Eclogite-facies polyphase deformation of the Drøsdal eclogite, Western Gneiss Complex, Norway, and implications for exhumation

Tectonophysics 398 (2005) 1 – 32 www.elsevier.com/locate/tecto

Eclogite-facies polyphase deformation of the Drbsdal eclogite, Western Gneiss Complex, Norway, and implications for exhumation Ruth Foremana,*, Torgeir B. Andersenb, John Wheelera a

Department of Earth and Ocean Sciences, University of Liverpool, 4 Brownlow Street, Liverpool, UK b Institute for Geology and PGP, University of Oslo, PO Box 1047, Blindern, Oslo 0316, Norway Received 24 October 2003; accepted 1 October 2004

Abstract Exhumation of the deep crust during orogenic extension is accepted as a geological phenomenon, but structures formed during burial and the earliest stages of exhumation are often overprinted by deformation events occurring at shallower crustal levels. The Western Gneiss Complex (WGC) of Norway comprises variably retrogressed high-pressure (HP) and ultrahighpressure (UHP) eclogite bodies enclosed in predominantly felsic amphibolite-facies rocks. The Drbsdal body, located in the Sunnfjord area, is one of the largest and best-preserved eclogites in the WGC, and comprises over ca. 3 km2 of exposed eclogite-facies rocks. It is an excellent example of an eclogite tectonite, and displays a wealth of structures formed during deformation at a minimum depth of 50 km and peak temperatures of ca. 800 8C. Large volumes of mylonites with the eclogitefacies assemblage garnetFclinopyroxeneFquartzFamphiboleFclinozoisiteFphengiteFkyaniteFrutile are preserved. Structures associated with the eclogite-facies metamorphism include E–W-trending isoclinal folds, boudinage, hinge-parallel lineations, and meter-scale kyanite-dominated veins. The Drbsdal body was tightly folded on the kilometer scale under eclogitefacies conditions. Subsequently, the shapes of the Drbsdal body and other mafic bodies in the Sunnfjord area were modified during eclogite-to-amphibolite-facies E–W-directed stretching, to give boudin-like lenses. The timing of formation of a pervasive eclogite-facies lineation, eclogite-facies folds, and kyanite bearing veins overlapped substantially, and this portion of the deformation history was dominated by a constrictional strain field. Partitioning of deformation occurred after a substantial amount of eclogite-facies deformation had already taken place, and resulted in the relative rotation of linear features in specific zones on the scale of ~400 m. The eclogite-facies lineations and fold hinges within the Drbsdal body are subparallel to amphibolite and greenschist-facies structures throughout the WGC. Although circumstantial, this suggests that structures belonging to each of these three metamorphic facies formed during one progressive deformation event corresponding to extensional exhumation. Our observations are consistent with models involving inhomogeneous deformation of the lower crust

* Corresponding author. E-mail address: [email protected] (R. Foreman). 0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2004.10.003

2

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

during early orogenic extension. We envisage that transtensional (noncoaxial) deformation dominated the extensional deformation of the WGC, but that coaxial deformation must have occurred locally. D 2004 Elsevier B.V. All rights reserved. Keywords: Eclogite; Exhumation; Extension tectonics; Folds; Transtension; Western Gneiss Region

1. Introduction Products of high-pressure (HP) and ultrahighpressure (UHP) metamorphism are exposed in orogenic belts worldwide, and are formed during episodes of subduction and extreme crustal thickening (Chopin, 1987; Andersen et al., 1991; Carswell et al., 2003; Rey et al., 1997; Liou et al., 2004). Unravelling the dynamics of their formation and subsequent journey to the Earth’s surface is essential if we are to understand the processes operative in the deep crust

during orogenesis. Large volumes of UHP and HP rocks are exposed over an area of approximately 45,000 km2 in the Western Gneiss Complex (WGC), Norway (Fig. 1). The WGC is largely composed of Proterozoic quartzo-feldspathic gneisses of granitic to granodioritic compositions; minor amounts of anorthosites, mafic rocks, ultramafic rocks, and metasediments are also present (Tucker et al., 1990). Extensive reworking occurred during the Scandian phase of the Caledonian orogeny (Kullerud et al., 1986; Milnes et al., 1997), and amphibolite-grade assemblages now

Fig. 1. Simplified geological map of Sunnfjord and the surrounding area, SW Norway (compiled from maps of Engvik and Andersen, 2000; Krabbendam and Dewey, 1998).

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

dominate the WGC. The UHP and HP metamorphism is Scandian in age (420–400 Ma; Gebauer et al., 1985; Griffin and Brueckner, 1980), and occurred during the collision between Baltica and Laurentia. The presence of such deeply formed rocks at the Earth’s surface requires exceptional geological circumstances, and must involve exhumation along a P–T path that permits metastable behaviour of particular lithologies over a wide range of metamorphic conditions (Austrheim and Engvik, 1997; Rubie, 1990; Wayte et al., 1989). Assuming that erosion alone is not sufficient to expose HP/UHP rocks at the surface, a contribution from deformation-related processes must be considered. Moreover, the large amount of deformation involved in the transfer of HP/UHP rocks to the uppermost crust appears, in many cases, to have dismembered crustal fabrics and enhanced retrogression to lower-grade assemblages. Accordingly, the majority of felsic gneisses and some of the mafic material of the WGC have experienced retrogression to amphibolite-facies assemblages (Krabbendam and Wain, 1997). Even so, large and superbly wellpreserved HP, and smaller UHP bodies characterise the WGC. Exceptionally, some of these bodies have maintained eclogite-facies structural relationships in their interiors; these features are rarely seen in UHP provinces worldwide. Many recent models for exhumation of the WGC involve extensional deformation in the upper crust, coupled with an overall coaxial deformation in the lower crust (Andersen and Jamtveit, 1990; Andersen et al., 1994; Jolivet et al., 1994). Krabbendam and Dewey (1998) proposed that exhumation occurred via transtensional (noncoaxial) deformation. Rey et al. (1997) proposed that plate divergence caused extensional exhumation of the WGC, and was triggered by the Variscan collision between Laurasia and Gondwana. There are two main issues to resolve concerning the HP and UHP rocks of the WGC: (1) exhumation of the HP rocks must be explained, and (2) links between the HP and UHP and their juxtaposition must be explained. Possible explanations for the close juxtaposition of preserved HP and UHP assemblages include those of Wain (1997) and Terry et al. (2000a,b). Wain (1997) defined a UHP province, using observations of relic coesite and pseudomorphs after coesite as evidence for UHP metamorphism of eclogites along with the surrounding gneisses, or din

3

situ.T In this scenario, the boundary between this UHP province and the HP portion of the WGC is interpreted as a folded or imbricated tectonic break, along which pressure jumps of ~3 kbar are recorded, giving a bimodal trend in P–T estimates in this dmixed zone,T and the HP and UHP units were juxtaposed prior to a late phase of overprinting. The model of Terry et al. also focuses on explaining the close juxtaposition of UHP and HP rocks (~40 and ~20 kbar, respectively), and follows Wain (1997) in separating the UHP and HP rocks into two distinct tectono-stratigraphic units with different P–T histories. This model can be summarised as a two-stage exhumation history in which the HP and UHP rocks originally reside in different, distinct crustal blocks separated by a major structural discontinuity. During the initial stage, the UHP block is exhumed from a maximum depth of 125 to ~60 km, and top-southeast thrusting and imbrication brings the UHP rocks into contact with the HP block. The second stage involves exhumation of the juxtaposed UHP and HP blocks to a depth of ~40 km, accompanied by reequilibration and top-to-west extensional shearing. Both models have drawn heavily on studies of the mineral chemistry, structure, and regional P–T distribution of the UHP and HP rocks. More recently, evidence of UHP metamorphism has been found in mafic bodies over a much broader area than previously anticipated, and UHP indicators have even been found in the Verpeneset body, hitherto considered a type example of an HP WGC eclogite (Carswell et al., 2003). In the light of this new evidence, it may be necessary to reappraise the models. Perhaps mixed HP and UHP rocks occur over a larger area than can be accounted for by models involving the tectonic juxtaposition of several distinctly different units; the possibility that there is no major tectonic break between bodies recording HP and UHP metamorphism in the WGC must be considered. The observation of a large coherent mixed zone in which mafic bodies preserve UHP or HP or amphibolitefacies metamorphism depending on their shapes, sizes, compositions, and distances from major detachment horizons is perhaps a closer match to the evidence. The implication is that metastability and nonreaction as documented in a number of papers on the eclogites of the Lind3s nappe in the Bergen area as well as in the WGC (cf. Austrheim and Engvik, 1997) may be more important than previously suggested, and it is therefore

4

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

difficult to explain the variable maximum pressure record of a large portion of the WGC by tectonic juxtaposition as discussed in previous papers (Terry et al., 2000a,b; Wain, 1997; Wain et al., 2001). Regardless of the details, the history of the UHP rocks of the WGC cannot be understood without reference to the HP history. The Drbsdal mafic body, in Sunnfjord (Fig. 2), is the focus of this study. It is an exceptional example of an HP tectonite because it displays several generations of mesoscale and macroscale structures that were formed at different times but all within eclogite-facies conditions. It is among the largest HP occurrences in the WGC, yet the body had not been mapped in detail until this study. In fact, it is, to our knowledge, one of the largest pristine eclogite bodies recorded worldwide. For example, the largest known eclogite body in the Dabie Shan HP/UHP terrane is also an HP body, and is 200 m1 km in extent (Castelli et al., 1998; L.

Jiang, personal communication). The Cabo Ortegal eclogite is the only recorded larger occurrence and is considerably larger (100 m thick and 17 km long; Abalos, 1997), but is heavily overprinted. Exceptional preservation of the V3rdalsneset eclogite body, a nearby HP tectonite, has been previously reported by Engvik and Andersen (2000). The V3rdalsneset eclogite body is similar to the Drbsdal body in many aspects of its petrography and has some similar structural characteristics (Engvik and Andersen, 2000); however, only a small area is exposed, and therefore the size and context of the body are unclear. In contrast, the whole of the Drbsdal body is more accessible and can therefore be placed in context. Structural analysis of the Drbsdal body is applied alongside petrological evidence, and reveals that this part of the WGC was both pervasively folded and extended on the kilometer scale during residence in the deep crust.

Fig. 2. Simplified geological map of Sognefjord–Dalsfjord region, SW Norway, showing the relationship of the Drbsdal mafic body (D) to related mafic bodies and regional-scale structures, including the V3rdalsneset eclogite body (V) (modified after Engvik and Andersen, 2000).

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

2. Geologic setting The Drbsdal mafic body is situated in Sunnfjord, western Norway, and is surrounded by amphibolitefacies granodioritic gneisses of the WGC (Fig. 1). The WGC constitutes the lowest exposed tectonic unit in the hinterland of the Norwegian Caledonides, and is characterised by the preservation of relatively small lenses of eclogite-facies material within large tracts of amphibolite-facies gneiss. The Drbsdal mafic body is situated in the southerly dipping limb of a regionalscale antiform (Hacker et al., 2003), with the lowgrade Devonian Solund and Kvamshesten Basins exposed in complimentary synforms ~5 km to the south and ~15 km to the north, respectively (Fig. 1). Orogen-scale extension on the Nordfjord–Sogn and Hornelen Detachment Zones (NSDZ and HDZ, respectively) is at least partially responsible for exhumation of the WGC, which now lies in its footwall (Fig. 3). The NSDZ is well exposed at key localities beneath the hanging wall. It has carried the hanging-wall basins down to the west, and in the Sunnfjord area, the detachment zone comprises ~3–5 km thickness of brittle–ductile mylonites (Andersen and Jamtveit, 1990). There is an indistinct transition with increasing depth from the NSDZ mylonites into the heterogeneous basement gneisses and eclogites of the WGC (Norton, 1987). The Drbsdal body belongs to a suite of mafic bodies that have been variably

5

tectonised, and range in character from relatively undeformed gabbros and coronitic eclogites (Engvik and Andersen, 2000; Mbrk, 1985), to true eclogites and eclogite tectonites of variable composition. The mafic bodies are included as pods and lenses in the amphibolite-facies gneisses (Figs. 2 and 3). These mafic bodies also preserve metastable mineral assemblages related to different stages in the P–T history of the WGC, including igneous material, granulite-facies assemblages, and variably deformed eclogites. In addition to P–T controls, fluid budget is thought to have had a large control on the initiation and progress of metamorphic reactions within these mafic bodies, and in turn on deformation. No precursors to the eclogite-facies rocks have been found within the Drbsdal mafic body, but the spatial distribution and petrographic characteristics of other mafic bodies in Sunnfjord suggest that these mafic bodies were intruded into felsic gneisses as a series of dykes, sills, and larger igneous complexes (Engvik et al., 2001), and passed through amphibolite- and granulite-facies before the onset of eclogite-facies metamorphism (Austrheim and Engvik, 1997; Engvik et al., 2001). Indeed, Krabbendam et al. (2000) and Wain et al. (2001) have demonstrated that some WGC rocks preserve Pre-Caledonian igneous and granulitefacies assemblages. Over 90% of the Drbsdal mafic body preserves eclogite-facies assemblages and structures, and retrogressive amphibolitisation is largely

Fig. 3. Schematic vertical section showing the structural relationships of the different crustal units in the Hyllestad–Sunnfjord area. Vertical distances calculated from maps presented herein and those presented by Hacker et al. (2003).

