
independent drainage systems that amalgamated during accelerated subsidence (e.g. the Cévennes Basin, Djarar et al. 1996).
Those basins located along approximately east-trending faults
(e.g. the Saint-Etienne Basin) tend to have pull-apart geometries
that are related to strike-slip movements along the basin margin
faults. Those basins located along approximately north-trending
faults tend to be half-grabens, some of which are strongly asymmetrical (Faure 1995; Mattauer & Matte 1998). Furthermore,
the extent of soft-sediment deformation from the Saint-Affrique
and Cévennes basins testifies to the syndepositional, extensional
to transtensional character of the basin margin faults (Legrand
et al. 1994; Djarar et al. 1996).
In the southern Massif Central distinct sequences are present
within the Autunian deposits. For example, in the Lodève Basin
(Laversanne 1978) sedimentary sequences between 8 and 15 m
thick occur within a unit that is c. 800 m in thickness (Fig. 19).
These sequences, which are continental in origin, comprise
fluviatile sandstones in the lower part, sandstones and bituminous
dolomites of palustrine or lagoonal origin in the middle part, and
calcareous silty floodplain mudstones and siltstones in the upper
part (which also contains evaporitic precipitates). Each basin
developed in a general tensional context. In many cases there is
no appreciable lithological change through the entire Upper
Permian sequence and, because of a lack of data, many workers
prefer to classify such rocks as Saxonian– Thuringian units. In
the Lodève Basin the Upper Group is present.
Saar – Nahe Basin. This Permo-Carboniferous basin extends from
SW Germany into France and is filled with exclusively continental
sediments (Fig. 20). The basin has a half-graben geometry, being
bordered to the north by the south-dipping Hunsrück Boundary
Fault (HBF), which, according to Henk (1993) is a detachment
soling out at mid-crustal levels (at a depth of c. 16 km). Transtensional subsidence of the partly inverted Saar – Nahe Basin, which
contains up to 5.6 km of Permo-Carboniferous clastic deposits,
accounts for a stretching factor of .1.36. Contemporaneous
extrusive activity reflects destabilization of its lithospheric
system (Ziegler et al. 2006). The evolution of the Saar – Nahe
Basin is closely related to the complex kinematics of the HBF,
which is part of a prominent suture zone separating two of the
main tectonostratigraphic units of the Variscan fold belt: the
Rheno-Hercynian and the Saxo-Thuringian zones. The sedimentary infill of the Saar – Nahe Basin is dominated by continental
clastic deposits (Schäfer 1989; Schäfer & Korsch 1998) with a
significant thickness of contemporaneous volcanic rocks.
Basin initiation occurred in the latest Namurian or possibly earliest Westphalian, and the absence of older Namurian sediments
suggests that this was an elevated area from the late Viséan to
the late Namurian. This period of non-deposition reflects Variscan
compression, uplift, exhumation and cooling, as indicated by
320– 335 Ma 40Ar/39Ar cooling ages from Mid-German Crystalline Rise (MGCR) basement rocks to the south (i.e. Odenwald
and Erzgebirge, Werner & Lippolt 2000; Schubert & Lippolt
2000). There is an angular unconformity between the Westphalian
D and Stephanian A, indicative of a reorganization of fault kinematics in the area. Major uplift in the southeastern part of the
basin elevated the MGCR basement to provide a new sediment
source (Korsch & Schäfer 1995). The volcanic rocks are related
to the crustal stretching that commenced in the earliest Westphalian, where a regional, rather than local, mechanism was responsible. The cause of this was probably dextral translation between
Laurussia and Gondwana, which was about to collide with
eastern to southeastern North America. Rotation and translation
led to the reactivation of old lineaments (such as the northern
boundary of the Saar – Nahe Basin, the Hünsrück Fault) and the
establishment of large-scale, new fault systems in Europe (e.g.
Arthaud & Matte 1977; Ziegler 1990). Magmatism led to a
thermal anomaly, leading to a subsequent phase of thermal
relaxation and crustal re-equilibration.
The west –east extensional stress regime that dominated during
the post-Westphalian basin formation (Stollhofen 1998) was
oblique to the pre-existing fault pattern; the common slip direction
was maintained by transtensional strike-slip movements on the
transfer faults and oblique-slip motion on the normal faults
(Stollhofen et al. 1999). An angular unconformity, indicative of
a reorganization of fault kinematics in the area, underlies the
synrift megasequence. The Stephanian prevolcanic synrift
sequence has a thickness of 3.8– 4.7 km and was deposited over
a period of c. 14 Ma. It comprises lacustrine, fluvio-deltaic and
fluvial sediments with minor limestones, coals and pyroclastic
fallout deposits (these last were derived from sources outside of
the basin; Stollhofen et al. 1999). This is overlain by the Lower
Permian volcanic synrift sequence (1.1 km thick, deposited
over c. 4 Ma); during this phase widespread bimodal calc-alkaline
magmatism occurred with subvolcanic intrusion of rhyolitic–
dacitic domes and basaltic to andesitic sills and dykes (von Seckendorff et al. 2004b). Magma generation can be attributed to
underplating and intrusion of pulses of mantle-derived melts into
the crust inducing partial melting, or perhaps a distal subduction
zone (Schmidberger & Hegner 1999). The extrusive rocks are
interbedded with fluvial sediments and minor lacustrine units
(Stollhofen 1994).
371
Fig. 19. Cross-sections through Permian basins in Central
France: Lodève Basin, south of the Massif Central (after
Cassinis et al. 1995), Autun Basin and Blanzy– Le Creusot Basin
(after Blès et al. 1989). (Locations are shown in Fig. 17.)
Fig. 20. Geological interpretation of
DEKORP 1C and 9N showing the
half-graben form of the Saar–Nahe Basin,
Germany (after Henk 1993).
372
T. MC CANN ET AL.
Iberian basins. The Iberian microplate was affected during the
Permian by the final stages of Variscan activity and by the early
Atlantic rifting (Fig. 21). These tectonic disturbances led to the
development of intracratonic basins. Two phases of Variscan
deformation have been described in northern Spain, an early
phase consisting of two east– west-oriented compressional
events in the period from the Namurian to the Stephanian that
resulted in arc-parallel folds and thrusts, and a later north –
south-oriented phase (Sakmarian – Early Permian) that marked
the final stage of Variscan deformation in northern Iberia and
resulted in plunging fold axes of the early arc-parallel folds
(Weil et al. 2000). Subsequent rifting episodes (e.g. Early
Permian– Triassic) resulted in the formation of basins with orientations ranging from NE– SW to east –west (Jabaloy et al. 2002).
