Kinematics of orogenic collapse in the Variscan Pyrenees deduced from microstructures in porphyroblastic rocks from the Lys-Caillaouas massif

Tectonophysics 238 (1994) 139-160

Kinematics of erogenic collapse in the Variscan Pyrenees deduced from microstructures in porphyroblastic rocks from the Lys-Caillaouas massif D.G.A.M. Aerden Laboratoire de Giophysique et Tectonique, Case Postal 060, Universit6 de Montpellier II, Place E. Bataillon, 34095 Montpellier Cedex 5, France

Received 10 June 1993; accepted January 7,1994

Abstract In the Pyrenees, a phase of bulk vertical shortening (DJ is indicated by a horizontal crenulation cleavage (S,), which is developed in a pre-existing schistosity 6,). Porphyroblasts of the Lys-Caillaouas massif regionally preserve: (1) the initial pre-D, orientation of S, in the form of a vertical and E-W-trending preferred orientation of their inclusion trails; and (2) a vertical L, stretching lineation defined by the elongation of the inclusions. These data are consistent with N-S compression and crustal thickening during D,. The angle between the inclusion trails (Q) and the (external) schistosity (S, J records the subsequent rotation of S, from vertical to gently dipping orientations. At higher crustal levels S, remained steep (suprastructure) and little D, deformation is evidenced. With depth, D, deformation progressively intensifies and zones can be distinguished with opposite dip directions of S,, opposite apparent porphyroblast “rotation” senses and opposite F3 crenulation asymmetries. These (micro)structures on a N-S section through the Lys-Caillaouas and the Garonne Dame massifs record a regional partitioning of bulk coaxial vertical shortening. This deformation partitioning was probably controlled by a not-exposed gneiss dome, which behaved relatively rigid as compared to its metasedimentary cover. In zones of non-coaxial D, deformation, S2 was extended, reactivated and forms the dominant schistosity. Where D, was locally more coaxial, S, was shortened, isoclinally crenulated and S, forms the macroscopic cleavage. The fact that D, extension in the Lys-Caillaouas massif immediately followed D, compression and that similar extension is recorded in large parts of the Variscan foldbelt, suggests gravitational collapse of a thickened orogen. This is consistent with the flattening type of D, strain recognized in a number of Pyrenean Variscan massifs. Peak high-T, low-P metamorphic conditions coincide with the onset of erogenic extension, when very little crustal thinning could have been accomplished yet. This suggests that heat-induced strain softening within the orogen may have triggered its collapse.

1. Introduction The importance of gravitational or extensional tectonics as opposed to, or as part of, plate tectonics has since long been recognized and dis0040-1951/94/$07.00 d

cussed (e.g., van Bemmelen, 1972; Elliott, 1976; Ramberg, 1977; Platt and Vissers, 1989; Bell and Johnson, 1989). In more recent years, earth scientists have recognized that the uplift a&l thickening of mountain belts is limited by their internal

0 1994 Elsevier Science B.V. All rights reserved

physical strength (Royden et al., 1983; Pfatt, 1987;

The kinematics (e.g., coaxial or non-coaxial) and transport directions of this erogenic extension are poorly constrained and, consequently, controversy surrounds the question what caused the extension. For example, rifting (Wickham and Oxburgh, 1986), diapirism (Soula, 1982; Pouget et al., 1988) and gravitational collapse of a thickened orogen (Vissers, 1992) are relatively recent explanations, Vissers (1992) considered that extension was mainly orogen-parallel (E-W) based on E-W rotations of porphyrobfasts, subhorizontal E-W stretching lineations and the geometry of Stephano-Permian fault-bound basins. However, porphyroblast rotations about N-S axes have also been described (Lister et al., 1986; de Bresser et al., 1986; Kriegsman et al., 1989) and subhorizontal stretching lineations have other than E-W orientations (including N-S) in many Pyrenean massifs (van den Eeckhout, 1986; Sofiva et al,, 1989; de Saint Bfanquat, 1990, 1994). Additionally, as the Variscan Pyrenees suffered complex polyphase deformation (up to six deformation events according to van den Eeckhout, 19861, subhorizontal stretehing lineations do not neces-

Dewey, 1988; Molnar and Lyon-Caen, 1988). This implies that when a mountain belt is uplifted beyond a critical level it will become gravitationally unstable, starts thinning by internal deformation and spreads outward. Plate convergence may continue during this process while outward spreading is accommodated by external thrusts Whereas the kinematics of erogenic extension by normal faulting, extensional shear-zone devefopment and core-complex formation have received wide attention (e.g., Davis et al., 1986; SCranne and Seguret, 1987; Malavieille et al., 19901, relatively Iittle is known about the fdnematics of the more homogeneous internal deformation of the extending Crust. In this paper, such deformation is considered in the Pyrenees, where Upper Carboniferous erogenic extension affected the Variscan crust (D,) following N-S compression CD,) and crustal thickening (Verhoef et al., 1984; Lister et al., 1986; van den Eeckhout, 1986, 1990; van den Eeckhout and Zwart, 1988; Kriegsman, 1989a, b; Kriegsman et al., 1989; de Saint Blanquat et al., 1990; Gibson, 1991).

F

R

A

N

C

E

FRANCE

I

-

Alpine fault

-A

I

Fig. 1. SimpHied map of the central Pyreneas, mctdifkd after Zwart (E979), shuwing the position of the Gaty%tneDome-( tys-Caillaouas maasif (LC) and a number of other %iscan mass&. Majar Alpine faults are labelled: NPF (North Pyreneatr Fault), GT (Gavarnie Thrust) and MF (Merens Fault), The Paiaeozoic rock8 south of the NPF constitute the “Axial~Zone” of the Pyrenees.

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

sarily indicate the extension direction during D,, especially considering possible reactivation of foliations during later deformation events (Bell, 1986). This paper attempts to reconstruct the kinematics of erogenic extension in the Pyrenees by looking in detail at the microstructural expression of this deformation phase in the Lys-Caillaouas massif. By means of selective microscopic strain measurements, D, and D, extension directions are reconstructed from the finite strain state. Asymmetry switches of D, microstructures in the Lys-Caillaouas and Garonne Dome massifs (Figs. 1 and 2), in particular of matrix-porphyroblast rotation senses, are shown to register a crustalscale pattern of coaxial vertical shortening, partitioned into zones with opposite horizontal shear senses. The strain patterns obtained by this method are used to develop a kinematic model

141

for the Lys-Caillaouas-Garonne Dome region in terms of gravitational collapse of a thickened orogen. This model integrates structures on all scales of observation, can be extended to a number of other Variscan massifs in the Pyrenees and explains the apparently conflicting data of previous workers.

2. The Variscan Pyrenees The Variscan Pyrenees form part of the Variscan orogen in Europe (e.g., Matte, 1991) and were uplifted in the core of the Pyrenean chain during the Alpine orogeny (Fig. 1). A number of massifs can be distinguished with intermediate- to high-grade metamorphism consisting mainly of lower Palaeozoic metasediments. Variscan orogenie activity was concentrated in the Late Car-

Fig. 2. Map of the Lys-Caillaouas and Garonne Dome massifs after Zwart (19791, Pouget et al. (1988) and Kriegsman et al. (1989). Bold cross-section lines correspond to Figs. 3 and 14a. The area concentrated upon in this paper is situated in the northern part of the Lys-Caillaouas massif (dashed box). The microstructure of Fig. 15 is from the location indicated with an asterisk in dhe southern part of the Garonne Dome.

142

D.G.A.M.