6

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

confined to zones of contact with the surrounding granodioritic gneisses.

3. Petrography and mineral chemistry In this section, our aim is to demonstrate that the rocks at Drbsdal preserve eclogite-facies assemblages; we will not give a comprehensive metamorphic history. 3.1. Petrography The interior of the Drbsdal body comprises rock of fine to medium grain size, and characteristically displays eclogite-facies assemblages with little or no evidence of retrogression. No protolith to the eclogitefacies rock has been observed in the Drbsdal body; eclogitisation was complete and accompanied by pervasive deformation. Garnet is always present in the eclogite-facies assemblage, along with omphacite, zoisite, quartz, and rutile. Amphibole, phengitic mica, and kyanite occur as accessory minerals in textural equilibrium with some or all of the eclogite-facies minerals. The assemblage varies considerably between different compositional bands, and can also vary within and between parts of the mafic body with different structural trends (Section 4). A gradual transition to a fine-grained green–black amphibolite is typically observed at the margins of the mafic body. This is interpreted to be a retrogressed variety of the eclogite. Over distances between 5 and 100 m, this mafic amphibolite passes into a lithology with the granodioritic amphibolite-facies assemblage quartz, plagioclase feldspar, potassium-rich feldspar, white mica, biotite mica, and, in some places, garnet and sphene. Brief descriptions of the rock types found within and around the Drbsdal mafic body and preliminary P–T estimates follow. Rock types are grouped

7

according to field observations, but descriptions also include optical microscope observations and chemical data obtained via microprobe analysis. Mineral analyses were performed on a Cameca SX100 electron microprobe (wavelength dispersive system) at the Department of Geology, University of Oslo, using natural and synthetic mineral standards. An accelerating voltage of 15 kV was used, with a beam current of 10 nA and counting time of 10 s for all analyses. A focused electron beam was used for all minerals except sheet silicates and feldspars, for which the beam was defocused to 10 Am diameter. Amphibole compositions are classified according to the nomenclature of Leake et al. (1997); clinopyroxene compositions are classified according to the nomenclature of Morimoto et al. (1988). Lithologies are grouped as follows: (1) gneissic eclogites, (2) partially retrogressed eclogites, (3) amphibolites, and (4) granodioritic amphibolite-facies rocks. Minor volumes of other rock types are found within the mafic body, and these are described briefly for completeness. 3.1.1. Gneissic eclogite The name gneissic eclogite covers the majority of rocks in the Drbsdal body, perhaps as much as 80% or 90% of the total volume (Fig. 4A and B). The gneissic eclogite is characterized by a strong compositional foliation and a strong pervasive type I lineation (Section 5.2). Within the gneissic eclogites, garnet, omphacite, quartz, zoisite, and rutile exist in textural equilibrium. Kyanite or amphibole may be present as part of an equilibrium assemblage, although may also have formed later. Garnet generally occurs as euhedral grains of up to 1 cm in size. Most garnet grains larger than 0.5 mm contain numerous inclusions of quartz and zoisite, and occasional pargasitic amphibole in their cores. Complex compositional zoning is a conspicuous feature of garnet in the gneissic eclogites, and is observed in all but the

Fig. 4. Photomicrographs showing typical mineralogy and textures in rocks from the Drbsdal eclogite body. Each pair shows the same area. Double-headed arrows show lineation direction. Field of view 4 mm width for all photomicrographs: (A) gneissic eclogite (sample P10) viewed in cross-polarised light (XPL); (B) same view in plane-polarised light (PPL); (C) partially retrogressed eclogite (sample R18a, viewed in XPL); (D) same view in PPL; (E) amphibolite (sample SK10, viewed in XPL); (F) same view in PPL; (G) granodioritic amphibolite-facies rock (sample OL6, viewed in XPL); (H) same view in PPL. Abbreviations: amphibole (Amp); biotite (Bt); diopside (Di); epidote (Ep); garnet (Grt); matrix of plagioclase, quartz, magnesiohornblende, alkali feldspar, and epidote (Mx); omphacite (Omp); phengite (Phg); plagioclase (Pl); quartz (Qtz); zoisite (Zo).

8

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

smallest of grains (b0.2 mm). Broadly, zoning varies from core (Alm42–56, Prp14–30, Grs20–27) to rim (Alm41–48, Prp26–39, Grs10–22) in a fashion commonly thought to represent prograde growth, but more detailed analysis reveals a much more complex pattern than can be explained by a simple model of prograde growth. Discussion of this problem is beyond the scope of the present work, and will be addressed in a separate study, although it is worthwhile to note that since any estimate of P–T assumes chemical equilibrium between the phases used in the calculation, P–T estimates for these rocks should be viewed with considerable caution. Omphacite grains (Jd29–45, Ae2–25; Na2O content 6.3–7.3 wt.%) generally show a strong shapepreferred orientation, and define a strong lineation, along with zoisite/clinozoisite and occasional kyanite. The pervasive centimeter-scale to meter-scale foliation defined by variations in the relative abundances and grain sizes of the eclogite-facies minerals is mimicked at the submillimeter scale. Kyanite in the gneissic eclogite is commonly associated with clinozoisite/zoisite-bearing layers, and it is in these layers that the shape-preferred orientation is strongest. Zoisite within these layers is strongly prolate in shape, and frequent fractures and zones with a high density of fluid inclusions, interpreted to be dhealedT fractures, transect zoisite grains at a high angle to the lineation direction. Omphacite grains are in many places surrounded by a rim of symplectite, especially so in quartz-rich samples. The symplectite consists of delicately intergrown diopside (Jd3–35, Ae1–14; Na2O content 1.48–6.6 wt.%) and plagioclase (Ab60–87, An13–25), and is interpreted to have grown during static posteclogitefacies conditions. Barroisitic amphibole often occurs in the matrix of the eclogite-facies assemblage. It is variably abundant, and generally shows a similar shape-preferred orientation to clinopyroxenes within the foliation. Commonly, poikiloblastic amphibole grains occur in the eclogite, with garnet (Alm40–43, Prp34–36, Grs19–20) as the only included phase. Kyanite occasionally occurs in these samples as large poikiloblasts, and carries inclusions of quartz, garnet (Alm41–45, Prp30–34, Grs17–24), and omphacite (Jd43–47, Ae0–4; Na2O content 6.5–7.0 wt.%) with a shape-preferred orientation subparallel to lineation.

Grains of phengitic mica are generally present in small amounts in most samples, and are aligned within the foliation. Phengite is more abundant in the margins of the mafic body and is particularly abundant in samples from the margins of eclogitefacies veins, especially those bearing kyanite and quartz. It is therefore likely that these phengites are products of eclogite-facies hydration reactions facilitated by the local availability of a fluid phase. Since additional potassium is also required, a potassium source such as a potassium-bearing metasomatic fluid is also required for new phengite to be produced. Rutile occurs as individual grains or strings of grains aligned parallel to lineation. 3.1.2. Partially retrogressed eclogites These rocks represent partially retrogressed lithologies occurring at the margins of the mafic body, and contain substantial amounts of minerals typical of the amphibolite-facies, along with minor amounts of greenschist-facies material (Fig. 4C and D). Samples typically contain garnet, plagioclase, diopside-rich clinopyroxene, amphibole, quartz, and phengite. Omphacite, zoisite, and rutile are present in small amounts and are interpreted to represent eclogitefacies remnants. Chlorite, epidote and sphene may be present as accessory minerals, and are retrogression products belonging to the amphibolite and possibly greenschist facies. The widespread presence of symplectites and reaction rims indicates a high degree of textural disequilibrium. Garnet grains (Alm48–58, Prp11–18, Grs20–24) in these rocks are zoned, indicating disequilibrium during retrograde metamorphism. They are generally subhedral to anhedral, and are in many cases surrounded by a kelyphitic mantle consisting of an inner rim of hastingsitic amphibole and an outer rim of chlorite. Grain size of the garnet varies from 0.3 mm to N1 cm, and larger grains have an assortment of inclusions. Amphibole (taramite to magnesiotaramite) and carbonate inclusions are common towards the rims of large grains. Zoisite and quartz inclusions are present in garnet cores. Omphacite (Jd34–42, Ae10–12; Na2O content 6.1– 7.5 wt.%) in the partially retrogressed eclogites is largely replaced by symplectic intergrowths of plagioclase (Ab87–93, An8–13) and diopside (Jd17–19, Ae9–11; Na2O content 3.7–4.2 wt.%). The extent of

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

symplectisation varies within and between samples on the millimeter scale. The widespread presence of delicate mimetic symplectite in these samples indicates static retrogression. Phengitic mica is more abundant than in the gneissic eclogites, and makes up as much as 15% of the rock in some samples. It is present as aggregates N2 mm in size, often oriented subparallel to the sample lineation, and wraps around large garnet grains, sometimes forming pressure shadows. Locally abundant phengite may signify the availability of a fluid phase and a potassium source to facilitate eclogite-facies hydration reactions (as hypothesized in Section 3.1.1), but does not strictly require the presence of a fluid phase since hydrous minerals can also exist when water activity is b1 (i.e., no free water). The abundant eclogite-facies phengite in these rocks may indicate that they were probably particularly hydrous under eclogite-facies conditions. Our interpretation is that the present assemblages and textures are the result of locally enhanced retrogression of marginal rocks with particularly hydrous mineral assemblages. 3.1.3. Amphibolites Amphibolites comprise the dark-coloured iron and magnesium-rich rocks associated with the Drbsdal mafic body (Fig. 4E and F). They occur at the contact of the eclogite-facies mafic rocks and the granodioritic gneisses surrounding them, and also as pods of b5 m2 found dfloatingT in the granodioritic material close to the Drbsdal mafic body. These rocks are often finegrained and in many cases contain plagioclase, amphibole, alkali feldspar, biotite, quartz, epidote, white mica, and sphene as part of the amphibolitefacies assemblage. Chlorite is occasionally present. The fabrics within these amphibolites vary from statically retrogressed symplectic and granoblastic fabrics after eclogite-facies textures to strongly lineated and mylonitic textures. Very similar assemblages and textures also characterise the fine-grained amphibolites developed at amphibolite-facies vein margins (symplectic and granoblastic) and in shear zones (lineated to protomylonitic) within the body (see Section 4.7 for structural description of shear zones). Garnet (Alm42–47, Prp22–37, Grs16–30) is largely replaced by chlorite, alkali feldspar, and magnesiohornblende. Relict grains are transected by

9

chloritised cracks and mantled by chlorite, alkali feldspar, and biotite. Retrogressive metamorphism has produced a medium- to coarse-grained matrix of plagioclase (Ab64–84, An15–37), quartz, magnesiohornblende, occasional alkali feldspar, and epidote, with occasional biotite aggregates. Rutile is commonly present, and grains are often mantled by a fine-grained aggregate of sphene. 3.1.4. Granodioritic amphibolite-facies rocks The granodiorites surrounding the mafic body (Fig. 4G and H) are made of plagioclase, quartz, alkali feldspar, biotite, white mica, and epidote. Chlorite is also commonly present, and garnet occurs in small amounts in the granodiorites closest to the eclogite body (approximatelyb100 m away). A felsic matrix of quartz, plagioclase (Ab85–91, An8–15), and alkali feldspar (Ab3–9, An0) constitutes between 65% and 85% of these samples, with the remainder made up of sheet silicates, garnet, and epidote. Where present, garnet (Alm51–64, Prp2–14, Grs27–41) grains appear to be in textural equilibrium with the amphibolite-facies matrix, and larger grains (N3 mm) contain inclusions of carbonates, sphene, zircons, and quartz. Biotite occurs in mimetic association with phengite, partially or completely replacing the phengite grains. Garnet porphyroclasts are often accompanied by quartz-filled pressure shadows and wrapped by sheet silicates to give a diffuse banding. 3.1.5. Other lithologies 3.1.5.1. Garnet–quartz layers. Occasional bands of garnet-rich rock occur within the eclogite body. Finegrained garnet of b0.5 mm grain size constitutes 70– 90% of the rock. Abundant inclusions of allanite, iron oxides, and rutile occur within garnet grains. The garnet forms a matrix with interstitial quartz, omphacite, phengite, rutile, and iron oxides. Rutile grains of up to 3 mm in diameter constitute 3–5% of the matrix, appearing as elongate grains with long axes parallel to the lineation of the gneissic eclogite. 3.1.5.2. Garnet–amphibole rocks. Rocks consisting almost entirely of coarse-grained amphibole and garnet occur locally at the margins of the mafic body. A key locality is on the eastern edge of Svardalsvatnet

10 R. Foreman et al. / Tectonophysics 398 (2005) 1–32 Fig. 5. Structural map of the Drbsdal eclogite body. Positions of detailed maps of Tinghaugen (see Fig. 10) and Teiges3ta (see Fig. 12), and cross-section line A–B (see Fig. 8) are indicated.