Additionally, magmatic events have been recorded. Well-dated
volcanic rocks and coeval granites crop out in the Pyrenees, the
Iberian Range and the central part of Iberia (Muñoz et al. 1986;
Sopeña et al. 1988). The basins are filled with red beds interdigitated with volcanic and volcaniclastic rocks. Thicknesses are variable and can be as much as 2000 m. These units appear as isolated
outcrops of generally small extent because they were deposited in
complex graben systems. Stephanian deformation involved NE–
SW compression and horizontal extension orthogonal to the compression direction. In the Early Permian there was major volcanic
activity. The fracture patterns associated with this activity in the
central part of Iberia indicate extensional deformation in this
region (de Vicente et al. 1986; González-Casado et al. 1996;
van Wees et al. 1998) and probably also in the Pyrenees. Structural
data are scarce, but appear to indicate that the central area was
characterized by a NNE –SSW- to NNW– SSE-oriented extension
(Arche & López-Gómez 1996; González-Casado et al. 1996),
although transcurrent deformation with a NE – SW orientation
continued in the SW of the Iberian Peninsula (Herraiz et al. 1996).
There are five regions in Spain where the Permian is well
known: the Cantabrian Mountains, the Pyrenees, the Central
System, the southern margin of Iberia and the Iberian Range
(Cassinis et al. 1995). The succession may be broadly subdivided
into a Lower and an Upper Group. The Lower Group, associated
with the formation of small basins controlled by strike-slip
faults, comprises one or more tectonosedimentary units that
unconformably overlie the Stephanian or older Palaeozoic rocks.
Its thickness ranges between 200 and 2000 m. Widespread andesitic and rhyolitic volcanism of calc-alkaline type is associated
with this group (Lago et al. 2004). The overlying Upper Group
sediments were deposited in an extensional cycle and are areally
more widespread. There is a marked unconformity between the
two groups.
The Permo-Carboniferous succession of the Central Pyrenees
has been considered to have originated from strike-slip dynamics
that developed during a compressional episode at the end of the
Variscan orogeny as established from facies analysis (e.g. Marti
1986, 1991; Besly & Collinson 1991) and from regional palinspastic reconstructions (e.g. Muñoz 1992; Casas et al. 1989). In
the Pyrenees, the oldest sediments comprise breccias, sandstones
and coal beds, and contain Stephanian flora (Fig. 22). These are
overlain by a widespread volcanic unit comprising andesitic pyroclastic deposits and volcaniclastic rocks (see Cassinis et al.
1995; Lago et al. 2004, for details), which is overlain by conglomerates and sandstones that contain a Stephanian –Autunian
flora. In the Anayet Massif, for example, there is an extensive
Fig. 21. Tectonic sketch map of the
Iberian Peninsula (after Van Wees et al.
1998). Inset map provides a
reconstruction of the basin-bounding faults
of the Iberian Basin (after Arche &
López-Gómez 1996).
373
Fig. 22. Upper Carboniferous and
Permian of the Pyrenees and Iberian
Ranges in Spain (after Cassinis et al.
1995). AU, Autunian macro- or
microflora; ST, Stephanian macro- or
microflora; TH, Thuringian microflora.
sedimentary record of Stephanian –Permian-age deposits. Here,
the succession is composed of four lithological units representing
a stratigraphic transition from Stephanian to Permian. In terms of
the sediments, these units represent a transition from humid to
arid climatic conditions (with fluvial sediments) (Gisbert 1983).
Contemporaneous with sedimentation, there were several intrusive and extrusive magmatic episodes favoured by the transtensional tectonic regime, which was also responsible for the
development of small listric-fault-bounded basins containing up
to 3 km of volcanoclastic and detrital rocks (Bixel et al. 1996).
These mixed magmatic– sedimentary successions were deposited
in isolated basins (five of which have thus far been identified;
Martı́ 1996).
The only outcrops of marine deposits of Stephanian and
Permian age to be found in Western Europe (with the exception
of the Eastern Alps and Sicily) occur within the Iberian Massif
in western Spain and Portugal (Martı́nez Garcı́a 1990). Compressional tectonics characterizes the evolution of the Stephanian
basins, whereas the Permian (and Stephanian C) shows a tensional
environment that reveals plate-tectonic conditions suggestive of
failed rifting. Two units, corresponding to different basins, have
been distinguished in the Permian succession of the Cantabrian
and Palential zones by Martı́nez Garcı́a (1983). The lowermost
of these comprises alternating clastic sediments, tuffs, volcanic
agglomerates, lava flows and shallow marine limestones. Overlying these are volcanic tuffs and ash up to 900 m thick and
with an increasingly acid character. Thick conglomerate units
attest to syndepositional tectonic activity. The unconformably
overlying succession comprises predominantly continental
clastic sediments.
The Permian succession of the present-day Iberian Range,
eastern Spain, is very well exposed (Figs 22 and 23). These sediments are of continental origin and filled isolated, small basins
that were initiated during latest Carboniferous –early Permian
times (Sopeña et al. 1988) and that evolved into a single
basin (i.e. the Iberian Basin) during late Permian times. This
period was one of tectonic readjustment of plates by transtensional faulting in the Chedabucto– Gibraltar and Bay of Biscay
areas and the development of a conjugate wrench zone, originating later in the Iberian Basin, following an ancient suture (NE –
SE-oriented) running across the microplate (Arthaud & Matte
1977; Salas & Casas 1993; López-Gómez et al. 2002). During
early Permian (Autunian) times, a series of intermontane
basins were filled in by alluvial fans, slope breccias and lacustrine deposits associated in their lower part with volcaniclastic
rocks of calc-alkaline affinities (Lago et al. 1992). Lacustrine
deposits at the top of the volcaniclastic succession record a
change from freshwater to saline lakes, indicating a progressive
aridity in Iberia. These early Permian sediments were mostly
unconformably deposited on rocks of early Palaeozoic age, but
also upon late Carboniferous (Stephanian B– C) sediments.
These comprise sandstones and coal-bearing successions that
filled the Henarejos intermontane small basin contemporaneously with those of the Cantabrian and Palential zones and
are related to the final, extensional collapse of the Variscan
orogeny, which also resulted in monzogranitic magmatism and
uplift in Central Iberia.
374
T. MC CANN ET AL.
During the late Permian, red bed successions up to 300 m thick
accumulated in grabens and half-grabens that formed during
a period of widespread extension within the Iberian Basin
(López-Gómez et al. 2002). The overall development shows a
fining-upwards trend in most of the basins, beginning with transverse alluvial fan deposits and followed by longitudinal braided
river system deposits. There was no volcanic activity during this
period in the Iberian Basin, where basin boundary faults have
shallow listric geometries at depths of only 13– 14 km.