Aerden / Tectonophysics

boniferous (Zwart, 1979). In a number of massifs (e.g., St. Barth&my, Agly, Aston-Canigou) preVariscan gneisses are exposed, the origin of which is still not fully resolved. They are either regarded as pre-Cambrian basement (Cavet, 1957; Guitard, 1970; van den Eeckhout, 1986), and/or as lower Palaeozoic granite intrusions (Zwart, 1979). Alpine deformation/ metamorphism vastly overprinted the massifs north of the North Pyrenean Fault (North Pyrenean massifs), whereas south of this fault Palaeozoic rocks constitute the “Axial Zone” of the Pyrenees (Zwart, 1979), in which Alpine deformation is restricted to narrow shear zones, bounding internally little-affected Variscan blocks (Fig. 1; Zwart, 1979; de Saint Blanquat, 1990, 1994). The Lys-Caillaouas massif and other massifs considered in this paper belong to this Axial Zone. The Axial Zone of the Variscan Pyrenees has been subdivided into a “suprastructure”, i.e., higher structural levels with greenschist-grade metamorphism and vertical foliations, and an “infrastructure”, i.e., deeper structural levels with amphibolite-grade metamorphism and predominantly subhorizontal foliation (Zwart, 1979). It is well established that the steep foliation in the suprastructure formed during a main phase of crustal thickening and N-S compression (e.g.,

238 (1994) 139-160

Zwart, 1979; Corstanje et al., 1989; Kriegsman et al., 1989; Kriegsman, 1989b; Vissers, 1992). In contrast, the foliation in the infrastructure (S,) is generally considered to be younger and to have formed during subsequent crustal extension (Verhoef et al., 1984; van den Eeckhout, 1986, 1990; de Bresser et al., 1986; Lister et al., 1986; van den Eeckhout and Zwart, 1988; Pouget et al., 1988; Kriegsman et al., 1989; Kriegsman, 1989a; Gibson, 1991; Vissers, 1992). However, new data in this paper show that a gently dipping foliation in the infrastructure is commonly a rotated and extended S,, instead of a younger S,. This has far-reaching implications for the kinematic reconstruction of crustal extension in the Variscan Pyrenees.

3. StNctuNl

deYebpment of the ~s&@a&Mias

massif The Lys-Caillaouas massif (Figs. 2 and 3) consists of Cambru-Ordovician metapelites (schists) and quart&es that are rich in biotite, and&u&e, cordierite and staurolite porphyroblasts. The ddest recognized foliation is a we&deveIoped s&istosity defined by oriented phyllosiiicates and flat quartz grains. It constitutes the dominant macro-

m

F4,w

/-

Fig. 3. Cross-section through the northern part of the Lys-Caillaouas massif showing the attitude of the S, and S3cleavages in the Cambro-Ordovician metasedimerits. Cross-section line and legend are given in Fig. 2. Circles in the c~section represent the projected locations of the microstructures shown in other figures (fig. numbers are given). F, folds and sedimentary layering have been omitted for clarity.

D.G.A.M. Aerden /Tectomphysics

scopic cleavage in large.parts of the Pyrenees and was termed “main-phase cleavage” or “S” by Zwart (1979). This cleavage is in this paper referred to as S,, because of evidence of earlier pre-Variscan deformation in the Axial Zone of the Byrenees (Hartevelt, 1970; Santanach, 1972; den Brok, 1989). Also, relics of a pre-S, cleavage were reported by Garcia Sansegundo and Alonso (1989) and Garcia Sansegundo (19921, but no associated folds. The labelling of cleavage generations in this paper can be directly correlated with that of van den Eeckhout (1986, 1990), van den Eeckhout and Zwart (1988), Kriegsman (1989a, b), Kriegsman et al. (1989) and Gibson (1991). In the study area, S, averages a southerly

238 (1994) 139-160

143

dip, but progressively steepens upwards, (Fig. 3) and is axial planar to E-W-trending folds of all scales (de Bresser et al., 1986; Kriegsman et al., 1989; Kriegsman, 1989b). S, is overprinted by a lower dipping to subhorizontal crenulation cleavage (S,), which is the foliation related by previous workers in the LysCaillaouas and in other Pyrenean massifs to crustal-scale vertical shortening or erogenic extension. S, is axial planar to recumbent mesoscopic folds that are relatively rare in bhe study area (Fig. 3). Locally, S, forms the dominant cleavage where S, is transposed by isoclinal crenulation. Both S, and S, are affected by upright folds of

Fig. 4. (a) Photomicrograph and line tracing of a symmetrically crenulated S,, formed by shortening perpendicular to S,. S3 is defined by the axial planes of crenulations. For location see Fig. 3. (b) Photomicrograph and line tracing of a differentiated S, crenulation cleavage with asymmetric ‘7 geometry crenulations. S, cleavage planes are defined by the crenulation lbng limbs, in which (shear) strain was concentrated and quartz was selectively dissolved. Sinistral shear components are indicated by the progressive curvature of S, from the quartz-rich short limbs into the mica-rich long limbs. For approximate location see Fig. 3.

144

LXG.A.M. Aerdcn / Tectunophysics 2% (1994) 139-160

all scales that are locally associated with a subvertical crenulation cleavage (S,). This cleavage is intensified near E-W to ENE-WSW-trending subverticai fault/shear zones with south-bl~k-up movement components (possible transcurrent components are unknown). These structures are tentatively correlated with the late-Hercynian steep foliations, shear zones and coeval folds described by Matte (1969, his D,>, Carreras and Cires (2986), van den Eeckhout (1986, his D,) and Garcia Sansegundo (1992, his D, and D4).

Porphyroblasts of the Lys-Caillaouas massif include the dominant schistosity (S,) as inclusion trails, which are generally straight or slight’ly curved. More rare, crenulated and sigmoidal inclusion trails that change asymmetry from one F, fold limb to the other, further establish an early syn-D, timing for po~hyroblast growth (de Bresser et al., 1986; Lister et al., 1986; Pouget et al., 1988; Kriegsman et al., 1989; Aerden, 1995). The metasediments of the Lys-Caillaouas massif were intruded by a large granite body (Fig. 2) after peak metamorphic conditions were reached, i.e., late syn- to post-D, Kriegsman et al., 1989), whereby a narrow contact-metamo~hic zone developed, characterized by the partial breakdown of regional metamorphic staurolite, renewed growth of andalusite and new growth-of siliimanite. In the northern part of the Lys-Caiflaouas massif, Ca~bro-Ord~ci~ schists are faulted against Silurian dark slates @ii. 2 and. 3) and a N-dipping thrust occurs, which is the continuation of the “Gavamie Thrust” more to the west. Mesozoic rocks along the Gavamie Thrust indicate an Alpine age (Zwart, 1986).

4. T4e nature of S, in the Lys=C!

Fig. 5. Line tracing of a differentiated Z&crendation cleavage wrapping andalusite ~~hy~blasts (shaded areas). Thinning and boudinage of a quartz veinlet, as it curves out of the larger andalusite porphyroblast, demoastmtes sinistrai shear strain in the matrix, with shearing concentrated in S, cleavage planes. Thus, the angle between Sz, and S,, and the open fold geometry of the quartz veinlet did not result from rolling of the porphyroblast, but from heterogeneous sinistral shear in the matrix. The lack of ~~h~ob~ast rotation relafive to S3 was balanced by concentration of rotational strain components in S, cleavage planes anastomosing around the porphyroblast. For approximate location of this mkrostrueture see Fig. 3.

massif

At high structural levels and moderate D, strain, S, is defined by symmetric F3 crenulations developed in a steeply dipping or subyertical S,. S, is not a true foliation in these zones, but just defined by the axial planes of crenulations (Fig. 4a). With increasing intensity and non-coaxial@ of D, strain, crenulations tighten, become more asymmetric, whereas S, averages a lower dip at a decreased angle with S,. In these areas, S, is a differentiated crenulation cleavage with quartzrich short-limb domains (microlithons) ~ntaining a steep S, and narrow mica-rich long-limb domains containing a subhorizontal S, and defining the S,-cleavage planes (Fig. 4b). Concentration of sinistral shear strain in the iong-limb zones (viewing east) and shortening components peipendicular to S, are indicated by: (1) the consistent sense in which S, curves out of the quartz-rich microlithons into the mica-rich S,-cleavage planes,

D. GA. M. Aerden / Tectonophysics 238 (1994) 139-160

producing “Z’‘-asymmetry crenulations viewing east (Figs. 4b and 5), hereby assuming that quartz dissolution was proportional to strain and that S, was initially steep; (2) progressive rotation, boudinage and offset of quartz veinlets across individual &-cleavage planes (Fig. 5); (3) early syn-D, porphyroblasts enclosing weak F, crenulations

145

that continue parallel but more tight and asymmetric in the matrix, the porphyroblasts being wrapped by S,-cleavage planes (Fig. 6). Thus, S, embodies a regular partitioning of bulk general shear (pure- and simple-shear components) with shear strain concentrated in S,cleavage planes (Bell, 1981; Bell and Cuff, 1989).