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

(Fig. 5), where the southern boundary of the mafic body is in contact with granodioritic gneiss. 3.1.5.3. Felsic bodies. Rocks of felsic composition make up approximately 20% of the mafic body, and consist largely of quartz and phengite with small amounts of garnet, rutile, and zoisite, and occasional omphacite. These felsic bodies are highly variable with regard to size and morphology, with dimensions in the range of 10 cm to N100 m length, and 2 cm to 15 m width. They are interpreted to have originated under eclogite-facies conditions, and their structural features are described in more detail in Section 4.5. 3.2. P–T estimates Temperature and pressure estimates have been calculated for the formation of the gneissic eclogite, according to the clinopyroxene–garnet thermometers of Ellis and Green (1979) and Powell (1985b), the garnet–amphibole thermometers of Powell (1985a) and Graham and Powell (1984), the garnet–clinopyroxene–phengite barometer of Waters and Martin (1993), and using the THERMOCALC program. Temperature calculations based on cpx–grt pairs for the two thermometers yielded either extremely large ranges (e.g., 420–715 8C for T max (Fet) for a single mineral pair), or no temperature at all. Large uncertainties are involved in the calculation of ferric iron contents for clinopyroxenes, especially in the specific compositional range of the omphacites studied here. These uncertainties arise because: (1) the calibrations of the calculations used for cpx–grt exchange thermometry are highly sensitive to variations in Fe2+ of omphacite, and (2) Fe3+ was estimated by stoichiometry and is therefore subject to further uncertainties related to analytical errors and nonstoichiometry. The temperature estimates obtained via cpx–grt thermometry for all methods are therefore rejected (Koons, 1984). Grt–amp thermometry using the methods of Graham and Powell (1984) and Powell (1985a,b) gives similar estimates of T max (Fet); 537–633 and 531–631 8C, respectively, identical within error. Pressures of 17F2 kbar were calculated using the calibration of Waters and Martin (1993) for grt–cpx– pheng grains in close proximity, assuming a temperature of 600 8C.

11

Using the THERMOCALC program, estimates of T=720–830 8C and P=19–21 kbar were obtained for gneissic eclogite samples. The following four assemblages were used in the THERMOCALC calculations: garnet+omphacite+zoisite+amphibole, garnet+omphacite+zoisite+amphibole+kyanite, garnet+omphacite+amphibole, and garnet+omphacite+zoisite+ phengite. THERMOCALC could not calculate P or T estimates for assemblages not containing amphibole or phengite due to a dlack of independent reactionsT. On a more cautionary note, uncertainties related to Fe3+ in omphacite and poorly understood complex zonation patterns in garnet from Drbsdal indicate that these P–T estimates may deviate substantially from the peak P–T conditions of these rocks. Relevant previous P–T work on eclogites in the Sunnfjord region (Fig. 2) includes estimates from the V3rdalsneset eclogite body (Engvik and Andersen, 2000), from localised eclogite-facies layers within the B3rdsholmen banded granulite-facies complex (Engvik and Andersen, 2000), and from smaller eclogite pods in the Lavik–Hyllestad area (Hacker et al., 2003). The V3rdalsneset rocks yielded estimates of T=677F21 8C and P=16F2 kbar using the cpx– grt geothermometer of Powell (1985b) and the geobarometer of Waters and Martin (1993), respectively; and the B3rdsholmen rocks gave estimates of T=455–490 8C after Powell (1985b) and P min=12 kbar using the method of Holland (1980), assuming T=500 8C. Labrousse et al. (2002, in preparation) recalculated these data using THERMOCALC, to give T=615F22 8C and a pressure of 22.7 kbar for V3rdalsneset, and T=525F46 8C and a pressure of 23.4 kbar for B3rdsholmen. The Lavik–Hyllestad samples gave estimates of ~700 8C and 23 kbar, using THERMOCALC (Hacker et al., 2003). All published pressure estimates obtained using THERMOCALC are higher than those obtained using exchange reactions. A possible reason for this is that THERMOCALC uses larger assemblages than the exchange reactions.

4. Structural relations The Drbsdal eclogite body is enveloped by granodioritic amphibolite gneisses (Fig. 5), and its E–W-trending long axis in map view is subparallel to

12

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

the linear trend of regional-scale late extensional features in this area (Andersen et al., 1994; Krabbendam and Dewey, 1998; Milnes et al., 1997). Structural mapping and analysis have allowed us to link grainand outcrop-scale structures with map-scale features, and thus infer the shape of the body in three dimensions. 4.1. Foliation and folding Throughout most of the Drbsdal mafic body, a strong compositional foliation is defined by variations in the modal abundances of the eclogite-facies minerals garnet, omphacite, quartz, zoisite, amphibole, and kyanite, or their amphibolite-facies equivalents in zones of retrogression. Compositional bands are typically b1 cm to ~1 m in thickness, and steeply dipping. Variations in orientation are associated with pervasive folding of the body. As shown in Figs. 5 and 6, consistent structural trends are observed within several map-scale domains (Section 4.3). In the eastern part of the body, the foliation trends ENE–WSW and is folded into tight to isoclinal

upright structures with wavelengths of 20 cm to 20 m (Figs. 5 and 6). These folds typically have ENE– WSW-trending subcylindrical hinges with plunges of 15–408 to the west, and can be traced laterally for distances of 3–10 m. All eclogite-facies folds lack axial planar cleavage or fabric. The folds affect the eclogites, and the amphibolitic and granodioritic gneisses in their immediate vicinity. The granodioritic gneisses surrounding the eastern end of the body contain a relatively weak ENE–WSW-oriented foliation, which is also subparallel to the lithological contact to the mafic body. Towards the western end of the body, the foliation is less steep. Its typical orientation is NNE–SSW in the interior, and close to E–W at the margins. Foliation within the granodioritic gneiss surrounding this part of the body also trends E–W, although structural features are somewhat weaker in the granodioritic lithologies. Pervasive folding of both the mafic body and the country rock is observed, and tight upright folds with subcylindrical hinges and wavelengths of 20 cm to 20 m are again typical. Fold hinges have a spread of orientations, varying from E–W-trending fold hinges

Fig. 6. Structural data for Drbsdal eclogite, subdivided into three (A–C) distinct structural zones as indicated by boxed domains. Stereonets plotted as lower hemisphere equal area projections.

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

13

Fig. 7. Field photographs of typical eclogite-facies folds from Drbsdal. (A) Isoclinal m-fold in gneissic eclogite. Note more felsic layers at top of photograph are folded more chaotically than mafic layers near pencil. Length of pencil, 15 cm. Profile view, looking west. (B) Large tight fold in gneissic eclogite. Length of hammer, ca. 80 cm. Profile view, looking west; outcrop dips gently towards viewer.

associated with plunges of 10–208 to the west, to N– S-trending fold hinges associated with plunges of 10– 208 to the south. The centimeter- to meter-scale folds of the Drbsdal body are commonly asymmetric, but local vergence is not always obvious. In the eastern end of the body, these folds are characterized by steep southerly dipping axial planes and tight to isoclinal shapes (Fig. 7A and B). Cross-sections through the body (Fig. 8) show that meter- and kilometer-scale folds are also asymmetric, and are also characterised by steep

southerly dipping axial planes, northward vergence, and tight fold shapes. 4.2. Lineations Two types of lineation can be distinguished at Drbsdal. Since these are somewhat similar in mineralogy and appearance, the two types will be referred to as types I and II. Type I lineations (Fig. 9) are strongly developed throughout the majority of the Drbsdal body, and are

Fig. 8. Vertical cross-section through eastern end of Drbsdal eclogite body (location shown on Fig. 5). Data shown as dip indicators. Vertical=horizontal scale. Height measured in meters above sea level.

14

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Fig. 9. Schematic block diagram showing the relationship between folds, eclogite-facies type I lineations, and boudinage in a typical block of the Drbsdal eclogite.

commonly defined by a grain shape fabric of the eclogite-facies minerals omphacite, zoisite, quartz, and/or kyanite. A polycrystalline lineation is also common, defined by elongate clusters or aggregates of omphacite, zoisite, quartz, or kyanite grains. Distinction between the grain shape and polycrystalline lineations is not always possible in the field, but observations at the thin section scale indicate that grain shape lineations defined by zoisite are stretching lineations (Section 5.2). In general, the eclogite has an LNS fabric; layering is compositional, and grains are only slightly elongate in sections perpendicular to lineation. Lineations measured in the amphibolite-facies granodioritic gneisses outside the mafic body are subparallel to those inside the mafic body, and are defined by polycrystalline alignments of elongated quartz or feldspar grains. As with the compositional foliation, lineations have a wide variety of orientations throughout the whole of the mafic body, but trends characteristic of specific structural domains can be distinguished at the map scale (Figs. 5 and 6). These lineations are interpreted as stretching lineations, and their modal orientation (in area C; shown in Fig. 6) coincides strongly with the regional linear E–W trend. Previously, the linear E–W trend has been reported in amphibolite-facies rocks (Andersen et al., 1994; Krabbendam and Dewey, 1998), but at Drbsdal, minerals defining the E–W-trending lineation clearly grew under eclogite-facies conditions, suggesting that the trend may have been initiated before the late amphibolite-facies stage (see discussion in Section 5.4). The lineations generally define a penetrative crystal fabric, and are aligned within the plane of the compositional foliation. In the western part of the body, at Teiges3ta and Tinghaugen (areas A and

B), the orientation of lineations in the eclogite is highly variable. In area A, for example, a progressive change in the orientation of lineations is observed from ENE–WSW-trending lineations with westward plunges to ESE–WNW lineations with eastward plunges (Figs. 5 and 6). However, the presence of centimeter-scale to kilometer-scale shear zones within and at the margins of the body complicates the structural trends further in this area (see Section 4.7). In the eastern part of the body, at Ramsgro, the observed lineations follow an ENE– WSW trend, and typically plunge 25–408 to the west or southwest. The second type of lineation (type II), described as a surface lineation, is observed locally, particularly a few tens of meters north of Svanetjorna (Fig. 10). It is defined by finely dcorrugatedT aggregates of micas, amphiboles, and zoisite crystals, and looks similar in many respects to the lineations described above. However, this type of lineation is not penetrative but confined to the surfaces of isolated foliation planes within the rock, and micas are conspicuously abundant on these planes. Linear mineral orientations of this type trend E–W, and plunge 5–258 to the east or west, and are therefore generally much shallower than the plunges of the dominant penetrative lineation (Fig. 11). Perhaps, the type II lineation is locally superimposed on the earlier fabric defined by type I lineations. 4.3. Map-scale variations in structure Map-scale observations indicate local disturbances to the east–west trend of mesoscale fold hinges and other linear features (Sections 4.1. and 4.2). Planar features such as foliation planes are also affected, locally becoming less steep. In the Teiges3ta area (Fig. 12), lineations, foliation, and fold hinges measured in the granodioritic rocks surrounding the Drbsdal mafic body exhibit subparallelism with the regional E–W trend of gneissic fold hinges, foliation, and lineations. The same is true of measurements from very marginal parts of the mafic body, but towards its interior, these are locally highly oblique to the regional E–W trend (Figs. 6 and 12). This gradual change in obliquity of linear structures towards the interior of the body indicates a rotation of the margins of the mafic body