The outcrops in the Iberian Range are the most widespread and
the best studied. The Molina de Aragon series (Ramos 1979;
Ramos & Doubinger 1979) is the most characteristic and the
one that is best correlated with the Permian of Central Europe
(Virgili et al. 1976, 1983; Fig. 23). The lowermost units
unconformably overlie Lower Palaeozoic rocks and consist of
breccias, volcano-sedimentary rocks and thin volcanic intercalations. These are overlain by 300 m of lacustrine and fluvial sandstones and black shales. The succession ends with lacustrine,
siliceous dolomites unconformably overlain by the Upper Group
(Permian). The onset of sedimentation of the Upper Group is
marked by an important and widespread unconformity reflecting
an important change in tectonic evolution.
Geophysical investigations
Subsequent to the European Geotraverse (EGT; Blundell et al.
1992), the increase in the number and quality of seismic data
Fig. 23. The Pálmaces– Riba and Molina
de Aragon sections from the Castilian
branch of the Iberian Ranges. AS,
Arandilla sandstones; HGC, Hoz de Gallo
sandstone; MB, Monstsoro beds; PB,
Prados beds; PLC, Pálmaces lower
conglomerates; PMS, Pálmaces mudstone
and sandstone; PS, Pálmaces sandstone;
PUC, Pálmaces upper conglomerates; RGS,
Rillo de Gallo sandstone; RSC, Riba de
Santiuste conglomerates; RSS, Riba de
Santiuste sandstones; VSC,
volcano-sedimentary complex (after
López-Gómez et al. 2002).
has greatly improved our understanding of the nature of the crust
that underlies many of these Permo-Carboniferous-age basins,
allowing us to ascertain the nature and relevance of any preexisting structures (e.g. the Tornquist– Teisseyre Zone and the
Sorgenfrei– Tornquist Zone) to the formation of the post-Variscan
basins (e.g. Polish Trough and Oslo Rift, respectively). The following section will outline some of the main advances and
results that have helped to improve our understanding of the evolution of these basins.
Germany
A series of DEKORP profiles were shot in the 1980s to image the
western margins of the Variscan deformation front close to the
German border with Belgium (Meissner & Bortfeld 1990). Of
greater interest, however, was the series of profiles shot in the
1990s, which imaged not only the North German Basin but also
the major structures to the north and south of it. Prior to obtaining
the deep seismic profiles, the precise nature of the North German
Basin (i.e. whether it was an extensional basin or not) had been
much discussed in the literature. As is clearly shown on the
DEKORP 9601 profile (DEKORP-BASIN Research Group
1999), the basin is intracratonic, and was initiated above a
region that had undergone significant rupturing as a result of the
wrench tectonics at the Permo-Carboniferous transition.
However, there is little evidence of these movements on the
seismic profiles. Instead, we have a clear image of the Moho (continuous across the entire profile) and the northern (where Avalonia
is obducted onto Baltica) and southern (with clear evidence of
Variscan thrusting) margins (Fig. 24). The depth to the Moho
is 30 km beneath the entire basin, except at the margins, where
it extends to 35 km. More detailed profiles are required to reveal
any evidence of Stephanian –Rotliegend tectonic activity.
However, tectonic activity in the form of continuous transpression
is evidenced by the diverse angular unconformities visible
between the Westphalian and Permian deposits on 3D seismic
profiles and in drill core (McCann 1998a).
In the Saar – Nahe Basin two deep seismic profiles (DEKORP
1C and 9 N) crossed parts of the basin (DEKORP Research
Group 1991; Korsch & Schäfer 1991). The southernmost part of
the basin was also imaged by the ECORS-DEKORP 9S profile
(Brun et al. 1991). These profiles reveal the asymmetric geometry
of the basin bounded by the South Hunsrück Fault, which has
been interpreted as having either a subvertical (Korsch &
Schäfer 1995) or listric (Henk 1993) geometry, although the
latter interpretation is the favoured one. The upper crust beneath
the basin shows a segmentation into three distinctively different
reflectivity patterns. The uppermost highly reflective package
represents the pre-rift sediments and Permo-Carboniferous basin
fill. The underlying wedge-shaped unit lacks major reflectors
and is interpreted as crystalline basement rocks of the
northern Saxo-Thuringian Zone (i.e. Mid-German Crystalline
Rise). Beneath is a thick, highly reflective zone, which may
represent remnants of sediments and oceanic crust from the
375
Rheno-Hercynian Ocean (Behr et al. 1984; Franke & Oncken
1990), or an anastomosing pattern of shear zones and duplex
structures (DEKORP Research Group 1991).
Heat-flow modelling has been carried out in a number of areas,
generally in conjunction with other forms of modelling. In NE
Germany, for example, Ondrak et al. (1997) integrated the structural models of Scheck (1997) to produce a regional thermal
model that allowed the determination of temperature distributions
and a depth-dependent estimation of the local heat-flow conditions. Temperature gradients within the model matched the
regional trends of heat-flow distributions, although the pattern is
more complex. There is a clear relationship between temperature
and the Zechstein salt layer, where high temperatures are related
to salt margins and regions where sediments with low thermal
conductivities cause local elevation of the isotherms.
Poland
The crust of central and northern Poland has been extensively
studied by deep refraction and wide-angle reflection seismic
studies. Older profiles have recently been reprocessed and reinterpreted (see Grad et al. 2003, for details) and integrated with the
new high-quality LT7 profile (Guterch et al. 1994), the TTZ
profile (Grad et al. 1999) and profiles from the POLONAISE 97
experiment (Guterch et al. 1999; Jensen et al. 1999; Janik et al.
2002). All of these seismic data have provided new information
on the deep structure of the transitional area between the East
European Craton and the Palaeozoic Platform in central and northern Poland. The depth to the Moho in this region is 32– 39 km
beneath the two-layered Palaeozoic Platform and 43– 45 km
beneath the three-layered crystalline crust of the East European
(Precambrian) Craton. The Trans-European Suture Zone (TESZ)
located between these two main crustal domains is characterized
by the presence of thick (up to 20 km) complexes of low-velocity,
sedimentary, metamorphic or volcanic rocks, and by a lower crust
characterized by high velocities. In this region, the Moho is
located at intermediate depths in the NW (30– 33 km) and
deepens to the SE. According to some interpretations, the lower
crust within the TESZ area in central and northern Poland represents the attenuated Baltica margin underthrust towards the
SW beneath the Avalonian accretionary wedge (Grad et al. 2002).
The results of gravity modelling based on deep seismic refraction data suggest that the crustal structure of the Mid-Polish
Trough (MPT), especially of its pre-Zechstein substratum, is
more complex than suggested by deep refraction data (Królikowski & Petecki 1997, 2002; Petecki 2002). The upper mantle
beneath the TESZ is dense, and within the upper crust a highdensity body was also identified, as well as a complex transition
zone between the crust and the upper mantle. Long-wavelength
gravity anomalies are associated with lateral density variations
within the upper mantle and lower crust (Królikowski & Petecki
2002; Petecki 2002). Results of magnetic modelling suggest that
the average depth of magnetic sources within NW Poland is of
the order of 18 km, and could be correlated with the crystalline
Fig. 24. Interpreted line drawing of BASIN 9601 profile and its offshore extensions PQ2-009.1 and PQ2-005 showing the main tectonic and stratigraphic features.