Fig. 6. Andalusite porphyroblast enclosing weak F, crenulations which continue straightly but more tightly and more asymmetrical in the matrix. Wrapping of S, around the porphyroblast indicates shortening components perpendicular to S,, hereas the crenulation asymmetry in the matrix indicates sinistral shearing components. For approximate location of this micros4 ructure see Fig. 3.

136

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

The local strain magnitude can thus be estimated from the angle between S, and S,; a low angle indicating high (shear) strain. Mesoscopic F, folds have an analogues geometry to F, crenulations, with short limbs at a high angle to S, (low strain) and long limbs at a low angle to S, (high shear strain). They are therefore similarly interpreted as a result of heterogeneous general shear with S, as flow plane. The scarcity of F, folds and the lack of macroscopic D, shear zones indicate that D, deformation was relatively homogeneous on the scale of the Lys-Caillaouas massif.

5. Porphyroblasts as kinematic indicators Microstructural data were collected in the northern, central part of the Lys-Caillaouas massif (Fig. 2) from: (1) N-S-trending vertical thin sections; (2) E-W-trending sections perpendicular to S, and parallel to L$ and (3) horizontal thin sections. Measurements of the orientation of porphyroblast inclusion trails in these sections revealed a strong subvertical and E-W-trending preferred orientation (Fig. 7). Aerden (1995) analyzed these data in detail and demonstrated that the preferred orientation of inclusion trails is due to a lack of porphyrol&st rotation r&&e to S, during D,. His main evidence was: (1) ~ralieiism between internal (within porphyroMpst$3 aad external S, (e.g., Figs. 6 and 8); and (2) idependent shear-sense cr&xia contra&&g w&h shearinduced rolling of porphyroblasts (e.g., Fig. 5). Rigid objects have been previously shown to have maintained a stable orient&ion relative to the flow plane due to strong @rmti partitioning effects (cf. Bell, 1981, 1985; Haruner and Passchier, 1991; Hayward, I!@>), whereby the Lack of object rotation was ~&&&XXIby i&en&i&on of shear strain in foiiattion planes an&w around it. The la& of the Lys-Caillaouas massif has the cations for kinematic analysis in the area. (1) The typical “Z” asymmetry of S, as it passes through porphyroblasts with straight inclusion trails does not record rotation of the porphyroblast relative to S,, but anticlockwise rotation

t

HORIZONTAL

SECTJONS

S UP

VERTICAL SECTIONS Fig. 7. (a) Rose diagram showing strong E-W alignment of parphyr&last inclusiontrails as -red in horizontal thin sections. Data rcgreoent 81 amklusite. staurolite and co&&e poEplq*dafBBts frosn six hand specimens. A 10 circle @aded area = scale) represents orie&&&Cb).iZose diagram showing a vertkal preferred or&H&on of @&ion trails for 348 por*if, measured in vertical pkqtwblaats of the Lyst& s+x-tions. The d#a were collected from 24 ~diatribn&d in the area indicated in Fig. 2. A lo” sqmimt of tbc psot cb$e (shat86d area = s&e) represents 8% of 1 red orientations. See Aerden (1993) for f-data. d&ailed-an

of the e&err& S, relative to S, due to sinistral shear @rain in the matrix (Fig. 51. Similarly, @a&Isresulted from gr&vth within developing short-iiinb domains of F3 crenulations or microfolds (Fig. 8). (2) As S, has maintained a relative& constant orientation through time, the inch&on tr& xlrave also. Consequently, their vertical preferred orien-

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

tation records the orientation in which S, formed in the Lys-Caillaouas massif. This is consistent with an origin of S, as a vertical foliation during orogen thickening and N-S-directed compression (van den Eeckhout, 1986, 1990; Corstanje et al., 1989; Kriegsman et al., 1989 and many others).

147

Fig. 9 graphically illustrates porphyroblastmatrix relationships as observed in different sets of thin sections. In horizontal thin sections porphyroblast inclusion trails trend E-W, approximately parallel to the external S,. In N-$-trending vertical thin sections, S,,i/S,,e relationships

Fig. 8. (a) Line diagram of an andalusite porphyroblast enclosing delicately crenulated, but on a larger scale, sigmoidal S2 inclusion trails that are continuous with the external foliation. This microstructure suggests initiation of porphyroblast growth early syn-D, in a nucleating F3 short limb, followed by progressive rotation of S, in the long limb towards the immaterial flow plane (S,) and into an extensional orientation. S, is preserved within the porphyroblast (short limb) but inferred to have been destroyed in the matrix (long limb) due to extension and reactivation of S2 at advanced D, deformation stages (cf. Fig. 11). (b) Microphotograph corresponding to the area indicated in Fig. 4a, showing crenulated inclusion trails that are continuous with the external foliation.

D.G.A.M.

138

Aarden / Tectonophysics

23X (1994) 139-lh0

tectonic transport based on porphyroblast “rotations” in E-W-trending thin sections of the LysCaillaouas massif (Zwart, 1979; de Bresser et al., 1986; Kriegsman et al., 1989; Vissers, 1992) is considered to be incorrect.

6. The effect of progressive D3 deformation on S,

P Fig. 9. “Unfolded” block diagram schematically illustrating porphyroblast-matrix relationships in 3-D. Inclusion trail geometries are cylindrical in E-W direction. Consequently, the opposite porphyroblast “rotations” observed in E-W-trending sections perpendicular to S, are cut effects that resulted from having sectioned close to anastomosing axes of relative matrix-porphyroblast rotation. Strain shadows of porphyroblast and flat biotite platelets @lack lenses) are observed in both N-S- and E-W-trending thin sections. The biotite platelets appear slightly longer in E-W- than in N-S-trending thin sections.

indicate sinistral shear strain in the matrix (Fig. 5). However, in E-W-trending sections, opposite senses of relative matrix-porphyroblast rotation were observed and highly variable rotation angles. These relationships indicate that inclusion trails are cylindrical in E-W direction and that the porphyroblast “rotations” recorded in the E-W-trending thin sections are cut effects that resulted from having sectioned subparallel to anastomosing F3 crenulation axes, The described geometric relationships are consistent with noncoaxial vertical shortening affecting an E-Wtrending vertical S, (Fig. 10). From the model of Fig. 10 it follows that relative matrix-porphyroblast rotation axes do not indicate stretching or tectonic transport directions, as assumed by workers envisaging shear-induced rolling of porphyroblasts. Consequently, the inference of E-W