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

15

Fig. 10. Structural map of the Tinghaugen area (see Fig. 5 for location).

relative to its interior in this area. Within the unrotated block, fold morphologies are less tight than those throughout the rest of the body; but otherwise, structures and mineralogies are typical for the Drbsdal eclogite. The observed lineation pattern can be explained by partitioning of strain into the rest of the body, and a resulting clockwise rotation relative to the Teiges3ta dblock.T The open folds at Tieges3ta would be preserved remnants within a low-

Fig. 11. Conceptual sketch of the relationships between surface lineation (depicted in black) and penetrative lineation (depicted in grey) in the Drbsdal eclogite.

strain lozenge. This shearing is compatible with the transtensional strain regime proposed by Krabbendam and Dewey (1998), who presented evidence that exhumation through amphibolite-facies occurred this way (see Section 5.7). A similar obliquity of structural features with respect to the regional E–W trend is observed in the Tinghaugen area (Fig. 10). Fold style at Tinghaugen is more chaotic than at Teiges3ta, at times being disharmonic, with poorly defined layering. Lithologies at Tinghaugen are much more retrogressed than at Teiges3ta, but thin section analysis shows that the zones of retrogression are dominated by symplectic material after eclogite-facies minerals such as omphacite and garnet. We conclude that retrogression of the rocks in the Tinghaugen area was largely static, so the obliquity was established at eclogite facies (as at Teiges3ta). Differences in deformation style between Tinghaugen and Teiges3ta may simply be the result of differences in bulk composition or distance from the upper or lower margins of the body (out of the map plane). The relationship between N–S-trending linear features at Tinghaugen and those at Teiges3ta is difficult to assess due to these fundamental differences

16

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Fig. 12. Detailed structural map of Teiges3ta area (see Fig. 5 for location).

in deformation style, and as a result of this, the timing of formation of structures in the two areas remains unresolved. 4.4. Boudinage Boudinage is most common in the eastern portion of the Drbsdal body and structures range from pinch-and-swell to completely detached boudins. The extension direction implied by the orientation of boudins trends SW–NE to E–W, and therefore coincides strongly with the trends of stretching lineation and fold hinges within the Drbsdal mafic body. Ductile deformation of less competent felsic layers is observed to have accommodated the boudinage of more competent mafic layers. Asymmetric truncation of the ends of layers within boudins is common, but both sinistral and dextral shear senses are observed, implying that no dominant shear sense is present (Goscombe and Passchier, 2003). Boudin necks are commonly sites of vein development (see Section 4.6). Both boudins and veins in boudin necks comprise eclogite-facies mineralogies.

4.5. Felsic bodies Rocks of felsic composition make up approximately 20% of the mafic body (see Section 3.1.5 for petrographic description). The felsic bodies themselves are often foliated, and some may also be traced around eclogite-facies folds. Some of the largest foliated felsic eclogite bodies, however, have a discordant cross-cutting relationship with the foliation in the surrounding mafic eclogite, and contain pods of mafic eclogite. Their discordant relationship with the surrounding eclogite indicates that they postdate the majority of eclogite-facies deformation, while their mineralogy indicates that they crystallized under eclogite-facies conditions. Little or no retrogression is observed either at the margins of these bodies or within the included pods (Fig. 13A and B). These features lead to the interpretation that most of these dyke-like felsic bodies were introduced under eclogite-facies conditions, possibly as a fluid phase or melt, and that pods and lenses of mafic eclogite were entrained during dintrusionT of these felsic bodies, but it is also possible that some are older than the eclogitefacies event. Many of these large felsic bodies are also

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

17

Fig. 13. Field photographs showing key structural features within the Drbsdal eclogite body. Locations are numbered and shown on Fig. 5. (A) Lens of unretrogressed eclogite enclosed within large foliated felsic dyke; notebook is 20 cm in length (location 1). Map view; looking east. (B) Protrusion at margin of felsic dyke; white markers indicate contact and black marker indicates foliation in surrounding unretrogressed eclogite. Note that the foliation runs broadly from the left to the right of the photograph, and is clearly cut by the protrusion, seen in the centre of the lower part of the photograph (location 1). Map view; looking northeast. (C) Folded kyanite-bearing vein in mylonitic eclogite (type I vein as described in text; foliation of eclogite runs from the left to the right of the photograph); black markers indicate opposite ends of an individual kyanite lath, which is oriented at a high angle to foliation. Length of pencil, 15 cm (location 2). Profile view; looking north. (D) Cross-cutting kyanite-bearing vein (type II vein) in folded mylonitic eclogite; individual laths of kyanite up to ~10 cm are visible in this view. Note that laths are oriented subparallel to foliation. Length of hammer, ca. 80 cm (location 3). Profile view; looking northwest. (E) Close-up view of kyanitebearing vein in boudin neck (type III vein). Note clustering of kyanite laths towards vein neck (location 4). Map view; looking north. (F) Amphibole–quartz vein in mylonitic eclogite. Note highly irregular vein margins with tapering protrusions into the wallrock. Pencil (15 cm in length) indicates position of ca. 2-cm amphibole crystal at vein margin (location 5). Looking north; outcrop dips gently towards viewer. (G) Network of E–W-trending plagioclase–quartz–amphibole veins near margin of eclogite body. Clinometer, 10 cm long (location 6). Looking north; map view. (H) Submeter-scale ductile amphibolite-facies shear zone within eclogite in the Teiges3ta area. Shear sense is sinistral. Pencil (15 cm in length) indicates orientation of eclogite-facies foliation outside the zone of influence of the shear zone (location 7). Map view; looking north. (I) Meter-scale brittle amphibolite-facies shear zone (vein with small shear displacement). Markers indicate limits of related retrogression in eclogite. Length of clinometer is 10 cm (location 8). Map view; looking east.

associated with the occurrence of large quantities of smaller felsitic bodies, occurring as net veins or disordered arrays of dykelets. This suggests that much

of the felsitic material may have been derived locally, by partial melting of the eclogite itself, or sourced from the surrounding granodioritic gneisses.

18

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

4.6. Veins A highly characteristic feature of the Drbsdal body is the spectacular development of abundant eclogitefacies veins with a variety of morphologies and compositions. Based on their mineral assemblages, these veins are divided into four main types. The types are ordered according to relative abundance: (1) kyanite–quartz–phengite, (2) quartz, (3) amphibole– quartz, and (4) feldspar–quartz–amphibole. 4.6.1. Kyanite–quartz–phengite veins Veins dominated by kyanite are present throughout the mafic body, and field observations indicate that they contain the mineral assemblage kyanite+quartz+white micaFomphaciteFgarnetFamphiboleFrutile. Individual kyanite laths may be up to 15 cm in length, and laths of 5–10 cm length are common. No alteration of the eclogite wall rock is observed at the margins of kyanite-dominated veins; therefore, we conclude that the veins formed at eclogite facies. Kyanite veins are divided into three types (I–III), on the basis of structural setting with respect to the surrounding eclogite. Type I veins are oriented subparallel to the compositional foliation, and can often be traced around eclogite-facies folds. Most type I veins contain a strong lineation, defined by the shape-preferred orientation of kyanite laths, which plunge 15–408 to the west or southwest (Figs. 13C and 14A). Where kyanite veins are folded, the kyanite laths may be concentrated into bundles at the hinge region, and are generally oriented subparallel to the fold hinge. Kyanite lineations in the dfold limbT sections of such veins may also be aligned along the fold limbs, at a high angle to the fold hinge and lineation in the eclogite. Type I veins are spectacular, varying from 5 cm to N3 m in total traceable length, and vary in thickness from 1 cm to 1 m. Type II veins cut the eclogite-facies foliation at various angles. They are generally relatively straightsided and taper towards each end (Figs. 13D and 14B). Lineations within these veins are again defined by a grain shape fabric of kyanite, and plunge 15–408 to the WSW. Although the veins are often oriented at a high angle to the eclogite-facies foliation, there may be a continuity of the eclogite-facies foliation through the kyanite laths from one side of the vein to the other.

Fig. 14. Schematic sketches showing characteristic relationships of kyanite-bearing veins with surrounding eclogite. (A) dType IT vein; characteristically subparallel to eclogite-facies foliation and can be traced around folds. (B) dType IIT vein; clearly cross-cutting eclogite-facies foliation and with kyanite laths oriented subparallel to compositional foliation in the surrounding rock. (C) dType IIIT vein; characteristically located in the neck regions of boudins, with kyanite laths subparallel to the local foliation.

These veins range from 2 to 50 cm in width, and from 10 cm to N4 m in length. Type III kyanite veins are developed in the necks of eclogite-facies boudins, with kyanite fibres up to 10 cm in length mimicking the lineation and foliation in the boudinaged eclogite itself, forming a fan-shaped aggregate on each side of the boudin neck. These aggregates are 2–15 cm in width and 5–30 cm in total length, and lineations within them have plunges of 20–408 to WNW, W, or WSW (Figs. 13E and 14C). 4.6.2. Quartz veins Quartz veins of 10 cm to 1 m in length and 5 mm to 40 cm in width are common throughout the mafic body. Occasional fine-grained (b1 mm) zoisite,

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

omphacite, and garnet grains are included in the vein fill, but are only observable in thin section. Vein margins are often enriched in white mica and sometimes in rutile. Most of the quartz veins are steeply dipping and oriented E–W to ENE–WSW, thus lying subparallel to the pervasive compositional foliation within the body, but occasionally they cut the foliation at a high angle. Some contain a weak lineation defined by grain shape fabric of quartz, plunging moderately or steeply to the NW or W. Veins that are oriented subparallel to the compositional foliation are commonly deformed within eclogite-facies structures such as folds or boudins (Sections 4.1 and 4.4). Alteration of the eclogite at the margins of quartz veins is rare, but occasionally there is a change in colour of the wallrock from green to grey or black towards the vein. Eclogite-facies mineralogy and structural relationships (i.e., folding) of the majority of the veins clearly indicate that they originated under eclogite facies. 4.6.3. Amphibole–quartz veins Amphibole-bearing veins are less common than kyanite-dominated veins, and are composed of barroisitic amphibole+quartzFgarnetFwhite micaF kyanite. These veins vary from 5 cm toN1 m in length and from 1 to 30 cm in width. In most cases, amphibole-bearing veins crosscut the eclogite-facies foliation, dipping steeply to the west or northwest. Lineations within these veins are defined by amphibole grain shape fabric, and plunge 15–408 to the WNW, W, or WSW. Vein margins are often highly irregular, with tapering protrusions into the wallrock (Fig. 13F), but field observations suggest that no alteration of the eclogite wallrock occurs at the vein margins. 4.6.4. Feldspar–quartz–amphibole veins Feldspar-bearing veins are divided into two groups, based on their orientation and mineral content. N–Strending alkali feldspar-bearing veins are observed throughout the eclogite body, and constitute one endmember of a continuum of structures from brittle veins with no apparent offset to mineralised ductile shear zones. These are described in more detail in Section 4.7. Plagioclase-bearing E–W-trending veins are common at the margins of the mafic body, and generally have moderate to steep dips. They run subparallel to

19

the lithological contact between the eclogite and the granodioritic gneiss, and are associated with retrogression of the eclogite to a fine-grained grey to black amphibolite. They vary in thickness from b1 to N20 cm, and commonly form a pervasive network of anastomosing veins, giving the rock a stripy appearance (Fig. 13G). 4.7. Shear zones Measurable shear zones at the outcrop scale are rare in the Drbsdal mafic body as a whole. However, a range of shear zone styles is observed (Fig. 13H), ranging discontinuously from brittle to ductile. In the brittle-style shear zones, the shear plane itself consists of a vein filled with an amphibolite-facies mineral assemblage of feldspars (these vary within and between different veins, and can be plagioclase or alkali feldspar), quartz, and, in some places, amphibole (Fig. 13I). The eclogite-facies compositional foliation is transected at a high angle and may be offset by distances of up to 50 cm. All types of kyanite veins are also transected by these brittle structures. This type of shear zone is generally accompanied by a diffuse zone of alteration, 10 cm to 1 m in width, in which the eclogite is replaced by a fine-grained grey rock towards the shear zone (Fig. 15A). Brittle shear zones occur in the interior of the mafic body, and are less common at its margins. Ductile shear zones are characterized by intensification and reorientation of the compositional foliation, and are also often accompanied by a change in colour and grain size of the eclogite. This deformation is most intense in the shear plane (Fig. 15B). Ductile shear zones occur at or near the margins of the mafic body, particularly where the eclogite has a relatively narrow outcrop width, or tapers laterally; they are not common within the interior of the eclogite body. For both ductile and brittle types of shear zones, where a displacement is observed, the dextral shear zones typically dip steeply or moderately to the SW, whereas sinistral shear zones generally dip steeply to the SE or SW. It is likely that these orientations represent conjugate sets of structures (Fig. 16) and that shear senses are as they appear (i.e., a mixture of top-towest and top-to-east). However, it is possible that the real shear sense of some of these shear zones is different from the apparent shear sense, since shear

20

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Fig. 15. Characteristics of amphibolite-facies shear zone type features: (A and B) Cartoons showing end-members of a continuum of amphibolite-facies shear zone features found within the Drbsdal eclogite (A, ductile end-member; B, brittle end-member).

zones are commonly viewed in section, and exposure of shear planes or lineations is rare.