376
T. MC CANN ET AL.
basement as evidenced by seismic data (Petecki 2002). Upper
crust fault zones revealed by gravity and magnetic data present
within the pre-Zechstein MPT basement played a significant role
during the Mesozoic evolution of the MPT (Dadlez 1997) as evidenced by regional interpretation of seismic reflection data (Krzywiec & Wybraniec 2003) and structural modelling (Lamarche
et al. 2003a).
Geothermal models in Poland have been more concerned with
examining the broad structural framework of the region, rather
than concentrating on the basin infill succession. For example,
Majorowicz et al. (2003) have shown that there is a sharp
change in heat-flow data between the East European Craton and
the Palaeozoic Platform. Numerical modelling of the crustal temperatures along several deep seismic profiles suggests extensive
crustal– mantle warming within the zone located between the
Sudete Mountains and the TESZ (Grad et al. 2003; Majorowicz
et al. 2003). This anomalous zone coincides with the location of
the Dolsk Fault and the Variscan Deformation Front. High heat
flow for the Palaeozoic Platform and related high temperatures
of the crust coincide with the reduced crustal thickness, whereas
the low heat flow of the East European Craton coincides with a
higher crustal thickness. Modelling results also suggest that high
mantle heat flow is required within the high heat-flow zone
located within the Palaeozoic Platform (characterized by the
100 km thermal lithospheric thickness) whereas cold crust and
cold mantle are typical of the East European Craton (characterized
by the 200 km thermal lithospheric thickness; Grad et al. 2003;
Majorowicz et al. 2003).
Thermal modelling of the Polish Basin performed by
Karnkowski (1999) suggested that the Rotliegend volcanism
began with high geothermal anomalies in the western part of
the basin. The anomalies are characterized by higher values
(100–150 mW m22) during the Late Permian –Early Triassic
interval, in relation to rifting in the Polish Basin. The Late Triassic
and Jurassic were a time of cooling, until the break-point in
thermal evolution of the Polish Basin at the Jurassic– Cretaceous
boundary as a result of uplift and erosion, after which the heat
inflow decreased.
Norway
The Oslo Graben and surrounding areas have been the subject of a
number of geophysical investigations (e.g. Ro et al. 1990a,b;
Kinck et al. 1991), which have revealed that there is a marked
crustal thinning beneath the graben, with the amount of thinning
increasing southwards. The depth to the Moho has been estimated
to range from 28 to 35 km beneath the Oslo rift system (Cassell
et al. 1983; Thybo 1997). Gravity data indicate the presence of
a dense 90 km wide body towards the base of the lower crust,
which has been interpreted as representing cumulates and gabbroic
rocks in a deep crustal magma chamber (Neumann et al. 1992).
Furthermore, gravity and seismic data imply a different crustal
structure along the rift from in the adjacent Precambrian terrane
(e.g. Wessel & Husebye 1987; Kinck et al. 1991). However,
more recent data have questioned these interpretations of gravity
data in the Oslo region (Ebbing et al. 2006). These data and modelling results show that the high-velocity layer present at the base
of the Oslo Graben is similar to the typical high-velocity layers
that are commonly found at the base of Proterozoic crusts
(Durrheim & Mooney 1991), suggesting that the idea of underplating beneath the Oslo Graben is a false one. It has been noted that
although Permian mafic intrusions in the crust may be present,
they do not reside in a hypothetical underplate, but most probably
in the middle crust (Ebbing et al. 2006).
Seismic refraction or wide-angle-reflection data have been used
to model variations in crustal thickness and structure in the Oslo
rift and adjacent parts of the Precambrian shield (e.g. Bungum
et al. 1971; Cassell et al. 1983; Ro et al. 1990a,b; Kinck et al.
1991). The subsequent Moho map shows that the crust is thicker
east of the Oslo Graben than west of it (Kinck et al. 1991;
Thybo 1997). The refraction or wide-angle reflection data of greatest relevance to the crustal structure of the centre of the Oslo
Graben segment include the profiles of Tryti & Sellevoll (1977)
(see Neumann et al. 1995, for details). Three distinct crustal
layers were noted, including a high-velocity (7.1 km s21) lower
crustal layer that had been interpreted as resulting from magmatic
underplating (see above) and/or a zone of magmatic cumulates
and residues (Neumann et al. 1995). Teleseismic studies carried
out along the 60th parallel (i.e. parallel to the modelled section
AA0 ) suggest that deepening of the base lithosphere occurs
abruptly from west to east in the Oslo region (Babuska et al.
1988). These finding are confirmed by a more recent teleseismic
study (Plomerová et al. 2001).
Modelling post-Variscan basin development
Although numerical methods have been applied to tectonic problems for over 30 years, relatively few have been concerned
with modelling the post-Variscan (i.e. Permian) basins of
Europe. Most Permian basins are buried at considerable depths
by Mesozoic and Cenozoic sediments, and their structure has
been obscured by younger tectonic events. As a result, the acquisition of accurate data for the modelling is very often difficult.
Nevertheless, the increasing collection of subsurface observations
(i.e. well and seismic reflection data) has allowed for the modelling of the post-Permian subsidence history in some areas (i.e.
Paris Basin, North Sea, NE German Basin, Danish Basin). In
other areas, which were uplifted in post-Permian times, direct
observation of the Permian basins (i.e. the Oslo Graben) allows
relevant data for thermo-mechanical modelling including
rheology and decompression melting to be acquired.
Models based on sediment succession –basin analysis
Models based on subsidence analyses have been the most widely
applied to reconstruct the pre-rift configuration and the rifting
history of Permian basins (e.g. Van Wees et al. 1998, 2000;
Frederiksen et al. 2001). These models are based upon the classical model of McKenzie (1978), and its numerous derivatives, in
which rapid synrift subsidence is followed by a more extended
phase of thermal subsidence. The pattern of basin subsidence is
strongly dependent on the initial configuration of the lithosphere
and the amount of stretching. One of the advantages of this
method is that it provides information on a previous rift event
from the analysis of post-rift sediments. Thus, direct observation
of the synrift sedimentary infill can be bypassed (see Allen &
Allen 2005, for further discussion on this topic).