During progressive “general shear” (both pureand simple-shear components) a pre-existing foliation, initially oriented at a high angle to the flow plane, can rotate from the shortening into the extensional field of the incremental strain ellipsoid. Microstructurally, this strain path may result in early crenulation (shortening) and subsequent progressive decrenulation (extension) of the foliation at advanced deformation stages (Fig. 11). If the extending foliation represents a structural anisotropy, it may eventually be reactivated as an internal “slip system” (Fig. 11; Bell, 1986; Aerden, 1991; Hanmer and Passchier, 1991; Hayward, 19921. In terms of folding mechanisms this process could be described as a change from heterogeneous shearing parallel to the axial planes of folds, to flexural slip along the anisotropy. Progressive rotation, extension and reactivation of S, during D, is in the Lys-Caillaouas indicated by the following. (1) The progressive fading out of Fi crenulations and microfolds from low to high D, strain zones (Fig. 12). (2) Some porphyroblasts with a prolonged growth history contain a steep crenulated S, in their core, but towards their external parts a progressively less crenulated and more inclined S, is continuous with a straight S2 in the matrix (Fig. 8). (31 !&-defining inclusions within porphyroblasts are elongate in vertical direction, whereas S,-defining grains in the adjacent matrix are elongate in horizontal E-W direction (see below). The difference in finite strain must be due to extension and reactivation of S, after porphyroblast growth, i.e., syn- or post-D,. (4) !&-parallel pegmatite and quartz veins are commonly boudinaged. Their exact parallelism with S, and the evidence of syn- to post-D,

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

magmatic activity in the area suggest that these veins formed parallel to S, instead of having rotated into parallelism with it during non-coaxial deformation. A likely timing of their formation is therefore early syn-D,, when the maximum shortening direction was oriented subparallel to S, (Fig. 11). The boudinage would then record subsequent extension and reactivation of S, later during D, (Fig. 11). Thus, it appears that, whereas F, crenulations and S, may be well developed in areas of moderate D, strain, this is commonly not so in zones of high D, strain. Care is therefore required when interpreting a single schistosity in areas of intense non-coaxial D,. In zones of relatively coaxial deformation, S, probably remained in the shortening field throughout the D, deformation history and could have been eventually transposed by a subhorizontal S,. In contrast, in zones with significant shear-strain components, progressive rotation of S, is inferred to have led to its extension and reactivation, whereby S, was destroyed. The general structure envisaged for the LysCaillaouas massif and the relationship between

149

infrastructure and suprastructure in general are schematically pictured in Fig. 13. The transition from suprastructure with subvertical S, to infrastructure is inferred to involve progressive increase of shear strain and consequent rotation, extension and reactivation of S,. The depth axis can also be regarded as a time axis as the infrastructure would represent a kinematically more advanced stage as the suprastructure. In both domains S, is the dominant foliation. The model explains what Kriegsman (1989b) already recognized in the Lys-Caillaouas massif, namely, that the angle between S, and S, progressively decreases with depth until, eventually, only one schistosity is observed.

7. D, kinematic axes in the Lys-Caillaours massif 7.1. Mesoscopic strain data S,-parallel extension in the Lys-Caillaouas massif is mesoscopically indicated by: (1) strain shadows of porphyroblasts; (2) boudinage of the

Fig. 10. 3-D Kinematic model illustrating how various microstructures resulted from non-coaxial vertical shortening affecting an E-W-trending vertical S,. The Ds maximum stretching direction is hypothetically chosen to be ENE-WSW. Note that this direction cannot be predicted from the direction of L: nor from the axes of relative porphyroblast-matrix rotations. Insets show the most common porphyroblast-matrix relationships in thin section.

IS0

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

already described &-parallel quartz and pegmatite veins; and (3) elongation of conglomerate pebbles. These features are equally well observed in N-S and E-W directions parallel to S, and hence, indicate a flattening type of strain (X > Y B Z). The early syn-D, timing of the porphyroblasts, which is also inferred for the quartz/ pegmatite veins (see above), implies that they record D, strain, at least, away from the D, deformation zones. Flattening strain during D, was previously determined by Pouget et al. (1988) in the Garonne Dome and by van den Eeckhout (1986) in the Hospitalet massif (Fig. 1). A pronounced E-W lineation is locally developed in the Lys-Caillaouas massif, which superficially resembles a stretching lineation. However, differently oriented thin sections show that this lineation is defined by the intersection of S, and S, foliations and by Fi crenulation axes. The lack of a pronounced L, stretching lineation is consistent with a fairly uniform flattening type of strain during D,.

7.2. Microscopic strain analysis The early D, timing of porphyroblast growth allows to separate out the kinematic axes of the D, and D, deformations on microscopic scale by using the following method. As porphyroblast inclusions were trapped early during D,, their shape anisotropy records mainly D, strain. The 3-D shape of these inclusions was disclosed by statistically measuring their geometry in orthogonal sections perpendicular to the inclusion trail plane. The results of this investigation are shown in Table 1, which gives length to width ratios of the measured opaque inclusions in a number of staurolite and cordierite porphyroblasts with vertical, E-W-trending inclusion trails. The inclusions appeared to be 20% more elongate in vertical than in horizontal thin sections revealing a steep or vertical L, stretching lineation, which is consistent with crustal thickening during D, (Zwart, 1979; Corstanje et al., 1989; Kriegsman et al., 1989; Kriegsman, 1989b; Vissers, 1992).

Fig. 11. Diagram illustrating the possible effect of progressive general shear on a pre-existing foliation, initially oriented perpendicular to the bulk flow plane. During early deformation stages, the foliation is oriented in the shortening field of the but& incremental straia ellipse and is being crennuiated. At advanced d&rmation stages, the foliation has too&d i&o We extensiona feld of the incremental strain ellipse, decrenulates and eventu& iweactivotedas an internal a& &em. Tfre wnse af raw! the anisotropy is antithetic to that of the bulk flow. Note that the bulk flow plane is no longer mate&d (expressed me& as SJ in the reactivation stage.

151

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

The D, kinematic axes were reconstructed by considering opaque minerals defining a reactivated S, in the matrix and comparing their shape with that of their counterparts within porphyroblasts, which did not suffer D, strain. Any difference in shape anisotropy should be due to D, deformation, at least away from the zones of local D, deformation. Measurements were made in two hand specimen containing a gently ‘dipping reactivated S,, from which N-S and E-W-trending thin sections were cut perpendicular to S, (Table 2). Opaque grains defining S, in these rocks now appeared to be only 9% more elongated in E-W direction than in N-S direction, indicating that the principle D, stretching direction was oriented somewhere between NE-SW and SE-NW, thus at right angles to the vertical D, elongation direction. In the same hand specimens a weak mineral lineation is defined by thin (10 to 30 pm thick) &-parallel biotite platelets, which appeared longer in E-W-trending thin sections than in N-S-trending thin sections.

/v,

1

8. Comparison with other Pyrenean massifs 8.1. Garonne Dome Foliations equivalent to S, and S, of this paper were described by Zwart (1979) and Pouget et al. (1988) as S, and S, in the Garonne Dome. In the following discussion, their S, and S, are directly translated to S, and S,, respectively. Pouget et al. (1988, fig. 1) interpreted the dominant schistosity in the amphibolite-facies part of the Garonne Dome as defining the dome structurally and being S,. S, would be completely obliterated in this area. However, Zwart (1979) interpreted the dominant schistosity in the Garonne Dome as S,, because he observed that it was continuous with porphyroblast inclusion trails (cf. Fig. 15) and because he did not observe any relics of an older foliation. Zwart (1979) described S, in the same area as a locally developed, gently dipping crenulation cleavage, which was associated with porphyroblast rotations. The

Flow plane

(S3)

1

2mm

Fig. 12. Line tracing of microstructural porphyroblast-matrix relationships interpreted to indicate local syn-D, reactivation of S,. D, “general shear” (pure- and simple-shear components) with an S,-parallel bulk flow plane is heterogeneously distributed as follows. In zero-strain (porphyroblasts) or low-strain domains, S, is oriented at a high angle to S, and F, crenulations may be observed. In high-strain areas the angle between Sz and S, is reduced and no crenulations are observed. In the transition from lowto high-strain zones individual crenulations fade out (arrow I) and F, microfolds progressively unfold (arrow 2). Note that S, is not subhorizontal at this location (Fig. 2). This is probably due to rigid block rotations during later D, folding (Aerden, 1893).