5. Structural and petrographic interpretation In this section, each type of structure will be discussed and interpreted, and then a large-scale model for the structural evolution of the Drbsdal eclogite body will be presented and explained.

Fig. 16. Orientation measurements of shear zones plotted as poles to shear planes, equal area projection, and upper hemisphere.

5.1. Eclogite-facies folding The Drbsdal mafic body is a coherent sheet of compositionally banded eclogite that is tightly folded on the kilometer scale. Folds at the kilometer scale are not seen clearly in map view; however, the vergence of outcrop-scale folds indicates that the body is folded on a larger scale. Investigation of the symmetry of meter- and centimeter-scale folds via the construction of cross-sections provides further evidence that kilometer-scale folds have affected the whole of the eastern part of the body. The larger-scale folds are interpreted to have morphologies resembling those of the mesoscale parasitic folds, which are generally asymmetric, tight to isoclinal structures. The apparent lack of large-scale fold axes in map view is a result of the steepness and bulk asymmetry of structures in the area; fold axes are not necessarily associated with obvious switches in dip direction of the foliation. The mesoscale folding affects layers with well-preserved eclogite-facies mineralogy, indicating that folding was active during metamorphism at eclogite facies. There is a possibility that folding began before the onset of eclogite-facies conditions, continuing into the eclogite-facies phase of metamorphism, but the lack of preserved protolith to the eclogite precludes further investigation. At the map scale, the pinching out of the mafic body at its extreme west and east, and the spatial relationships between the Drbsdal eclogite and other

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

mafic bodies in its immediate vicinity, indicate that boudinage of mafic layers and lenses has contributed to the present distribution of these mafic bodies, and has also modified their shapes (see Figs. 1 and 2). Zones of retrogressive amphibolite-facies material are strongly associated with local shearing of the Drbsdal body. These shear zones are concentrated in the Teiges3ta area and in the termination of the Drbsdal body to the NW of Trausdalsvatnet. This indicates that eclogite-to-amphibolite facies E–W-directed stretching and boudinage were partially accommodated by movement on submeter-scale shear zones, and that this deformation further altered the original shape of the tightly folded eclogite layer to give the observed lensoid shape. Additional evidence of this boudinage is found further to the west, in the eclogite around Svanetjorna. Folding is relatively well constrained for the eastern end of the body. Notably, there is a finger of granodioritic amphibolite-facies rock protruding into the mafic body to the south of Butjorna, which is interpreted as the core of an antiform (Figs. 5 and 8). This feature appears again as a dwindowT to the west, where a steep westward-facing hillside has cut through the eclogite exposing the granodioritic core of the fold (Fig. 5). This antiform is also seen as the westward protrusion of eclogite to the north of Trausdalsvatnet (Fig. 5). The curve in the eclogite boundary slightly further north, at Svanetjorna, is due to the presence of a westward-plunging synform. The elongate, bifurcated shape of the body is therefore the result of folding on the kilometer scale, modified by boudinage (see Section 5.7; Figs. 17 and 18). The Drbsdal body is currently situated in the southerly dipping limb of a large-scale antiformal culmination between the Kvamshesten and Solund Devonian basins (Fig. 1). Folds in this system typically have wavelengths in the 10–15 km range. The relationship of the eclogite-facies folding to this regional-scale structure is discussed in more detail below. 5.2. Eclogite-facies type I lineation The parallelism of eclogite-facies fold hinges with lineations defined by eclogite-facies minerals indicates a link between these structure types. This relationship between lineations and fold hinges is

21

found in structural domains dominated by a N–S linear trend and also in those dominated by and E–W linear trend (Figs. 5 and 6). Grain shape lineations defined by omphacite, zoisite, and kyanite are interpreted as stretching lineations for a number of reasons. Firstly, the opening direction of fractures and healed fractures within zoisite grains always coincides with lineation direction. Textural evidence suggests that the zoisite and omphacite are in equilibrium at the thin section scale and therefore represent the same phase of eclogite-facies deformation. Secondly, this lineation coincides with the orientation of kyanite laths and amphibole fibres precipitated as vein fillings during eclogite-facies metamorphism. These grains can often be traced through veins with opening directions that cut foliation and lineation of the surrounding eclogite at a high angle. These veins therefore opened with an E–W extension direction (in the present frame of reference) during eclogite-facies metamorphism. Thirdly, hinge parallel boudinage of the eclogite and associated development of kyanite veins in boudin necks are further evidence for large amounts of eclogite facies E–W stretching in the eastern end of the Drbsdal body, since no change in metamorphic conditions appears to accompany the boudinage. A mechanism for developing fold axes parallel to a linear stretching feature must therefore be inferred. Folds with hinges parallel to a lineation may develop in the following ways: (1) an array of folds with either random or common orientation may be subjected to a strain large enough to rotate the folds towards parallelism with a subsequent stretching direction (Escher and Watterson, 1974; Sanderson, 1972; Skjernaa, 1980); (2) in situations where the intermediate strain orientation Y is constrained to be perpendicular to the original layering, and deformation takes place under plane strain conditions, buckle folds may form with axes parallel to the regional or local stretching direction and therefore have similar orientations to simultaneously formed linear stretching features (Grujic and Mancktelow, 1995; Watkinson, 1975); (3) a rock mass containing a preexisting linear fabric may be folded; for hinge lines to form parallel to the fabric, or dbending anisotropy,T the fabric must be sufficiently well defined to exert a mechanical control on the orientation of new fold axes (Cobbold and Watkinson, 1981); (4) folds may initiate

22

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Fig. 17. Schematic 3D diagram showing stage 3 of the structural model for formation and deformation of the Drbsdal eclogite body (see Section 5.7 of text). c=shear strain.

with hinges parallel, oblique or perpendicular to the transport direction, as a result of differential movement on a shear plane, and subsequently oblique or perpendicular hinges experience passive rotation towards the transport direction with an increase in strain intensity (Coward and Potts, 1983); (5) stretching lineations and fold hinges may form simultaneously with a common orientation, specifically within a constrictional strain field (Krabbendam and Dewey, 1998). The high degree of compositional heterogeneity between layers and the widespread presence of class 3 folds are indicative of high competence contrasts. The morphologies of linear and planar eclogite-facies structures suggest that high strains were achieved and that ductile behaviour dominated the folding and boudinage. Although omphacite–zoisite-rich layers of the rock contain a strong lineation that may possibly have formed before nucleation of the earliest folds, the

bending anisotropy due to this lineation is unlikely to have been marked enough to affect the orientations of new folds (Cobbold and Watkinson, 1981). The production of folds in this way is only likely in situations where the linear feature causing the bending anisotropy is composed of deformation-resistant rods or fibres that are embedded in a much dweakerT matrix and do not deform. It is likely that the large variations in mineralogy observed between layers led to significant competence contrasts during deformation, and that layer competence therefore had a much larger control on the mechanism and morphology of the folding than the linear features themselves. We therefore conclude that mechanism (3) (Cobbold and Watkinson, 1981) is unlikely to have caused the parallelism of lineations and fold hinges in the Drbsdal eclogite. If the fold hinges developed parallel to the extensional lineation according to the model of Coward and

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Potts (1983), the fold hinges would not necessarily have the same orientations as the lineations. The model dictates that folds may form with hinges oblique, perpendicular or parallel to the transport direction, and that oblique or perpendicular folds may subsequently rotate into parallelism with the transport direction as the shear zone advances. If fold hinges in the Drbsdal eclogite had been reorientated postfolding but synchronously with the lineation according to this model, we would expect to see a statistical deviation around the modal orientation for fold hinges, but a smaller deviation or no deviation from the modal orientation for lineations. Stereonets show tight clusters for orientations of both fold hinges and lineations for the eastern portion of the Drbsdal body, and for the western portion of the body, we see a large range in orientations of lineations and a comparatively smaller range in orientations of fold hinges. The model also requires large shear strains (c=20 or more) to rotate folds initiated at high angles to the transport direction into subparallelism with the shear direction (Alsop and Holdsworth, 1999; Skjernaa, 1980). Given the large amounts of deformation intrinsic to any tectonic model for exhumation, it is possible that the Drbsdal body has undergone shear strain of this order of magnitude. However, the large spread in lineation orientations is not easy to reconcile with this model. It is also unlikely that the fold hinges and lineations formed coevally and parallel to a regional or local stretching direction, under plane strain conditions with Y constrained to be perpendicular to the original layering as described by Grujic and Mancktelow (1995) and Watkinson (1975) (mechanism (1)). Folds produced in such a manner are observed to have low amplitude–wavelength ratios, even at high strains. This is not consistent with our observations. The strength of the L-fabric and its coincidence with stretching direction of fold hinges and the long axes of boudins suggest that formation of folds and lineations was simultaneous. The mechanism involving a constrictional strain field (Krabbendam and Dewey, 1998) may be responsible for the parallelism of fold hinges, lineations, and long axes of boudins. Populations of fold axes initiated under purely constrictional strain conditions would perhaps be expected to have randomly oriented axial planes. However, it is likely that the orientation of layering (D1) present in the Drbsdal rocks before folding began

23

would also have a large effect on the formation of the fold population. Consider the orientations of the three principle strain axes for the folding event (D2): X is oriented parallel to present fold hinges (broadly E–W on map), Y is oriented perpendicular to X but within the axial surface (in and out of map plane), and Z is perpendicular to fold hinges and axial surfaces (N–S on map). If X/YJY/Z, and YNZ, then folds forming in a dhorizontalT multilayer (i.e., layer subparallel to X–Z plane) with large competence contracts between layers would develop with hinges parallel to the X direction and amplify in the Y direction whilst shortening in the Z direction. Fold axial planes produced in these conditions would be consistently steep to upright in orientation, as shown for the eclogite-facies folds at Drbsdal. It is also necessary to consider that an existing population of upright folds could be progressively rotated into parallelism under constrictional conditions, and thus also be parallel to simultaneously formed stretching features such as lineations and long axes of boudins. The preferred interpretation for the initial stage of fold formation is the constrictional model of Krabbendam and Dewey (1998), which is compatible with our data for cases where X/YJY/Z. Perturbations to this strain field, perhaps due to evolving rheological contrasts at the margins of the body, or volume changes due to metamorphic reactions taking place after initial formation of folds and stretching lineations could explain the obliquity of linear structures in the Teiges3ta and Tinghaugen areas. However, a more likely explanation of the observed obliquities is the onset of top-west–directed overshear (Fig. 17). Evidence of top-west shear is seen throughout the WGC in amphibolite- and greenschist-facies rocks (Andersen and Jamtveit, 1990; Fossen, 1992; Eide et al., 1999; Hacker et al., 2003). We propose that the onset of a precursor to this amphibolite- and greenschist-facies shear began at eclogite facies. This top-west-directed overshear (Fig. 17) tightened folds in area C, and rotated fold hinges in areas A and B clockwise in the present-day map view (area C is the area above plane B in Fig. 17, and areas A and B are below plane B in Fig. 17; these also correspond with areas A, B, and C in Fig. 6). The fold rotation models of Sanderson (1972), Escher and Watterson (1974), and Skjernaa (1980) may also apply, but are considered to be less important than the constrictional