A variety of basins both within and to the north of the broad
zone of Variscan-age deformation have been studied with this
method. Basins within the deformation zone include the Paris
Basin (e.g. Brunet & Le Pichon 1982; Prijac et al. 2000), the
Mid-Polish Trough (e.g. Dadlez et al. 1995), the Iberian Basin
(e.g. van Wees et al. 1998), and the Southern Permian Basin
(e.g. van Wees et al. 2000); those to the north of it include the
Danish Basin (e.g. Sørensen 1986; Frederiksen et al. 2001), the
offshore arm of the Oslo Rift (e.g. Pedersen et al. 1991), and
the Northern North Sea (e.g. Odinsen et al. 2000). In detail, the
results differ somewhat from one basin to another, reflecting the
complexity of the structure of the European lithosphere in
Permian times and the interplay between different subsidence
mechanisms. For example, Permian rifting and post-Permian subsidence in the Paris Basin can be explained by the collapse of the
Variscan mountain chain and the slow decay of the associated
thermal anomaly (Brunet & Le Pichon 1982; Prijac et al. 2000).
In the Variscan foreland such a mechanism cannot be invoked.
In contrast, Dadlez et al. (1995) and van Wees et al. (2000) argued
for strong destabilization of the lithosphere by late Variscan
wrenching involving deep fracturing and decompression melting
of the underlying mantle. This interpretation explains the large
thicknesses of Permo-Carboniferous magmatic rocks and post-rift
sediments that accumulated in the absence of major normal faulting of the crust. Modelling results concerning the precise age and
timing of the Permian rifting event differ from one study to
another. This reflects differences in data coverage between the
various studied areas, and differences in the methods used, but
also probably different responses of the lithosphere in terms of
its structure and past history. Whatever the precise mechanisms
associated with the Permian event, it is clear that subsidence
modelling studies agree on two major points: (1) latest Carboniferous– early Permian rifting (290–305 Ma) was a widespread and
dramatic event in Europe; (2) the thermal signature of the
Permian rifting was a significant control on the subsequent
Mesozoic and Cenozoic evolution of the European lithosphere.
In particular, the second conclusion has serious implications for
oil exploration in the North Sea area (Sørensen 1986; Pedersen
et al. 1991; Odinsen et al. 2000) in terms of reassessing the
reconstructed thermal history of the region and, by extension,
predictions concerning levels of organic maturation.
Ziegler et al. (2006) have noted that subsidence curve modelling
suggests that there was a period of Permo-Carboniferous ‘stretching’ from 300 to 280 Ma. This involved decoupled crustal extension and mantle-lithosphere attenuation. Such an assumption is
compatible with the concept that during the Permo-Carboniferous
re-equilibration of the crust – mantle boundary crustal extension
played a significant, but local, role. This can be concluded from
the fact that some important Permo-Carboniferous troughs
(e.g. Massif Central, Bohemian Massif) do not coincide with
major Late Permian depocentres, and depocentres such as the
Southern Permian Basin are not underlain by major PermoCarboniferous basins (Ziegler 1990; Ziegler et al. 2006). This
would suggest that during the Permo-Carboniferous tectonomagmatic cycle uniform and/or depth-dependent lithospheric extension was, on a regional scale, only a contributing factor and not
the dominant mechanism of crustal and mantle-lithosphere thinning. Mechanical stretching of the lithosphere played a subordinate role whereas thermal thinning of the mantle-lithosphere and
magmatic and erosional thinning of the crust dominated, providing
the principal driving mechanism for the Late Permian and later
subsidence of intracratonic basins.
Thermo-mechanical models
Rheological numerical models investigate the way in which
materials deform in response to stresses. Rocks are considered
to deform in different ways (i.e. by elasticity, plasticity or viscosity) depending on the duration of the load, the petrological
composition, the temperature, and the confining pressure. The
rocks can display a very complex rheology, as it is possible that
various modes of deformation occur at the same time. An
additional control is the depth to the brittle– ductile transitions,
which can be present within each lithospheric layer and are
strongly controlled by the geotherm. Thus, rheological modelling
implicitly involves thermal calculations and can include routines
to determine possible melt volumes.
Application of these rheological models is not straightforward,
as they require various, and relatively accurate, datasets. The
Oslo Rift is one of the few Permian basins in Europe where this
requirement is met and, consequently, it has been the primary
target for 2D numerical modelling (Ro & Faleide 1992; Pedersen
& van der Beek 1994; Pascal & Cloetingh 2002; Pascal et al.
2004). The Oslo Rift presents the paradox of showing little extension (i.e. b , 1.3) in association with huge volumes of synrift
magmatic rocks (i.e. .100 000 km3, Neumann et al. 2004). A
377
possible explanation for this is the presence of an underlying
thermal anomaly (i.e. a mantle plume) below the European
lithosphere in Permian times, which in turn could also explain
the observed widespread rifting and magmatism. However, for
various reasons (see discussion by Pedersen & van der Beek
1994; Pascal et al. 2004) this hypothesis is questionable.
Melting modelling of the Oslo Graben was carried out by Ro &
Faleide (1992) and Pedersen & van der Beek (1994). From a
model in which crust and lithospheric mantle are equally
stretched, Ro & Faleide (1992) argued for the mantle plume
hypothesis. In contrast, Pedersen & van der Beek (1994) showed
that the volumes of melts of the Oslo Rift can be accounted for
by differential stretching between crust and lithospheric mantle
(i.e. the lithospheric mantle is more stretched than the crust) and
reduced melting temperatures for the mantle owing to the presence
of volatiles (i.e. water and CO2).
Based on geophysical observations, Pascal & Cloetingh (2002)
proposed a rheological model that considers lithosphere thickness
heterogeneities in the Oslo region (Fig. 25). Their modelling
shows that such heterogeneities could have resulted in strong
localization of deformation in the Oslo Rift. A similar study by
Pascal et al. (2004) showed that the introduction of lithosphere
thickness contrasts in the models results in pronounced differential
stretching between crust and mantle lithosphere, which, in turn,
leads to decompression melting of the mantle over relatively
short time periods subsequent to the onset of rifting. In
summary, the models of Pascal & Cloetingh (2002) and Pascal
et al. (2004), in which the mechanical behaviour of the rocks
and a more realistic configuration for the lithosphere are included,
complement the study by Pedersen & van der Beek (1994).
Although modelling results are very often more suggestive than
firmly conclusive and need to be compared with nature, whenever
it is possible, they appear here to go against a plume hypothesis for
the Permian rift event in Europe.
Henk (1999) used rheological modelling of Permian basins of
Europe to examine the post-convergence evolution of the region.
The purpose of his modelling approach was to explore whether
the Variscides simply collapsed following the end of the orogenesis, thus leading to Permian rifting, or whether the region was
also influenced by far-field extension. Various 2D models were
presented by Henk (1999), and he concluded that far-field extension superimposed on gravity stresses are required to overcome
the strength of the post-Variscan lithosphere.
Along the LT-7 deep seismic refraction profile in the NW Polish
Basin (Guterch et al. 1994), 1D rheological modelling using a simplified petrological model of lithospheric layering was completed.