D. G.A. M. Aerden / Tectonophysics 238 (I 994) 139- I60

I I

.

BUlK FLOW PLANE (HORIZONTAL)

L

J

J

\

J

J3

Fig. 13. Conceptual diagram of the transition from suprastructure to infrastructure in the Lys-Caiilaouas massif. The intensity of D, “general shear” deformation increases with depth. At higher structural levels S, is crenulated, but towardodeeper levels of hi&er (shear) strain, Fj crenulations are obliterated due to extension and reactivation of S, (see Fig. 11). Porphyroblast microstructures are schematically shown in moderate- and high-strain domains and can be compared with Figs. 5 and 12, respectively. Porphyroblasts did not rotate relative to S, due to the effects of small-scale deformation partitioning as shown. Consequently, they regionally preserve the initial steep orientation of Sz as inclusion trails. The diagram is valid for smaller scales of observation as well. The depth axis also represents time as the deep structure represents a kinematically more advanced stage than the shallow structure.

same microstructural sequence was described by him and later workers for the Lys-Caillaouas massif (e.g., de Bresser et al., 1986; Lister et al., 1986; Kriegsman et al., 1989). As there is general

agreement that porphyrobiast inclusion trails in the Lys-Caillaouas and Garonne Dome massifs in these represent S, , the external schist&y massifs should also be considered S, (Figs. 5, 6, 8,

Table 1 L/W ratios of opaque porphyroblast inclusions. The number of measured grains is given between parentheses Sample

E-W sections

L/W ratio

N-S sections

L/W ratio

34.2.1

L = 26.6 pm (115) W = 12.2 pm (35)

e = 2.2

L = 25.6 pm (111) W = 9.2 pm (58)

e = 2.8

34.2.1

L = 33.3 pm W = 11.0 pm

(73)

e = 3.0

L = 32.0 pm (173) W = 9.9fim (56)

e = 3.2

48.6

L = 15.5 pm W = 6.5/1m

(74) (90)

e = 2.4

L = 23.8 Frn (114) W = 6.9 pm (108)

e = 3.4

46.2.1

L = 16.9 km W = 5.4pm

(85) (60)

e=3.1

L = 19.7 grn (92) W = 5.8 firn (60)

e = 3.4

Length (L) and width IW) of opw inclupioas measured in N-S and E-W sections. Inclusions are~20% more eiortgate in the N-S vertical direction than in the E-W direction.

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

153

Table 2 L/W ratios of opaque matrix grains aligned parallel to a syn-D, reactivated Sa. The number of measured grains is given between parentheses Sample

E-W sections

L/W ratio

N-S sections

L/W ratio

19.3

L = 25.9 pm W = 6.6 pm

(56) (38)

e = 3.9

L = 25.5 pm (38) W = 6.1 pm (33)

e = 4.2

48.6

L = 27.5 pm (111) W = 6.9 pm

e = 4.0

L = 30.4 pm (92) W = 6.9 pm (96)

e = 4.4

Length (L) and width (W) of opaque grains in the matrix aligned parallel to a reactivated S,, measured in N-S and E-W sections. Opaque grains are 9% more elongate in the N-S direction than in the E-W direction.

Caillaouas massif (S,), but also indicate that porphyroblast inclusion trails have a vertical preferred orientation also in this area (Fig. 15).

12, 15). Preliminary observations by this author in the Garonne Dome not only confirm that the dominant schistosity is the same as in the LysGARONNE COME reinleerpreted afterFwgel

CENTRAL LV!?-CAlLLAOUAS MASSIF

81 al. 1988

mshrdv

N

S

m

m

2000

2000

1000

1000

0

W

0

WESTERN

I’

LVS-CAIILLAOUAS Listeretal. 1988

S-N 2km

Zkm

N-S

YASSIF

d)

lKsFTrMEryIsElF

VandmE&hout1966

2km

Fig. 14. (a) Composite cross-section through the Lys-Caillaouas (LCM) and Garonne Dome (GD) massifs. Cross-section lines are given in Fig. 2. The structure of the LCM is continuous with that of the southern GD. S2 is interpreted as steepening upwards. The GD is an open antiform defined by the crenulation cleavage (S,). See text for further discussion. (b) Structure proposed ,by Lister et al. (1986) for the western part of the LCM, which is similar to the structure proposed here for the GD. (c) Interpretation of the “Orri Dome” by Hartevelt (1970), who inferred an “infrastructure dome” at depth (stippled pattern) to explain the attitude of S, and F, fold vergence at the surface. (d) Composite, schematic cross-section of van den Eeckhout (1986) showing the Hospitalet gneiss dome being cross-cut by S,. The vergences of F, and Fs folds are opposite on either side of the dome.

154

D.G.A.M. Aerden / Tectanophysics 238 (1994) 139-160

Returning to the work of Pouget et al. (19881, it appears that their interpretation of the main schistosity as S, was based on the observation of crenulated porphyroblast inclusion trails, the axial planes of which are approximately parallel to a single, external schistosity. From this they argued that F, crenulations were trapped within porphyroblasts early syn-D, and that continued S, development in the matrix completely destroyed S,. However, they failed to demonstrate

Fig. 15. Line-tracing of andalusite porphyroblasts (stippled pattern) from Cambro-Ordovician rock in the southern part of the Garonne Dome, at the location marked with an asterisk in Fig. 2. S, inclusion trails are subvertical and are continuous with the external main foliation. This is the same situation as in the Lys-Caillaouas massif (compare with Figs. 8 and 12). The lack of F3 crenulations in the matrix, despite the reiatively steep dip of S,, may reflect a hbh vorticity of the buik flow (large shear-strain components). This would have caused rapid rotation of S, into an extensional orienriltion without much earlier crenelation. Neverth&ss, tracts of S3 can be observed adjacent to porphyrobtasts edges, where D, strain is locally increased.

that the inclusion trails (Si) represent an older foliation than the external foliation (S,), for example, with relics of a transposed matrix foliation, or by showing truncation of Si by S,. In contrast, this author observed the same microstructures as described by Pouget et al. (19881, in which continuity between S, and Si could be demonstrated (Fig. 8). Such microstructures indicate that porphyroblasts overgrew S, when this foliation was oriented at a high angle to the D, flow plane and was being crenulated (Fig. 11). The lack of S, in the matrix can be explained by subsequent rotation of S, into an extensional orientation causing its progressive decrenulation and reactivation (Figs. 11 and 13). Although considering the main schistosity in the Garonne Dome and Lys-Caillaouas massifs to be S,, this author observed that in zones where S, is intensely crenulated, S, may locally form the macroscopic cleavage (cf. Verhoef et al., 1984; van den Eeckhout, 1986). Such transposition zones can be interpreted as domains where D, strain was locally more coaxial and where S, maintained a shortening orientation throughout the D, deformation history. Also S, has been reported to form the dominant macroscopic foliation in some parts of the “suprastructure” of the Garonne Dome (Matte, 1969, his S,). However, both in D,- and D,-transposition zones S, remains distinguishable at least in thin section. When the dominant foliation on the cross-section of Pouget et al. (1988, fig. 1) is reinterpreted as S, and projected to the cross-section of the Lys-Caillaouas massif (Fig. 14a), a consistent structure emerges. S, in both massifs is axial planar to the dominant folds and is overprinted by a gently dipping crenulation cleavage (S,), which is particularly well developed in the centre of the Garonne Dome (Zwart, 1979; Pouget et al., 1988). Instead of defining a dome, the main foliation in the Garonne Dome ($1 is interpreted to steepen upwards, as observed in the Lys-Caillaouas massif (Fig. 3; Kriegsman, 1989b). In the present interpretation, the Garonne Dome is thus not defined by the main schism&y but by the less pronounced S, crenulation cleavage (Fig. 14a). The centre of this Garonne Dome separates zones with opposite dip directions of S,, opposite cren-

D.G.A.M. Aerden / Tectonophysics 238 (1994) X39-160

ulation asymmetries and opposite relative porphyroblast-matrix rotation senses (Fig. 15; Hassing, 1985; Pouget et al., 1988). As outlined earlier, these elements record rotation of S, from subvertical to inclined positions during vertical compression (Figs. 5, 8, 12, 15).