24

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

transtension model of Krabbendam and Dewey (1998) in the present case. In summary, we propose that a strong preexisting planar foliation was folded to produce a population of folds with similarly oriented axial planes. These folds were subsequently subjected to sufficient constrictional strain to ensure parallelism of their axes with each other and with lineation, and eventually result in hinge-parallel boudinage. The eclogite-facies folds, stretching lineations, and boudins were formed at different but perhaps overlapping stages of one progressive event. It follows that the folds trending N–S in area A (Fig. 6) record an early phase of eclogite-facies deformation. Subsequent deformation was partitioned, and only involved those parts of the Drbsdal body outside area A. The early folds have subsequently been tightened, particularly in the eastern end of the body and its margins. As a result of this partitioning, the majority of the eclogite body and its surroundings has undergone a clockwise rotation relative to the dTeiges3ta block.T In map view, the structures in area A are oriented N–S to NE–SW, and the structures in the rest of the body are oriented broadly E–W. The E–W- to ENE–WSW-trending eclogite-facies structures within the rest of the Drbsdal body are oriented subparallel to the amphibolite-facies structures in the surrounding granodioritic rocks. This indicates that although ductile deformation of the Drbsdal body must have ceased by the onset of amphibolite-facies conditions, no major rotation of the Drbsdal body relative to its surroundings occurred after this time. Although structures within the whole of the WGC probably rotated due to transtensional deformation during exhumation, the fact that structures within the body are subparallel to those outside it indicates that rotation of the whole Drbsdal body relative to its surroundings was limited. This lack of rotation may be a result of the large size and elongate shape of the Drbsdal body, features that are undoubtedly little changed since eclogite facies. Smaller eclogite bodies would perhaps have experienced more substantial passive rotation since eclogite facies. 5.3. Eclogite-facies type II lineation (surface lineation) The confinement of the dtype IIT or surface lineation to foliation planes may indicate that it

formed by slip on the foliation surfaces. Where present, the surface lineation is oriented at a high angle to the pervasive, penetrative lineation, and this strongly suggests that the two types of structure formed at a different times. If the surface lineation formed as suggested, via slip on foliation planes, it cannot be an earlier feature than the penetrative lineation, as formation of the pervasive, penetrative lineation should have overprinted most small-scale linear features, especially those defined by hydrous phases. 5.4. Relationship of L at Drbsdal to regional linear dexhumationT features The orientations of fold axes, boudin long axes, and lineations in the eastern end of the Drbsdal body correspond closely to those of regional linear features formed during extension, such as the crenulation cleavage and amphibolite- to greenschist-facies shear sense indicators associated with movement along the NSDZ (Andersen et al., 1994; Krabbendam, 1998; Krabbendam and Dewey, 1998), and the eclogite- to amphibolite-facies stretching lineations described by Engvik et al. (2000). They are also parallel to the extensional fabric observed in the nappe units (Fossen, 1993). Although structures within the Drbsdal body and those further north within the WGC and overlying mid- to upper-crustal units and detachment zones were formed in different spatial settings, they may all have formed within a short time frame and corresponding to a stage of relatively fast burial followed by rapid extensional exhumation. Recent U–Pb ages from zircons and 40Ar/39Ar closure data from WGC and Hyllestad Complex localities ~3 km south of Drbsdal indicate that a mere N5–10 Ma may have passed between maximum burial of the WGC (ca. 410–400 Ma) and exhumation to upper crustal levels (ca. 403 Ma) (Chauvet and Dallmeyer, 1992; Hacker et al., 2003). If this is the case, then the Drbsdal structures could have formed either in several separate regional events with short total duration or during one short-lived phase of progressive deformation. However, evidence that links the formation of eclogite-facies stretching lineations in the lower crust to the amphibolite- and greenschist-facies extensional structures of the mid- and upper crust is largely

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

circumstantial. The presence of several structurally different domains within the Drbsdal body itself implies that deformation within the lower crust was gradually partitioned during passage through eclogitefacies P–T space, and was influenced by different factors such as local composition of the eclogite, thickness of the mafic body, and distance from the margins of the body. The domains containing E–Wtrending structures therefore represent zones of reworking through which the early-formed folds have been tightened, stretched, and rotated by subsequent deformation, perhaps during early exhumation. Rey et al. (1997) postulated extension with top-west-directed shear operating at eclogite facies as well as lower grades, and our data support this possibility. The parallelism of structures formed in vastly different parts of P–T space may be explained if no major changes in deformation regime occurred during the exhumation from eclogite-facies to amphibolitefacies conditions. This parallelism of structures formed over a range of metamorphic facies is not typical of the WGC rocks. For instance, Krabbendam et al. (2000) show that granulite-facies and eclogite-facies lineations are almost perpendicular in the Nordfjord area, and Engvik and Andersen (2000) show that dearlyformedT eclogite-facies lineations are consistently at high angles to dlater-formedT eclogite-facies lineations. Where lineations are consistently oblique, this indicates that an earlier part of the eclogite-facies history is seen locally, and where lineations within small granulite or eclogite bodies have variable lineations, this is likely to indicate subsequent rotation of these small bodies relative to their surroundings. The Drbsdal body is large, and the majority of its smallscale structures are subparallel to those in the surrounding rocks, as are its boundaries. The oblique fabrics at Tiegesata and Tinghaugen represent parts of the body that ceased to deform at an earlier stage. Partitioning of strain into the rest of the eclogite body resulted in a clockwise rotation of structures in area C relative to areas A and B. This obliquity of fabrics does not necessarily reflect a different strain regime at the crustal scale. Moreover, although the whole of the Drbsdal body could have rotated since it was present at eclogite-facies depth, this is unlikely due to the volume of rock involved and its elongate nature. Indeed, the idea of a constant strain regime during the exhumation from eclogite-facies conditions to amphibolite-facies

25

conditions fits well with the short-lived exhumation event proposed by Hacker et al. (2003) for the WGC. A sequential partitioning of deformation into domains of different rheologies would explain the preservation of the deepest-formed structures in the largest mafic bodies and the preservation of shallowly formed structures in granodioritic or lithologically variable units such as the felsic gneisses or finely interleaved mafic and felsic rocks. There is a conspicuous lack of systematically oriented outcrop-scale eclogite-facies shear sense indicators preserved in the Drbsdal mafic body. This indicates that coaxial flow may have dominated the deformation regime(s) active within the Drbsdal body during the latter stages of residence at eclogite facies, and possibly the whole of the eclogite-facies history. This interpretation agrees with those of Andersen and Jamtveit (1990), Andersen et al. (1994), and Jolivet et al. (1994), who argued for coaxial deformation of the lower crust during early extensional collapse. The model of Krabbendam and Dewey (1998) involves bulk noncoaxial deformation of the lower crust, with exhumation controlled largely by sinistral transtension. Lattice-preferred orientation (LPO) of omphacite grains in gneissic eclogite samples from the Drbsdal body (Foreman and Wheeler, in prep.) shows no asymmetry—an observation that is consistent with models involving coaxial deformation of the lower crust. However, it is also possible that deformation of the WGC as a whole was noncoaxial, but that deformation was inhomogeneous, and coaxial deformation occurred locally. It follows that the Drbsdal body may have rotated anticlockwise in a large-scale shear affecting the WGC, while deforming internally in a coaxial manner. However, the subparallelism of the latest eclogite-facies structures with amphibolitefacies structures outside the body indicates that posteclogite-facies rotation of the body relative to its immediate surroundings was limited (see Section 5.2). The evidence presented in the present contribution shows that a number of stages of pervasive ductile deformation affected the Drbsdal mafic body during residence at eclogite-facies conditions. It is therefore likely that the E–W-oriented eclogite-facies fabrics preserved at Drbsdal were developed during the early stages of exhumation, since there is no evidence for a major change in orientation of the strain field between eclogite facies and amphibolite facies.

26

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

5.5. Amphibolite-facies deformation features The presence of mafic amphibolite-facies mineral assemblages as vein fillings in brittle shear zones, and the replacement of eclogite-facies minerals by amphibolite-facies material in both brittle and ductile shear zones indicate that the deformation they represent occurred under amphibolite faces conditions, during exhumation. There are two possible interpretations to explain the relative timing of ductile and brittle shear zones. The first is that the ductile shear zones predate the brittle ones and therefore formed earlier in the exhumation history. If this is the case, brittle deformation processes would have become dominant as exhumation progressed and veins would form with progressively more brittle features. The alternative is contemporaneous formation of brittle and ductile shear zones, but in different parts of the mafic body. The presence of a fine-grained, weakly foliated, amphibolite-facies carapace around the mafic body indicates that metamorphic reactions and deformation were possible in mafic material at the margins of the body during the amphibolite-facies stage, but that no deformation and very little retrogression occurred at this stage within the interior of the body. It is therefore likely that amphibolite-facies shear zones formed in the felsic material outside the mafic body during exhumation and in the mafic material at the very margins of the body, where deformation was enhanced by reaction softening and possibly also (in the mafic margins) by an influx of fluid from the surrounding felsic rocks, and brittle veins formed simultaneously in the interior of the body where ductile deformation of the eclogite had already ceased. Top-west shear fabrics are observed in the granodioritic amphibolite-facies rocks less than 5 m from the margin, indicating that rotational deformation began to occur during the amphibolite-facies exhumation stage. The presence of numerous amphibolite-facies shear zones at the two western terminations of the body and the coincidence of the long axes of eclogite-facies mafic bodies in the area with both eclogite-facies stretching features and regional amphibolite to greenschist-facies stretching features indicate that the regional-scale boudinage of mafic material occurred before or during the early amphib-

olite-facies stage of exhumation. Following this boudinage event, deformation was largely accommodated in the surrounding granodioritic lithologies. During the amphibolite-facies stage of exhumation, the mafic bodies behaved as rigid rafts of material within a more deformable granodioritic matrix, and therefore retained much of their eclogite-facies mineralogy and structure. 5.6. Relationship of late regional-scale folding to structures at DrØsdal Regional-scale folds with E–W-trending hinges dominate the present outcrop pattern of the WGC. Modelling based on Ar spectra from K-feldspar (Eide et al., 1999) and muscovite (Andersen et al., 1998) imply that although decompression-related cooling of UHP and HP rocks from ~700 to ~350 8C had already occurred by 390F10 Ma (Andersen et al., 1998), there was a marked increase in cooling rate during Late Devonian–Early Carboniferous time (360–340 Ma; Eide et al., 1999). This phase of rapid cooling was in turn followed by slower cooling during the Permian and Late Jurassic–Early Cretaceous (300–140 Ma). The Late Devonian–Early Carboniferous cooling event correlates well with the final stages of N–S shortening affecting the entire existing crustal sequence (Eide et al., 1999) (see also Torsvik et al., 1986; Krabbendam and Dewey, 1998; Osmundsen and Andersen, 2001). It is important to note, however, that the Middle Devonian rocks were deposited in separate topographically constrained basins, and not as a continuous sheet that has been subsequently folded (Andersen et al., 1998), so some N–S shortening must have occurred prior to their deposition. As part of the WGC, the Drbsdal mafic body and its surrounding gneisses were undoubtedly reoriented as a result of the dlateT E–W folding discussed immediately above. However, there is no evidence to suggest that structures within either the granodioritic rocks of the WGC or the mafic rocks they envelope were in any way reworked during this event (see also Section 5.7 and Fig. 18). The Drbsdal mafic body is currently situated in the southerly dipping limb of a regional-scale antiform, close to the NSDZ. It records predominantly WSW-plunging eclogitefacies lineations, and lineations preserved within the

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

27

Fig. 18. Schematic 3D diagram showing geometry of the Drbsdal eclogite body after stage 4 of the structural model (see Section 5.7 of text).

surrounding gneisses are similarly oriented (Fig. 5). According to its position on the map, the V3rdalsneset eclogite body is situated in the northerly dipping limb of the same antiform, and is also close to the NSDZ (Figs. 1 and 2). Engvik and Andersen (2000) report subhorizontal N–W-trending lineations for V3rdalsneset. Unfolding of the regional antiform between the Solund and Kvamshesten basins reveals that the eclogite-facies lineations preserved within the Drbsdal and V3rdalsneset bodies match closely, provided that the lineations originally plunged shallowly to the SW or WSW. This not only implies that the two bodies record simultaneous eclogite-facies events, but that minimal rotation of the mafic bodies as individual dpodsT has occurred since formation of the eclogite-facies structures. Although circumstantial, this observation adds weight to the suggestion that the Drbsdal eclogite body records a sequence of structures formed at a range of times within its eclogite-facies history. In particular, structures formed during early residence at eclogite facies are partially overprinted or reoriented by structures developed during passage through the eclogite-facies portion of the exhumation path.