The results suggest that the lithosphere, except for the East
European Craton (EEC), is mechanically decoupled, and that the
upper crust is separated from the upper mantle by extremely
weak and ductile middle and lower crustal layers (c. 20 km
thick). Only within the Tornquist– Teisseyre Zone and the EEC
can the lower crust remain strong. The lithosphere of the EEC
is probably entirely coupled except for the edge of the craton,
where, with the low strain rates, mechanical discontinuity may
occur at the middle– lower crust or lower crust– mantle boundaries. Laterally, the cumulative strength of the lithosphere
changes by more than an order of magnitude (Jarosinski et al.
2002; Grad et al. 2003).
Tectonic and structural models
Based on geological and geophysical data, tectonic and structural
modelling of an object usually summarizes and tests the admissibility of combined information measured and observed in the
field and laboratory. Balanced sections thus provide geologically
reasonable constraints (Dahlstrom 1969), a concept that has
been widely used in the hydrocarbon industry (Bally et al. 1966;
Rowan & Kligfield 1989), but also is used to reveal the nature
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T. MC CANN ET AL.
Fig. 25. Numerical modelling of the Oslo Rift, involving rock rheology and heterogeneity in lithosphere thickness (after Pascal et al. 2004). The thickness of the
lithosphere in the left and right parts of the model is initially equal to 125 km and 180 km, respectively. The modelled line is 500 km long at t ¼ 0 Ma. The model
is stretched using a velocity of 1.6 cm a21. The upper panel presents the horizontal strain distributions (i.e. 1xx) 1 Ma and 9 Ma after rift initiation. (Note the strong
strain localization at the middle of the model and the Earth surface depression simulating basin formation.) The lower panel presents the thermal evolution (i.e.
isotherms) of the lithosphere. Note the rise at t ¼ 9 Ma of hot mantle rocks below the area that is depressed at the surface. The finite-element grid used for
the computations is also shown. U.C., upper crust; L.C., lower crust; L.M., lithospheric mantle.
of tectonic processes and kinematic evolution in the area of
interest (e.g. Oncken 1989). In the Central European Variscides,
extensive studies were carried out to determine the pre-Variscan
and Variscan evolution (see summary by Franke et al. 2000),
but only few comprise 2D and 3D geometric and tectonic modelling of late, Variscan (e.g. Plesch & Oncken 1999; Oncken et al.
2000, and references therein; Schäfer et al. 2000) or even postVariscan development (Tanner et al. 1998).
In the NE German Basin, the only palinspastic reconstructions
available are by Kossow & Krawczyk (2002), based on results
from the BASIN96 and commercial seismic surveys (Krawczyk
et al. 1999; Kossow et al. 2000). The flexural cantilever model
(see Kusznir et al. 1991, for model details) was also applied for
forward modelling of the initial phase of NEGB formation in
combination with detailed analysis of core material (Rieke et al.
2001; Fig. 26).
NE German Basin formation was initiated during the Early
Permian and was largely controlled by normal faulting related to
deep-seated ductile shearing, with a steep and faulted eastern
and a gently dipping western basin margin. A post-rift subsidence
phase of 35 Ma immediately followed this east– west extension.
The cantilever model predicts a stretching factor of b ¼ 1.2 in
the basin centre and 1.0 at the margins, which would have only
a slight effect on the crustal structure. The resulting smooth
Moho uplift would fit well with the observed seismic data
(Krawczyk et al. 1999). Restoration of the subsequent postZechstein kinematic evolution of the NEGB along a 260 km
long NE – SW cross-section further indicates two major uplift
periods at the Jurassic– Cretaceous and the Cretaceous– Tertiary
boundaries (Kossow & Krawczyk 2002). Quantification of geological processes yields a total basement subsidence of 2850 m
in the basin centre from end-Zechstein to present, maximum
erosion of 860 m during the Cretaceous – Tertiary event at the
southern NEGB margin, and at least 9 km of basin shortening.
Interestingly, there is a clear correlation between the deformation
intensity and the amount of uplift and erosion associated with the
Cretaceous –Tertiary deformational period in the NEGB.
Deformation intensity decreases from south to north, as do uplift
rates, thus suggesting compression from the south, which was
probably related to Alpine-induced intraplate deformations
(Kossow & Krawczyk 2002).
The Permian – Mesozoic development and tectonic inversion of
the Polish Basin has been modelled using a 3D structural model
combining analysis of 3D depth views and thickness maps
(Lamarche et al. 2003a; Lamarche & Scheck-Wenderoth 2005).
The model confirms earlier ideas that the Polish Basin and the
379
Fig. 26. Schematic cross-section across
the northern part of the NE German
Basin showing the faulted basement
comprising Permo-Carboniferous volcanic
units, which were subsequently overlain
by Rotliegend sediments (after
Rieke et al. 2001).
Mid-Polish Swell are genetically related to the Teisseyre – Tornquist Zone, which seems to have tectonically controlled the development of the area through time (e.g. Kutek & Glazek 1972;
Dadlez et al. 1995; Kutek 2001). When the Mid-Polish Trough
started to form, the Teisseyre –Tornquist Zone constituted a
zone of crustal weakness that was prone to extensional
deformation. Crustal thinning along the Teisseyre – Tornquist
Zone, rifting, and the following Mesozoic subsidence resulted in
additional weakening along the zone. As a result, when the
stress conditions changed from transtensional to compressional
at the end of the Cretaceous, the Teisseyre – Tornquist Zone was
preferentially deformed, inducing the inversion of the Mid-Polish
Trough and the uplift of a central NW –SE-elongated anticlinorium along the former basin axis, as well as the formation of two
bordering marginal troughs (see Krywiec 2002a; Lamarche et
al. 2003a for details). This geometry is the surface expression of
the tectonic squeezing of the Teisseyre – Tornquist Zone, which
played the role of an intra-continental zone of crustal weakness
as modelled by Nielsen & Hansen (2000), Hansen et al. (2000)
and Gemmer et al. (2002). Although the stress magnitudes may
have significantly decreased after the climax of the tectonic inversion, the stress pattern remained compressional, as indicated by
the Cenozoic central horst and marginal troughs developed
above the Mid-Polish Swell (Lamarche et al. 2003a; Lamarche &
Scheck-Wenderoth 2005). The Teisseyre –Tornquist Zone can
be considered as a regional weakness zone within which the deformation was localized. A strong tectonic inheritance of Palaeozoic
and Precambrian basement structures influenced the deformation
during the tectonic inversion (Krzywiec 2004). As a result of the
mosaic nature of its Palaeozoic basement, the southwestern flank
of the Mid-Polish Trough was tectonically unstable during the
Mesozoic, in contrast to the stability of the Precambrian East
European Craton beneath the northeastern part of the Mid-Polish
Trough. The model of Lamarche & Schech-Wenderoth (2005)
and tectonostratigraphic models based on seismic reflection data
(Krzywiec 2004) also show the Zechstein salt-bearing layer
acting as a decoupling level between the pre-Zechstein basement
and the Mesozoic cover in the central and northern segments of
the Polish Basin, inducing disharmonic deformation during
the tectonic inversion.