155

by S, based on similar observations as Pouget et al. (1988), namely, a single foliation in the matrix but F, crenulations preserved within porphyroblasts. However, without having documented the detailed relationship between Si and S,. this observation does not distinguish between the external foliation being a reactivated S, or S:,.

8.2. Western Lys-Caillaouas massif 8.3. Hospitalet massif Lister et al. (1986) described a directly comparable structure as proposed here for the Lys-Caillaouas-Garonne Dome area, in the western Lys-Caillaouas massif (Fig. 14b), which they also interpreted in terms of bulk vertical shortening. At deeper levels they inferred transposition of S,

The Hospitalet massif is of particular interest in the present study, because: (1) a gneiss dome is exposed at the base of the Cambro-Ordovician metasediments; (2) the massif as a whole represents a deeper structural level than the areas

01 F2

F2 gneiss

dome

Fig. 16. Kinematic model for erogenic extension (D,) in the central F’yrenees based on the numerical model for inhDmogeneous “general shear” of Malavieille and Taboada (1991). Gravitational collapse is proposed as the driving force of the extension and relatively rigid F, gneiss anticlines are inferred to have controlled the regional partitioning of this deformation. Isograds and SJ (subhorizontal lines) were draped around the gneiss dome, whereas S, was deflected away from it. F, crenulation and fold asymmetries as well as relative porphyroblast-matrix rotations are opposite on either side of the dome and indictate opposite horizontal shear-strain components (see insets). Depending on the amount of strain, either of the two shown dorphyroblast microstructures are encountered (see also Fig. 13), both preserving a subvertical S, as inclusion trails. Zones of opposike horizontal shearing are separated by a zone of coaxial strain at the dome centre. Granites intruded the dome late-D,, during ot post-D,.

156

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

considered so far; and (3) its structure can be directly extrapolated below the Garonne Dome area, along the general E-W structural trend of the Axial Zone of the Pyrenees (Fig. 1). Van den Eeckhout (1986) in an extensive study of the Hospitalet massif, showed that the exposed gneiss dome in this massif is cross-cut by a subhorizontal foliation (S,), which he showed formed during vertical shortening (cf. van den Eeckhout and Zwart, 1988; van den Eeckhout, 1990). He further observed that the vergence of both parasitic F,- and F,-folds were opposite on either side of the E-W long axis of the dome (Fig. 14~1, suggesting a possible D, origin of the gneiss dome as an F, anticline. The changes in structural asymmetry across the Hospitalet gneiss dome are analogues to those recorded on either side of the metamorphic Garonne Dome and suggest the presence of a not exposed gneiss dome in that area as well. In the next section a kinematic model will be proposed that describes the development of the Hospitalet and Garonne Dome massifs in a single crustal cross-section.

14d); and (3) the opposite D, shear senses on either side of the domes as indicated by porphyroblast-matrix relationships. A primitive version of the here proposed model was already advanced by Hartevelt (1970) who inferred relative rise of a not exposed “infrastructure” dome below the Orri Dome region to explain the fold and cleavage pattern at the surface (Fig. 14d). The domes in S, and the isograds in the Hospitalet and Garonne Dome massifs can be explained by assuming that these elements were initially subhorizontal and became draped around relatively rigid gneiss/basement domes during D, vertical shortening. D4 folding would have further tightened and modified the D, domes. As the metamorphic Garonne Dome occupies only the southern part of the stratigraphic Garonne Dome window (Fig. 2), later uplift of the metamorphic dome is suggested, relative to a less metamorphosed northern block, probably during D4 and/or Alpine deformation.

10. Discussion 9. Kinematic model for extension in the central Pyrenees A Hospitalet massif type gneiss dome at deeper levels in the Garonne Dome, with relatively rigid behaviour as compared to its metasedimentary cover, would exactly predict the deformation pattern observed at the surface, as illustrated in Fig. 16. The kinematic model of Fig. 16 envisages inhomogeneous crustal-scale vertical shortening, driven by gravitational forces and is based on the numerical strain model of Malavieille and Taboada (1991). Due to a more rigid behaviour of pre-existing F, miss anticlines, deformation is partitioned into doma~ins with opposite horizontal shear sense, separated by a zone of coaxial strain. The initially vertical S, is symmetrically deflected away from the dome centre during this deformation. This model predicts: (1) the attitude of S, and S, in the Hospitalet, Garonne Dome and Lys-Caillaouas massifs (Fig. 14a); (2) the vergence of parasitic F, folds being towards and of F3 folds being away from the dome centres (Figs. 14a and

10.1. Diapirism or gravitational oqenic

cohbpse

The symmetrical disposition of the S, and S, on either side of the Bosost leucogranite body (Fig. 14a) led Pouget et al. (1988) to relate S, to diapiric rise of a granite and/or gneiss dome (cf. Soula, 1982). Their main argument was that the dome-shape in S, is not observed in S, and, consequently, S, was “born” in a dome shape. However, in the present structural interpretation, envisaging upward steepening of the main cleavage, this argument would no longerbe valid. The is that second argument of Pougetet .aI. (I D, strain in the centre of the Garcwtlae Dome is characterized by relatively uniform extension, which is consistent with diapiric rise below this zone. However, this is also consistent tational c&apse around the dome as this paper. Furthermore, S, and S, dip teiivards the Lys-Caillaouas granite, which are the opposite relationships as observed near the r@n granite body in the Garonne Dome and hew, ineonsistent with diapirism (Figs 14a). The fact that S,

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160

is developed throughout the Lys-Caillaouas and Garonne Dome massifs (Fig. 14a), as well as in many other Pyrenean massifs, suggests relatively homogeneous deformation at the scale of the orogen, rather than local horizontal stretching above diapirs. 10.2. Origin of the Hospitalet gneiss dome Van den Eeckhout (1986) also argued that the attitude of cleavages around the Hospitalet gneiss dome, in particular of S, (Fig. 14d), refuted an origin of the dome by diapirism or by fold interference. Considering the by him recognized switch in vergence of parasitic F, folds across the dome (Fig. 14d), it is somewhat surprising that he did not interpret the dome as an F, anticline. However, he favoured an extensional origin (syn-D,) of the dome as a crustal-scale boudin or extensional flexure, because the intensity of S, increases towards the dome, which would suggest a genetic link. However, progressive flattening strain during D,, in combination with relatively rigid behaviour of a pre-existing (D,) dome, would also be consistent with intensification of D, strain against the dome. Additionally, extension and reactivation of S, could have partially destroyed S, adjacent to the dome (Fig. 16). The non-cylindricity of the domes can be readily explained as a result of inhomogeneous D, compression with a vertical stretching lineation and/ or by refolding effects of later deformation phases. In contrast, a D, origin of the Hospitalet gneiss dome and, potentially of other Pyrenean gneiss domes, is supported by the attitude and asymmetries of D, and D, structures in the here considered massifs (Fig. 16). Such an F, gneiss dome would have been opened to some degree during erogenic extension (D,) and again tightened during D, compression. 10.3. Gravitational collapse of the Variscan orogen This study consolidates previous work showing that crustal thinning and horizontal foliation development followed horizontal shortening and thickening of the Variscan orogen (e.g., Echtler and Malavieille, 1990; Malavieille et al., 1990;