5.7. The model 5.7.1. Stage 1 Formation of an eclogite sheet, probably southerly dipping, occurs. The precursor to this eclogite sheet was probably an igneous sheet of basic to intermediate composition that had already passed through amphibolite and granulite facies, but may or may not have chemically equilibrated at these metamorphic facies (see also Section 2). The eclogite sheet has developed a foliation defined by compositional differences, and also has both shape fabric and lineation. Although this initial stage of sheet formation and/or deformation at eclogite facies is compatible with a constrictional strain regime, direct evidence for transtension this early in the strain history is lacking. 5.7.2. Stage 2 The eclogite sheet becomes folded on all scales ranging from b1 m to N1 km, with smaller folds parasitic to larger ones. The folds form with variable hinge orientations, generally plunging W to SW, and with axial surfaces generally dipping gently to the south. Lineations form coevally, broadly parallel to

28

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

fold hinges, implying a constrictional strain field with the maximum stretching direction oriented roughly east–west, and shortening in the north–south and vertical directions, in the present reference frame. These structures may have formed under transtensional strain conditions, as described by Krabbendam and Dewey (1998) with reference to amphibolite- and greenschist-facies structures. It is interesting to note that eclogite-facies constrictional fabrics related to transtension have not been observed elsewhere in the WGC, which implies either (1) that the Drbsdal area has a different local strain history to other parts of the WGC, or that (2) the other eclogite bodies in the WGC ceased to deform internally at an earlier stage than the Drbsdal body. 5.7.3. Stage 3 Boudinage of the eclogite sheet occurred under continuing constrictional strain, presumably in the same transtensional regime. Long axes of the kilometer-scale boudins were broadly E–W-trending (Fig. 17). Orientation of boudin necks is difficult to ascertain, but they are arbitrarily shown perpendicular to lineation (i.e., stretching direction) in Fig. 17. The onset of top-west-directed overshear then caused rotation of the W-plunging antiform–synform pair, which made up the upper part of the Drbsdal sheet at this time (Fig. 17). The rotation would appear clockwise in present-day map view. This rotation tightened the existing folds throughout area C, and but caused rotation only at the edges of the eclogite body in areas A and B (area C is the area above plane B in Fig. 17, and areas A and B are below plane B in Fig. 17; these also correspond with areas A, B, and C in Fig. 6). This top-west-directed overshear probably occurred after boudinage but could have begun before the sheet was boudinaged, and is a precursor to the top-west shear evidenced throughout the WGC in amphibolite- and greenschist-facies rocks (Fossen, 1992; Andersen et al., 1998; Krabbendam and Dewey, 1998). Folds mainly predate the onset of top-west-directed shear. It should be noted that since folds were already quite tight with many hinges not far from their final orientation, as shown by fold style in area A (the NW part of the Drbsdal body), large additional shear strains would not have been necessary to reach the present-day structural geometries.

5.7.4. Stage 4 Reorientation of the Drbsdal eclogite body into its present-day position (Fig. 18) occurred by folding around the Solund synform. The Solund synform developed during regional N–S shortening, which affected the whole crustal sequence during Late Devonian–Early Carboniferous times, and which can also be related to transtension (Krabbendam and Dewey, 1998; Osmundsen and Andersen, 2001). 5.8. Regional implications Two main unresolved questions concerning the HP and UHP rocks of the WGC were put forward in the introduction to this contribution: (1) What was the mechanism of exhumation of the HP rocks? (2) What is the relationship between the HP and UHP rocks, and how did they become juxtaposed? Based on the compilation of thermochronological data, it can be demonstrated that the UHP rocks constitute the lowermost structural units in western Norway (Hacker et al., 2004). The UHP rocks cooled below Ar blocking in muscovite after 380 Ma, whereas higher structural levels in the WGC also containing eclogites cooled between 400 and 380 Ma. The large-scale EW folds also fold the muscovite Ar isochrons in the WGC (Hacker et al., 2004). The model presented above suggests a structural development of an HP eclogite body under conditions of constrictional transtension, with an additional component of top-west-directed overshear at a relatively late stage of the eclogite-facies history. Krabbendam and Dewey (1998) presented a transtensional model for the exhumation of the WGC, in which relatively homogeneous constriction produced linear fabrics in the deeper part of the orogen, including lineation-parallel folding. However, the model was based on observations of late-orogenic amphibolitefacies structures; indeed, the authors state that a gap of about 5–10 Ma exists in the structural record (Krabbendam and Dewey, 1998). This dgapT relates to the time interval between the formation of eclogitefacies structures (Andersen et al., 1991, 1994) and the late-orogenic amphibolite-facies structures on which the transtensional model is based (Krabbendam and Dewey, 1998). Since the linear (constrictional) structures preserved within the Drbsdal eclogite body are subparallel to the regional principle stretching direc-

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

tion of the transtensional strain field of Krabbendam and Dewey (1998), we propose that these exceptional eclogite-facies structures fill the dgapT in the structural record by providing a record of constrictional deformation in transtension at conditions of T=720– 830 8C and P=19–21 kbar. Additionally, we have shown that at least some of the eclogite bodies in the WGC continued to deform significantly after the onset of exhumation. Previously, data related to eclogite-facies deformation in western Norway were very limited, and could not be easily linked to the amphibolite- and greenschist-facies portions of the exhumation path. This contribution allows this link to be made at least in one area. The relationship between HP and UHP units cannot be directly elucidated from the data presented in this contribution. Even so, a few useful discussion points can still be raised regarding possible links between the HP rocks and UHP rocks. We have presented a model involving exhumation of an HP eclogite body under conditions of constrictional transtension; however, horizontal stretching of the orogenic pile (including a mantle wedge) via transtension would not give us the observed structural sequence. This stretching alone would not remove the mantle wedge from above the WGC; it would just become thinner. Perhaps extensional shearing at an early stage in the exhumation history facilitated the juxtaposition of UHP and HP units in a thickened crustal welt, in a fashion similar to that proposed for the UHP and HP rocks of the Piemonte zone by Reddy et al. (1999) and Wheeler et al. (2001), and the WGC by Andersen et al. (1991), Rey et al. (1997), and Terry et al. (2000a,b). Since the Drbsdal mafic body does not appear to preserve UHP assemblages or structures, it cannot provide answers to questions about the UHP history. Nevertheless, it provides a much-needed link between the eclogiteand amphibolite-facies portions of the exhumation path of the WGC, since the structures it preserves are broadly coaxial with the later amphibolite-facies structures as outlined above.

!

!

!

! 6. Conclusions ! The Drbsdal mafic body is a coherent sheet of compositionally banded eclogite. It is tightly folded on the scale of hundreds of meters; folds

29

are asymmetric with upright axial planes, and their hinges plunge moderately to the WSW. These folds formed under conditions of constrictional transtension at eclogite facies. Eclogite- to amphibolite-facies E–W-directed stretching followed by top-west-directed overshear after the onset of exhumation further modified the shape of the Drbsdal body and other mafic bodies in its immediate vicinity. This phase of deformation was partially accommodated by movement on submeter-scale shear zones, giving rise to the present boudin-like distribution and elongate shapes of eclogite lenses in the area. Long axes of the lenses, and the orientations of lineations within them are consistent with stretching features in the surrounding amphibolite-facies gneisses, indicating continuity of deformation regime during the transition between eclogite- and amphibolite-facies exhumation. Pervasive eclogite-facies stretching lineations within the Drbsdal eclogite lie subparallel to eclogite-facies fold hinges, and to eclogite-facies lineations within veins (defined by kyanite and omphacite). Structural and metamorphic data indicate that the timing of formation of the pervasive lineation, the eclogite-facies folds, and the kyanitebearing veins overlapped substantially during the eventful eclogite-facies history dominated by a constrictional strain field, with X/YNY/Z and YNZ, and Yb1. The X-axis of the strain ellipsoid was oriented broadly E–W, the Z-axis was oriented N– S, and the Y-axis was subvertical in the present reference frame. Eclogite-facies fold hinges and lineations in the Tinghaugen and Teiges3ta areas are oblique to the present E–W-dominated structures observed throughout the rest of the Drbsdal body. These oblique fabrics represent parts of the body that ceased to deform at an earlier stage. Partitioning of strain into the rest of the eclogite body resulted in a relative clockwise rotation of the majority of the eclogite body. This observed obliquity was achieved by syneclogite-facies top-west-directed overshear. Fold axes, boudin long axes, and lineations in the majority of the Drbsdal body are subparallel to the regional linear features formed during extension. These include the crenulation cleavage and shear sense indicators associated with movement along the NSDZ (Andersen et al., 1994; Krabbendam,

30

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

1998; Krabbendam and Dewey, 1998), the eclogite- to amphibolite-facies stretching lineations described by (Engvik and Andersen, 2000), and the extensional fabric observed in the nappe units (Fossen, 1992, 1993). Circumstantial evidence from Drbsdal suggests that the subparallel eclogite-, amphibolite-, and greenschist-facies structures of the Sunnfjord area formed during one progressive deformation event corresponding to extensional exhumation. Our observations fit well with a model involving N–S shortening coupled with E–W (broadly horizontal) extension in the lower crust during early extensional exhumation. Local coaxiality of deformation during eclogite facies is demonstrated by coaxial structures at Drbsdal, but noncoaxial transtensional deformation probably characterized the large-scale behaviour of the WGC. Sequential partitioning of deformation into domains of different rheologies during exhumation may explain the preservation of varied mineralogies and structures in this part of the WGC, despite the shared P–T history. Broad antiforms have obviously rotated the rocks within the WGC. However, the close match in orientation of lineations from the Drbsdal and V3rdalsneset eclogites implies that minimal rotation of mafic bodies relative to each other has occurred since formation of the eclogite-facies structures.

Acknowledgements This work has benefited considerably from the constructive reviews of P. Rey and an anonymous reviewer. G. Potts and A. McCaig are also thanked for constructive comments, which improved the manuscript. H. Austrheim, M. Erambert, and M.G. Lund are thanked for their help with chemical analyses, and numerous useful discussions. NERC funding for R.F. is acknowledged.

References Abalos, B., 1997. Omphacite fabric variation in the Cabo Ortegal eclogite (NW Spain): relationships with strain symmetry during high-pressure deformation. Journal of Structural Geology 19 (5), 621 – 637.