Thus, the idea is that the TTZ was a zone of weakness allowing
the Polish Trough to form. Such an idea is supported by the fact that
long-lived shear zones (in the crust, but probably also in the mantle)
tend to focus strain without regard to the past tectonic context of the
area. This is a fact, and is totally independent of theoretical models.
For example, the border faults of the Viking Graben are at present
the loci of a high degree of micro-seismic activitiy (e.g. Olesen
et al. 2004). This observation is in clear contradiction to the idea
that crustal thinning implies (following thermal relaxation) lithospheric strengthening (with respect to nearby non-rifted areas).
Furthermore, recent advances in fault zone rheology suggest that
repetitive deformation of the fault zone results in the development
of an in situ mylonitic foliation and concentration of weak phases,
which imply a drastic decrease in the coefficient of friction in the
fault zone and potentially a local drop in crustal strength (Bos &
Spiers 2002; Holdsworth 2004).
Discussion
The Variscan orogen was characterized by a particularly long
period of intracontinental deformation, associated with the collision of Gondwana and Laurussia. The post-collisional evolution
of Europe (i.e. within the latest Carboniferous– Early Permian
time frame) was characterized by the formation of a series of rift,
and wrench-induced, basins across the continent, together with significant magmatic events. From the above outline it can be seen that
although we have a reasonable understanding of the broad evolution of the late stages of the Variscan orogenic event and the subsequent period of wrench fault activity that was widespread across
both the internal and external Variscan provinces, there are many
problems relating to our understanding both of the underlying
mechanisms that controlled the various observed events and of
the detailed integration of the various observations. In particular,
there are problems relating the internal and external zones, which
have, at times, remarkably similar evolutionary histories (e.g.
coeval graben formation and associated volcanism in northern
Spain, Italy and northern Germany). Although certain events
may be interpreted in terms of plume-related activity, how do we
interpret similar successions thousands of kilometres apart?
Although it is clear from the above outline that the post-Variscan
period in Central, Southern and Western Europe was a period of
intense tectonic, magmatic and sedimentary change, any attempt
at summarizing these changes must, by necessity, try to assess
the various possible driving mechanisms involved in the generation
of the post-Variscan basins. It is clear that the coincidence of tectonic activity (both compressional and extensional), magmatic
activity and basin formation (with subsequent sedimentation)
was very different from the periods immediately before and after.
The geological evolution of the region, however, is problematic
given the relative lack of significant Early Permian extensional
structures. The large amounts of crustal-derived and crustalcontaminated volcanic rocks are also problematic. The processes
controlling the post-orogenic modification of the Variscan lithosphere have been variably attibuted to such mechanisms as slab
detachment, delamination of the mantle-lithosphere, crustal extension and plume activity during the Stephanian– Early Permian
phase of wrench faulting and magmatism that overprinted the Variscan orogen and its foreland (see Ziegler et al. 2006, for references). The following sections will attempt to examine the main
controlling mechanisms within the basin, to try to isolate those
that are of greatest importance in terms of overall basin evolution.
Rifting history
The examination of a variety of basins across Europe has allowed
us to compare and contrast the various successions within the
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T. MC CANN ET AL.
basins, as well as features such as basin form, controls on basin
formation, and the magmatic, tectonic and infill history. The
observed contrasts suggest that the underlying processes that controlled the post-Variscan evolution of Europe were very different
between those areas located in the former Variscan foreland basin
and those within the thrust front. The various modelling studies
carried out on the post-Variscan Permian basins suggest very
different mechanisms for each area. Permian rifting in the
former Variscan hinterland seems to have been strongly controlled
by the collapse of the mountain chain (Brunet & Le Pichon 1982;
Prijac et al. 2000) with a possible far-field extension component
also being plausible (Henk 1999). This process may also have
been modified by the slow decay of the associated thermal
anomaly (e.g. Paris Basin). In contrast, rifting in the former Variscan foreland appears to have been dominated by late Variscan
wrench tectonics (van Wees et al. 2000), particularly along the
boundary between Precambrian and Phanerozoic Europe (Dadlez
et al. 1995). Numerical modelling highlights the role of such lithospheric discontinuities in controlling rifting (Pascal & Cloetingh
2002; Pascal et al. 2003).
Butler et al. (1997) have noted that pre-existing heterogeneities
in the continental lithosphere are thought to influence its response
to subsequent deformation. From the late Early Carboniferous
onwards, Laurussia was transected by the Arctic –North Atlantic
Rift System, which was partially superimposed on the Caledonian
suture zone (Ziegler 1990). Indeed, Variscan exploitation of older
Caledonian structures has been reported from other areas (e.g. offshore Ireland; for details, see Shannon 1991; McCann 1996). In
Cornwall, early Variscan thrusts were reactivated as late Variscan
extensional faults (Shail & Alexander 1997). Additionally, the
interaction of the Variscan structures with the pre-Variscan
east– west dextral (Badham 1982) transform fault system
(running from the Uralides through Europe (Pitra et al. 1999) to
the Appalachians) and the NNW – SSE-trending wrench fault
system produced a complex series of conjugate shear zones and
pull-apart structures in the Cornwall area (Willis-Richards &
Jackson 1989) that remained active throughout the early
Permian. It is, therefore, highly likely that, within the area under
discussion, older structures, both Caledonian and Variscan, were
reactivated by later Variscan tectonic activity. However, more
recent work (Ebbing et al. 2006) has suggested that even older
structure may be involved. In their study of the Oslo Graben
they suggested that the rifting in the region is coupled to a reactivation of Precambrian fault systems, and indeed, the very location
of the Oslo Graben is more strongly dependent on the pre-rift
structure of the area than previously assumed.
One factor of note is that Permian wrench activity was not
merely limited to ‘accreted’ Europe, but is also evident in other
parts of the craton where there is sufficient stratigraphic evidence.
In particular, there is evidence of late Carboniferous– early
Permian transtensional tectonic activity in the Dniepr– Donets
Basin (Stovba & Stephenson 1999) and even further afield on
the margins of the East European Craton (Saintot et al. 2006).