157

Vissers, 1992). Together with a flattening type of D, strain in the Lys-Caillaouas, Garonne Dome and Hospitalet massifs, this suggests gravitational collapse of a thickened orogen as a probable cause of this erogenic extension (Vissers, 1992). Three possible causes of the dramatic collapse of the Variscan orogen are thought of. (1) A sudden change in plate motion that changed the internal stress state of the belt (e.g., Hoogerduijn Strating, 1994). (2) Detachment of a lithospheric slab could have produced a thermal pulse, increased uplift rates and, consequently, gravitational collapse (e.g., Platt and Vissers, 1989). (3) If temperatures within the orogen increased slowly as compared to the initial rates of crustal thickening, the orogen could have reached elevations that could no longer be supported at a time that rising temperature caused a critical level of internal strain softening. The coincidence of porphyroblast growth, peak metamorphic conditions and the very earliest stages of S, development in the Lys-Caillaouas massif, when very little crustal thinning could have been accomplished yet, favours a causal relationship, thus the second and/or third mechanism. The ensuing crustal thinning and decompression, with possibly an increased heat flow from the mantle due to lithospheric detachment, was probably responsible for partial melting of the deeper crust, emplacement of lateqorogenic granite plutons and volcanic activity (Innocent et al., 1994). 10.4. Position of the Pyrenees in the Variscan foldbelt

Theoretically, an erogenic belt collapsing under its own weight should spread radially outwards, following surface elevation gradients. This has been, for example, proposed for the Alpine orogen in the Alboran Sea region Watt and Vissers, 1989). The topographic slope of the orogen generates horizontal lithostatic pressure gradients and horizontal forces within the orogen, which compete with the external forces applied by the converging lithospheric plates. &tension directions in a collapsing orogen can’ thus be expected to vary as a function of the position

1%

D.G.A.M.

Aerden / Tectonophysics

within the orogen. At the orogen centre shortening would be ideally uniform, whereas outwards tectonic transport and an increasing difference between the X and Y strain axes (horizontal) would be expected. If the orogen encounters less resistance to gravitational spreading in lateral direction than in the direction of plate convergence, orogen-parallel extension will dominate. Basement culminations within the orogen do not necessarily coincide with the highest surface elevation and will therefore not determine the orogen-scale extension directions. However, as shown in this paper, they probably cause a regional partitioning of orogen-scale deformation (Fig. 16). The presented preliminary strain data indicate a relatively uniform flattening type of strain in the study area during a phase of erogenic collapse, albeit with slightly greater extension in E-W direction. Together with the abundance of Variscan granite, this suggests an internal position of the Pyrenees in the Variscan foldbelt. However, the Pyrenees are generally regarded as having a southern, external position in the Variscan foldbelt (Matte, 1991). A possible explanation is that the Variscan foldbelt was composed of two or more suborogens with largely independent, although interacting dynamic behaviour.

238 (1994) 139-160

D, strain in the Lys-Caillaouas massif, with a vertical shortening axis and somewhat greater extension in E-W direction than in N-S direction. Comparable extension has been previously determined in other Pyrenean massifs. The fact that crustal extension appears to have followed crustal thickening in many massifs of the Variscan foldbelt suggests gravitational collapse of an over-thickened orogen. The onset of D, deformation in the Lys-Caillaouas massif coincides with peak metamorphic conditions and porphyroblast growth, suggesting that such a gravitational collapse was possibly triggered by thermally induced internal strain softening.

Acknowiedgements This study has been financed by post-doctoral fellowship MRT 91 R 0509 from the French Ministere de la Recherche et de la Technologie. I would like to thank M. Mattauer for providing additional thin sections and for he&pfuI discussions. I also thank J. Malavieihe, M. SCranne, Ph. Matte, J.-M. Tubia and an anonymous referee for their most appreciated comments and suggestions.

11. Conclusions The attitude of S, and S, foliations and switching microstructural asymmetries on a section through the Lys-Caillaouas and Garonne Dome massifs reveal large-scale coaxial vertical shortening, partitioned into zones of non-coaxial deformation with opposite shear senses. This deformation partitioning was probably controhed by F2 gneiss domes at depth, which behaved relatively rigid as compared to their cover. During the extension horizontal isograds in the cover were draped around the gneiss domes, producing concentric isograd patterns at the surface, which were modified by later Variscan and Alpine compression. D, boudinage of quartz veins in N-S and E-W directions and preliminary strain analysis on opaque minerals, indicate a flattening type of

Aerden, D.G.A.M., 1991. Foliation boudinage control on the formation of the Rosebery Pb-Zn orebody, Tasmania. J. Struct. Geol., 13: 759-755. Aerden, D.G.A.M., 1995. Porphyroblast non-rotation during crustal extension in the Variscan Lys-CaiHsotias Massif, Pyrenees. J. Struct. Geol., in press. Bell, T.H., 1981. Foliation development: the contribution, geometry and significance of progressive bulk inhomogeneous shortening. Tectonophysics, 75: 273-296. Bell, T.H., 1985. Deformation partitioning and porphyrobiast rotation in metamorphic rocks: a radical reinterpretation. J. Metamorph. Geol., 3: 109-118. Bell, T.H., 1986. Foliation development and refraction in metamorphic rockq reactivation of earlier fohations and decrenulation due to shifting patterns of&formation partitioning. J. Metamorph. Geoi., 4: 421-444. Bell, T.H. and Cuff, C., 1989. D&solution, solution transfer, diffusion vs fluid tlow and volume loss during deformation/metamorphism. J. Metamorph. GeoL, 7: 425-447.

D.G.A.M. Aerden / Tectonophysics 238 (1994) 139-160 Bell, T.H. and Johnson, S.E., 1989. Porphyroblast inclusion trails: the key to orogenesis. J. Metamorph. Geol., 7: 279-310. Carreras, J. and Cires, J., 1986. The geological significance of the western termination of the M&ens Fault at Port Veil (Central Pyrenees). Tectonophysics, 129: 99-114. Cavet, P., 1957. Le Paleozoique de la Zone Axiale des Pyretrees orientales franqaises entre le Rousillion et I’Andorre. Bull. Serv. Carte Giol. Fr., 55/254: 303-518. Corstanje, R., Klepper, C., Rutgers, B., van der Wal, I.J. and van den Eeckhout, B., 1989. Quantification of finite strain in the Pyrenean Slate Belt; a first assessment using R,/+, method. Geol. Mijnbouw, 68: 313-322. Davis, G.A., Lister, G.S. and Reynolds, S.J., 1986. Structural evolution of the Whipple and South Mountains shear zones, southwestern United States. Geology, 14: 7-10. De Bresser, J.H.P., Majoor, F.J.M. and Ploegsma, M., 1986. New insights in the structural and metamorphic history of the western Lys-Caillaouas massif (Central Pyrenees, France). Geol. Mijnbouw, 65: 177-187. Den Brok, S.W.J., 1989. Evidence for pre-Variscan deformation in the Lys-Caillaouas area, Central Pyrenees, France. Geol. Mijnbouw, 68: 377-380. De Saint Blanquat, M., 1990. Petrological arguments for high-temperature extensional deformation in the Pyrenean Variscan crust (Saint Barthilemy Massif, Aribge, France). Tectonophysics, 177: 245-262. De Saint Blanquat, M., 1994. Late Variscan crustal thickening, thinning and rifting in the Pyrenees (abstr.). In: J. Malavieille and M. Seranne (Editors), Late Orogenic Extension in Mountain Belts. Dot. BRGM, 219: 169. Dewey, J.F., 1988. Extensional collapse of orogens. Tectonics, 7: 1123-1139. Echtler, H. and Malavieille, J., 1990. Extensional tectonics, basement uplift and Stephano-Permian collapse basin in a late Variscan metamorphic core complex (Montagne Noire, Southern Massif Central). Tectonophysics, 177: 125-138. Elliott, D., 1976. The energy balance and deformation mechanisms of thrust sheets. Philos. Trans. R. Sot. London, A283: 289-312. Garcia Sansegundo, J., 1992. Estratigrafia y estructura de la zona axial pirenaica en la transversal de1 Valle de Aran y de la Alta Ribagorca. Ph.D. thesis, Universidad de Qviedo. Publicaciones Especiales de1 Instituto Technologico Geominero de Espana, 339 pp. Garcia Sansegundo, J. and Alonso, J.-L., 1989. Stratigraphy and structure of the southeastern Garonne Dome. Geodin. Acta, 3(2): 127-134. Gibson, R.L., 1991. Hercynian low-pressure-high-temperature regional metamorphism and sub-horizontal foliation development in the Canigou massif, Pyrenees, Franceevidence for crustal extension. Geology, 19: 380-383. Guitard, G., 1970. Le metamorphisme hercinien mesozonal et les gneiss oeilles du massif du Canigou (Pyrbnees Orientales). M&m. BRGM 63, 353 pp. Hanmer, S.K. and Passchier, C.W., 1991. Shear sense indicators: a review. Geol. Surv. Can., Pap. 90-17, 72 pp.