Alsop, G., Holdsworth, R., 1999. Vergence and facing patterns in large-scale sheath folds. Journal of Structural Geology 21, 1335 – 1349. Andersen, T.B., Jamtveit, B., 1990. Uplift of deep crust during orogenic extensional collapse: a model based on field studies in the Sogn–Sunnfjord region of western Norway. Tectonics 9 (5), 1097 – 1111. Andersen, T., Jamtveit, B., Dewey, J., Swensson, E., 1991. Subduction and eduction of continental crust; major mechanisms during continent–continent collision and orogenic extensional collapse. Terra Nova 3, 303 – 310. Andersen, T.B., Osmundsen, P.T., Jolivet, L., 1994. Deep crustal fabrics and a model for the extensional collapse of the southwest Norwegian Caledonides. Journal of Structural Geology 16 (9), 1191 – 1203. Andersen, T.B., Berry, H.N., Lux, D.R., Andresen, A., 1998. The tectonic significance of pre-Scandian Ar-40/Ar-39 phengite cooling ages in the Caledonides of western Norway. Journal of the Geological Society 155, 297 – 309. Austrheim, H., Engvik, A., 1997. Fluid transport, deformation and metamorphism at depth in a collision zone. In: Jamtveit, B., Yardley, B.W.D. (Eds.), Fluid Flow and Transport in Rocks: Mechanisms and Effects. Chapman and Hall, London, pp. 123 – 137. Carswell, D.A., Cuthbert, S.J., Krabbendam, M., Medaris Jr., L.G., Bruekner, H.K., 2003. Guidebook to the Field Excursions in the Nordfjord–Stadtlandet–Almklovdalen Area. Geological Survey of Norway, Trondheim, p. 137. Castelli, D., Rolfo, F., Compagnoni, R., Xu, S.T., 1998. Metamorphic veins with kyanite, zoisite and quartz in the Zhu-Jia-Chong eclogite, Dabie Shan, China. Island Arc 7, 159 – 173. Chauvet, A., Dallmeyer, R.D., 1992. 40Ar/39Ar mineral dates related to Devonian extension in the southwestern Scandinavian Caledonides. Tectonophysics 210, 155 – 177. Chopin, C., 1987. Very-high-pressure metamorphism in the Western Alps: implications for subduction of continental crust. Philosophical Transactions of the Royal Society of London. A 321, 183 – 197. Cobbold, P., Watkinson, A., 1981. Bending anisotropy: a mechanical constraint on the orientation of fold axes in an anisotropic medium. Tectonophysics 72, T1 – T10. Coward, M.P., Potts, G.J., 1983. Complex strain patterns developed at the frontal and lateral tips to shear zones and thrust zones. Journal of Structural Geology 5 (3/4), 383 – 399. Eide, E., Torsvik, T., Andersen, T.B., Arnaud, N.O., 1999. Early Carboniferous unroofing in Western Norway: a tale of alkali feldspar thermochronology. Journal of Geology 107, 353 – 374. Ellis, D.J., Green, D.H., 1979. An experimental study of the effect of Ca upon garnet–clinopyroxene Fe–Mg exchange equilibria. Contributions to Mineralogy and Petrology 71, 13 – 22. Engvik, A., Andersen, T., 2000. Evolution of Caledonian deformation fabrics under eclogite and amphibolite facies at Vardalsneset, Western Gneiss Region, Norway. Journal of Metamorphic Geology 18, 241 – 257. Engvik, A., Austrheim, H., Andersen, T.B., 2000. Structural, mineralogical and petrophysical effects on deep crustal

R. Foreman et al. / Tectonophysics 398 (2005) 1–32 rocks of fluid-limited polymetamorphism, Western Gneiss Region, Norway. Journal of the Geological Society 157 (1), 121 – 134. Engvik, A., Austrheim, H., Erambert, M., 2001. Interaction between fluid flow, fracturing and mineral growth during eclogitisation, an example from the Sunnfjord area, Western Gneiss Region, Norway. Lithos 57, 111 – 141. Escher, A., Watterson, J., 1974. Stretching fabrics, folds and crustal shortening. Tectonophysics 22, 223 – 231. Foreman, R., Wheeler, J. (in preparation). Fossen, H., 1992. The role of extensional tectonics in the Caledonides of south Norway. Journal of Structural Geology 14 (8–9), 1033 – 1046. Fossen, H., 1993. Linear fabrics in the Bergsdalen Nappes, southwest Norway: implications for deformation history and fold development. Norsk Geologisk Tiddskrift 73 (2), 95 – 108. Gebauer, D., Lappin, M., Grunenfelder, M., Wyttenbach, A., 1985. The age and origin of some Norwegian eclogites: a U–Pb zircon and REE study. Chemical Geology 52, 227 – 247. Goscombe, B.D., Passchier, C.W., 2003. Asymmetric boudins as shear sense indicators—an assessment from field data. Journal of Structural Geology 25 (4), 575 – 589. Graham, C.M., Powell, R., 1984. A garnet–hornblende geothermometer: calibration, testing, and application to the Pelona Schist, Southern California. Journal of Metamorphic Geology 2, 13 – 31. Griffin, W.L., Brueckner, H.K., 1980. Caledonian Sm–Nd ages and a crustal origin for Norwegian eclogites. Nature 285, 319 – 321. Grujic, D., Mancktelow, N., 1995. Folds with axes parallel to the extension direction: an experimental study. Journal of Structural Geology 17 (2), 279 – 291. Hacker, B.R., Andersen, T.B., Root, D.B., Mehl, L., Mattinson, J.M., Wooden, J.L., 2003. Exhumation of high-pressure rocks beneath the Solund Basin, Norway. Journal of Metamorphic Geology 21, 613 – 629. Hacker, B.R., Root, D.B., Walsh, E.O., Young, D., Johnston, S.M., Andersen, T.B., 2004. Genesis and exhumation of ultrahigh pressure rocks in Norway. (Abstract) Proceedings of the 32nd International Geological Congress (32IGC), Florence, Italy. Holland, T.J.B., 1980. The reaction albite=jadeite+quartz determined experimentally in the range 600–1200 8C. American Mineralogist 65, 129 – 134. Jolivet, L., Daniel, J.M., Truffert, C., Goffe, B., 1994. Exhumation of deep-crustal metamorphic rocks and crustal extension in arc and back-arc regions. Lithos 33, 3 – 30. Koons, P., 1984. Implications to garnet–clinopyroxene geothermometry of non-ideal solid solution in jadeitic pyroxenes. Contributions to Mineralogy and Petrology 88, 340 – 347. Krabbendam, M., 1998. Structural and metamorphic evolution of eclogite gneisses during exhumation in SW Norway. Unpublished PhD thesis, University of Oxford. Krabbendam, M., Dewey, J., 1998. Exhumation of UHP rocks by transtension in the Western Gneiss Region, Scandinavian Caledonides. In: Holdsworth, R., Strachan, R., Dewey, J. (Eds.), Continental Transpressional and Transtension Tectonics, vol. 135. Geological Society of London, London, pp. 159 – 181.

31

Krabbendam, M., Wain, A.L., 1997. Late-Caledonian structures, differential retrogression and structural position of (ultra) high pressure rocks in the Nordfjord–Stadlandet area, Western Gneiss Region. Norges Geologiske Undersbkelse Bulletin 432, 127 – 139. Krabbendam, M., Wain, A.L., Andersen, T.B., 2000. Pre-Caledonian granulite and gabbro enclaves in the Western Gneiss Region, Norway: indications of incomplete transition at high pressure. Geological Magazine 137 (3), 235 – 255. Kullerud, L., Torudbakken, B.O., Ilebekk, S., 1986. A compilation of radiometric age determinations from the Western Gneiss Region, South Norway. Norges Geologiske Undersbkelse Bulletin 406, 17 – 42. Labrousse, L., Jolivet, L., Agard, P., Hebert, R., Andersen, T.B., 2002. Crustal Scale boudinage and magmatization of gneiss during their exhumation in the UHP Province of Western Norway. Terra Nova 14 (4), 263 – 270. Labrousse, L., Jolivet, L., Andersen, T.B., Agard, P., Maluski, H., Schaerer, U., in preparation. Pressure–Temperature–Time– Deformation history of the exhumation of ultra-high-Pressure rocks in the Western Gneiss Region, Norway. Leake, B., Wooley, A., Arps, C., Birch, W., Gilbert, M., Grice, J., Hawthorne, F., Kato, A., Kisch, H., Krivovichev, V., Linthout, K., Laird, J., Mandarino, J., Maresch, W., Nickel, E., Rock, N., Schumacher, J., Smith, D., Stephenson, N., Ungaretti, L., Whittaker, E., Youzhi, G., 1997. Nomenclature of amphiboles: report of the Subcommittee on Amphiboles of the International Mineralogical Association, Commission on New Minerals and Mineral Names. American Mineralogist 82, 1019 – 1037. Liou, J.G., Tsujimori, T., Zhang, R.Y., Katayama, I., Maruyama, S., 2004. Global UHP metamorphism and continental subduction/ collision: The Himalayan model. International Geology Review 46 (1), 1 – 27. Milnes, A.G., Wennberg, O.P., Skar, O., Koestler, A.G., 1997. Contraction, extension and timing in the South Norwegian Caledonides: the Sognefjord transect. In: Burg, J.P., Ford, M. (Eds.), Orogeny Through Time, vol. 121. Geological Society of London, London, pp. 123 – 148. Morimoto, N., Fabries, J., Ferguson, A.K., Ginzburg, I.V., Ross, M., Seifert, F.A., Zussman, J., Aoki, K., Gottardi, G., 1988. Nomenclature of pyroxenes. American Mineralogist 73, 1123 – 1133. Mbrk, M.B.E., 1985. A gabbro to eclogite transition on Flemsby, Sunnmbre, western Norway. Chemical Geology 50, 283 – 310. Norton, M., 1987. The Nordfjord–Sogn Detachment, W. Norway. Norsk Geologisk Tiddskrift 67, 93 – 106. Osmundsen, P.T., Andersen, T.B., 2001. The Middle Devonian basins of western Norway: sedimentary response to large-scale transtensional tectonics? Tectonophysics 332, 51 – 68. Powell, R., 1985a. Geothermometry and geobarometry: a discussion. Journal of the Geological Society of London 142, 29 – 38. Powell, R., 1985b. Regression diagnostics and robust regression in geothermometer/geobarometer calibration: the garnet–clinopyroxene geothermometer revisited. Journal of Metamorphic Geology 3, 231 – 243.

32

R. Foreman et al. / Tectonophysics 398 (2005) 1–32

Reddy, S.M., Wheeler, J., Cliff, R.A., 1999. The geometry and timing of orogenic extension: an example from the Western Italian Alps. Journal of Metamorphic Geology 17, 573 – 589. Rey, P., Burg, J.P., Casey, M., 1997. The Scandinavian Caledonides and their relationship to the Variscan belt. In: Burg, J.P., Ford, M. (Eds.), Orogeny Through Time, Geological Society, London, Special Publication, vol. 121, pp. 179 – 200. Rubie, D.C., 1990. Role of kinetics in the formation and preservation of eclogites. In: Carswell, D.A. (Ed.), Eclogite Facies Rocks. Blackie, Glasgow, pp. 111 – 140. Sanderson, D.J., 1972. The development of fold axes oblique to the regional trend. Tectonophysics 16, 55 – 70. Skjernaa, L., 1980. Rotation and deformation of randomly oriented planar and linear structures in progressive simple shear. Journal of Structural Geology 2 (1/2), 101 – 109. Terry, M.P., Robinson, P., Hamilton, M.A., Jercinovic, M.J., 2000a. Monazite geochronology of UHP and HP metamorphism, deformation, and exhumation, Nordoyane, Western Gneiss Region, Norway. American Mineralogist 85, 1651 – 1664. Terry, M.P., Robinson, P., Ravna, E.J.K., 2000b. Kyanite eclogite thermobarometry and evidence for thrusting of UHP over HP metamorphic rocks, Nordoyane, Western Gneiss Region, Norway. American Mineralogist 85, 1637 – 1650. Torsvik, T.H., Sturt, B.A., Ramsay, D.M., Kisch, H.J., Bering, D., 1986. The tectonic implications of Solundian (Upper Devonian) magnetization of the Devonian rocks of Kvamshesten, Western Norway. Earth and Planetary Science Letters 80, 337 – 347.

Tucker, R.D., Krogh, E.J., Raheim, A., 1990. Proterozoic evolution and age-province boundaries in the central part of the Western Gneiss Region, Norway: results of U–Pb dating of accessory minerals from Trondheinsfjord to Geiranger. In: Gower, C.F., Rivers, T., Ryan, B. (Eds.), Mid-Proterozoic Geology of the Southern Margin of Proto-Laurentia-Baltica, Geological Association of Canada, Special Publication, vol. 38, pp. 149 – 173. Wain, A., 1997. New evidence for coesite in eclogite and gneisses: defining an ultrahigh-pressure province in the Western Gneiss Region of Norway. Geology 25 (10), 927 – 930. Wain, A.L., Waters, D.J., Austrheim, H., 2001. Metastability of granulites and processes of eclogitisation in the UHP region of western Norway. Journal of Metamorphic Geology 19, 609 – 625. Waters, D.J., Martin, H.N., 1993. Geobarometry in phengite-bearing eclogites. Terra Abstracts 5, 410 – 411 (updated calibration at http://www.earth.ox.ac.uk/davewa/ecbar.html). Watkinson, A., 1975. Multilayer folds initiated in bulk plane strain, with the axis of no change perpendicular to the layering. Tectonophysics 28, T7 – T11. Wayte, G.J., Worden, R.H., Rubie, D.C., Droop, G.T.R., 1989. A TEM study of disequilibrium plagioclase breakdown at high pressure: the role of infiltrating fluid. Contributions to Mineralogy and Petrology 101, 426 – 437. Wheeler, J., Reddy, S.M., Cliff, R.A., 2001. Kinematic linkage between internal zone extension and shortening in more external units in the NW Alps. Journal of the Geological Society, London 158, 439 – 443.