Mantle plume dynamics
Another important issue addressed by modelling of Permian
basins, and in particular of the Oslo Rift (Ro & Faleide 1992;
Pedersen & van der Beek 1994), is the eventual role played by a
mantle plume (although this idea has recently been questioned;
see Ebbing et al. 2006, for details). Despite significant differences
in the tectonosedimentary setting and the type of magmatic
activity within the various basins examined, the Stephanian–
Autunian volcanic rocks in the internal Variscides comprise a
high proportion of pyroclastic deposits and are generally of intermediate to felsic composition, of calc-alkaline character, and often
have a significant crustal component, as shown by Sr – Nd isotope
data and the presence of crustal xenoliths, magmatic garnet, and
(locally) topaz, and (for the volcanic rocks in the NE German
Basin) the large amount of inherited zircons necessitating sensitive
high-resolution ion microprobe (SHRIMP) dating (Breitkreuz &
Kennedy 1999). The calc-alkaline character may reflect the derivation of the melts from a subduction-modified mantle source,
extensive assimilation of crustal material, or perhaps inheritance
resulting from the melting of older calc-alkaline, crustal sources
(such as Cadomian basement). However, the relative scarcity of
more primitive mafic melts precludes a more precise interpretation
of the mantle source compositions. In addition, numerical studies
suggest that huge volumes of magmas can be produced with small
amounts of stretching and without the need for any underlying
thermal anomaly (Pedersen & van der Beek 1994).
Crustal processes
The Stephanian–Autunian magmatic rocks in the internal
Variscides comprise a high proportion of pyroclastic rocks and are
generally of intermediate to felsic composition. Their generally
calc-alkaline character suggests a subduction-related origin. With
the possible exception of some magmatic rocks in the Alpine basement, this contradicts their intracontinental setting and the fact that
the Variscan oceans had closed by mid-Carboniferous times.
However, Sm–Nd isotope data and the presence of garnet and
crustal xenoliths indicate that many contain a significant crustal
component. This is corroborated by the predominantly negative
1Nd(t) values of the 290–300 Ma volcanic and intrusive rocks of
felsic to intermediate composition: 22.1 to 26.0 for the Krkonoše
Basin (Ulrych et al. 2002), 22.7 to 26.1 for the Intra-Sudetic
Basin (Ulrych et al. 2004); 20.8 to 27.0 for the rhyolites of the
Halle Volcanic Complex (Romer et al. 2001), 24.3 to 27.5 for
the granites in Cornwall (Darbyshire & Shepherd 1994), and 20.6
to 25.7 for the Saar–Nahe Basin (Schmidberger & Hegner 1999;
von Seckendorff et al. 2004a, and references therein). The parent
magmas of the granitoids, rhyolites and andesites may, therefore,
have assimilated large amounts of crustal material, or alternatively,
be derived from mantle sources that had been modified by earlier
subduction events (e.g. Cabanis & Le Fur-Balouet 1989; Schmidberger & Hegner 1999; Innocent et al. 1994; Cortesogno et al. 1998).
As in the North German Basin, the granites and rhyolites may be
of crustal origin, and their calc-alkaline signature inherited through
partial melting of calc-alkaline basement (Schaltegger 1997b;
Romer et al. 2001). The possible mechanisms for mantle melting
in the internal Variscides may have been the break-off of subducted
oceanic crust (e.g. Schaltegger 1997b; Cesare et al. 2002) or even the
oblique subduction of the mid-ocean ridge of Palaeotethys beneath
the active Eurasian margin (Stampfli 1996). Regional extension
leading to lithospheric thinning and decompressional melting of
updoming asthenosphere may have been a contributing factor in
the late Carboniferous–early Permian period. Compared with the
foreland, Stephanian–Autunian mafic rocks are much rarer in the
internal Variscides, which suggests that the mantle-derived parent
melts were unable to reach the surface, but stalled at lower to midcrustal levels. This may have been due to the large contrast
between the density of the parent melt and a low average density
of thinned Variscan crust. Only after fractionation and assimilation
of sufficient amounts of crustal material did the melts attain a low
enough buoyancy to be able to escape the magma chambers and
erupt on the surface.
Magmatic – tectonic activity
The relatively short and widespread pulse of Stephanian –
Autunian magmatism is likely to have taken place in response to
changes in the regional stress field at the Westphalian – Stephanian
boundary and subsequent thermal equilibration of the lithosphere.
The change of stress may have been due to a change in Viséan–
Westphalian crustal shortening and orogen-parallel extension,
and to Stephanian– Autunian gravitational collapse of the
Variscan orogen. The latter process was possibly superimposed
and aided by a far-field dextral extensional stress-field that was
due to the collision of Gondwana with eastern southeastern
North America and concomitant dextral translation (Torsvik &
Van der Voo 2002). The invocation of far-field effects is something that has previously been noted in discussions of postVariscan tectonics (e.g. discussion on the origin of the NE
German Basin; see DEKORP-BASIN Research Group 1999, for
details). In terms of the magmatic history of the post-Variscan
there are some indicators that far-field effects might also have
played an important role. For example, the alkaline composition,
style of volcanism and the presence of abundant megacrysts and
mantle xenoliths in the Early Permian mafic rocks in Scotland
indicate derivation by low-degree melting of local mantle
sources and rapid, vertical transport. In contrast, the sub-alkaline
mafic dyke and sills complexes (such as the Whin Sill Complex)
indicate higher degrees of mantle melting, and do not necessarily
reflect a mantle thermal anomaly of the same extent. The geometry
and orientation of the dyke swarms suggest a magmatic focal region
in the vicinity of the Denmark – Skagerrak region, which suggests
that magma transport may have been horizontal, westwards into
the North Sea and Scotland. Thus, the position, trend, number
and size of the dykes may have been controlled by the far-field
dextral extensional stress field.
Conclusions
The end Carboniferous–early Permian history of Europe represents
a period of crustal instability and re-equilibration throughout
Western and Central Europe. An extensive and significant phase
of Permo-Carboniferous magmatism led to the extrusion of thick
volcanic successions across the region. Coeval transtensional
activity led to the formation of more than 70 rift basins, which
differ both in form and infill according to their position relative to
the former Variscan Orogenic Front as well as to the controls that
acted on basin development. Despite the fact that no unified model
for the Permian event can at present be unequivocally proposed
from the results of the various modelling studies, recent studies do
agree on two fundamental and relevant points: (1) Permian rifting
was widespread in Europe with progressively propagated development; (2) its signature strongly influenced the evolution of the European lithosphere during Mesozoic and Cenozoic times (Sørensen
1986). It may not, however, be possible to provide more detailed
models for the evolution of the region. Numerical modelling of lithospheric rifting, for example, requires numerous parameters, among
which the pre-rift crust and mantle-lithosphere structure are
crucial. Because the pre-Permian lithosphere structure has been
obscured by repetitive tectonic phases in most parts of Europe,
lithosphere-scale modelling of the Permian event remains difficult
and modelling results need to be treated with a high degree of circumspection. The best approach, therefore, to elucidating the tectonosedimentary and magmatic history of the region is to adopt a broad
approach, examining the various basins at a range of scales and
making use of a variety of techniques.
This manuscript was greatly improved by the reviews of two anonymous
reviewers. T. Beilfuss is thanked for the production of the excellent diagrams.
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Post-Variscan (end Carboniferous–Early Permian) basin evolution

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