159

Hartevelt, J.J.A., 1970. Geology of the upper Segre and Valira Valleys, central Pyrenees, Andorra/Spain (Ph.D. thesis). Leidse Geol. Meded., 45: 167-236. Hassing, L.J., 1985. De geologie van de noordelijke Garonne dome tussen Bosost en Melles (Centrale Pyreneeen, Spanje/Frankrijk). Unpublished MSc. thesis, ‘University of Utrecht, 81 pp. Hayward, N., 1992. Microstructural analysis of the classical spiral garnet porphyroblasts of south-east Vermont: evidence for non-rotation. J. Metamorph. Geol., 1Ck567-587. Hoogerduijn Strating, E.H., 1994. Extensional faulting in an intraoceanic subduction complex-a working hypothesis for the Palaeogene of the Alps-Appenine system. In: M. Stranne and J. Malavieille (Editors), Late Orogenic Extension. Tectonophysics, 238: 225-273 (this volume). Innocent, C., Brique, L. and Cabanis, B., 1994. Sr-Nd isotope and trace element geochemistry of the late Variscan volcanism in the Pyrenees: magmatism in post-erogenic extension? In: M. Seranne and J. Malavieille (Editors), Late Orogenic Extension. Tectonophysics, 238: 161-182 (this volume). Kriegsman, L.M., 1989a. Deformation and metamorphism in the Trois Seigneurs massif, Pyrenees-evidence against a rift setting for its Variscan evolution. Geol. Mijnbouw, 68: 335-344. Kriegsman, L.M., 1989b. Structural geology of the Lys-Caillaouas massif, central Pyrenees. Evidence for large-scale recumbent folding of late Variscan age. Geodyn. Acta, 3(2): 163-170. Kriegsman, L.M., Aerden, D.G.A.M., Bakker, R.J., den Brok, S.W.J. and Schutjens, P.M.T.M., 1989. Variscan tectonometamorphic evolution of the Eastern Lys-Caillaouas massif, Central Pyrenees-evidence for late-orogenie extension prior to peak metamorphism. Geol. Mijnbouw, 68: 323-333. Lister, G.S., Boland, J.N. and Zwart, H.J., 1986. Step-wise growth of biotite porphyroblasts in pelitic schists of the western Lys-Caillaouas massif (Pyrenees). J. Struct. Geol., 8: 543-562. Malavieille, J. and Taboada, A., 1991. Kinematic model for postorogenic basin and range extension. Geology, 19: 555558. Malavieille, J., Guihot, P., Costa, S., Lardeaux J.M. and Gardien, V., 1990. Collapse of the thickened crust in the French Massif Central: Mont Pilat extensional shear zone and St. Etienne Late Carboniferous basin. Tectonophysics, 177: 139-149. Matte, P., 1969. Le problbme du passage de la schistositt horizontale B la schistosite verticale dans le dome de Garonne (Paleozoic des Pyrentes Centrales). C. Rend. Acad. Sci. Paris, 268: 1841-1844. Matte, P., 1991. Accretionary history and crustal evolution of the Variscan belt in Western Europe. Tectonophysics, 196: 309-337. Molnar, P. and Lyon-Caen, H., 1988. Some simple physical aspects of the support, structure and evolution of mountain belts. Geol. Sot. Am., Spec. Pap., 218: 179-207.

160

1%G.A.M. Aerden / Tecronophysics 238 (1994) 139-160

Platt, J.P., 1987. The uplift of high-pressure-low-temperature metamorphic rocks. Philos. Trans. R. Sot. London? A 321: 87-103. Platt, J.P. and Vissers, R.L.M., 1989. Extensional collapse of thickened continental lithosphere: a working hypothesis for the Alboran Sea and Gibraltar arc. Geology, 17: 540543. Pouget, P., Lamaroux, C. and Debat, F., 1988. Le dbme de Bosost (Pyr&es centrales): r&nterpr&ation majeure de sa forme et de son &olution tectonom6tamorphic. C. Rend. Acad. Sci. Paris, 307(H): 949-955. Ramberg, H., 1977. Some remarks on the mechanism of nappe movement. Geol. F&en. Stockholm FGrh., 99: 110117. Royden, L., Horvath, F. and Rumpler, J., 1983. Evolution of the Pannonian Basin, 1. Tectonics. Tectonics, 2: 63-90. Santanach Prat, P.F., 1972. Sobre una discordancia en el Paleozoic0 inferior de1 Pirineo entre La Cerdaiia y el rio Ter. Acta Geol. Hisp., 5: 129-132. Stranne, M. and SCguret, M., 1987. The devonian basins of western Norway: tectonics and kinematics of an extending crust. In: M.P. Coward et al. (Editors), Continental Extensional Tectonics. Geol. Sot. London, Spec. Publ., 28: 537548. Saliva, J., Salel, J.F. and Brunei, M., 1989. Shear deformation and emplacement of the gneissic Canigou thrust nappe (Eastern Pyrenees). Geol. Mijnbouw, 68: 357-366. Soula, J.-C., 1982. Characteristics and mode of emplacement

of gneiss domes and plutonic domes in central-eastern Pyrenees. J. Struct. Geol., 4: 313-342. Van Bemmelen, R.W., 1972. Geodynamic Models; an Evaluation and Synthesis. Elsevier, Amsterdam, 265 pp. Van den Eeckhout, B., 1986. A case study of a mantled gneiss antiform, the Hospitalet massif, Pyrenees, Andorra, France (Ph.D. thesis). Geol. Ultraiectina, 45: 1-193. Van den Eeckhout, B., 1990. Evidence for large-scale recumbent folding during infrastructure formation in the Pyrenees: the structural geology of part of the eastern Hospitalet massif. Bull. Sot. Gtol. Fr. (8), VI(2): 331-338. Van den Eeckhout, B. and Zwart, H.J., 1988. Hercynean crustal-scale extensional shear zone in the Pyrenees. Geology, 16: 13.5-138. Verhoef, P.N.W., Vissers, R.L.M. and Zwart, H.J., 1984. A new interpretation of the structural and metamorphic history of the western Aston massif (central Pyrenees, France). Geol. Mijnbouw, 63: 399-410. Vissers, R.L.M., 1992. Variscan extension in the Pyrenees. Tectonics, 11: 1369-1384. Wickham, S.M. and Oxburgh, E.R., 1986. A rifted teetonic setting for Hercynian high-thermal gradient metamorphism in the Pyrenees. Tectonophysics, 129: 53-69. Zwart, H.J., 1979. The geology of the central Pyrenees. Leidse Geol. Meded., 33: 191-254. Zwart, H.J., 1986. The Variscan geology of the Pyrenees. Tectonophysics, 129: 9-27.