Eclogite meso- and microfabrics: implications for the burial and exhumation history of eclogites in the Tauern Window (Eastern Alps) from P-T-d paths

TECTONOPHYSICS ELSEVIER

Tectonophysics 285 (1998) 183-209

Eclogite meso- and microfabrics: implications for the burial and exhumation history of eclogites in the Tauern Window (Eastern Alps) from P - T - d paths Walter Kurz a,*, Franz Neubauer a, Edgar Dachs b ~ lnstitut fiir Geologie und Paliiontologie, Paris-Lodron-Universitiit Salzburg, Hellbrunner Strq[3e 34, A-5020 Salzburg, Austria I~lnstitut fiir Mineralogie, Paris-Lodron- Universitiit Salzburg, Hellbrunner StraJ3e 34, A-5020 Salzburg, Austria

Received 24 April 1997; accepted 12 June 1997

Abstract The Eclogite Zone of the central-southern Tauern Window comprises eclogites and associated high-pressure metasediments that are intercalated between Penninic basement units in the foot-wall and an imbricate stack in the hanging-wall. This nappe stack consists of continental basement units, cover sequences of a distal continental slope, and the main part of the ophiolitic Glockner Nappe. Textures and microfabrics of eclogites from the central-southern Tauern Window allow the establishment of the eclogitic and post-eclogitic deformation histories from burial by subduction to subsequent exhumation. The metamorphic evolution of the eclogites is documented by: (1) initial greenschist to blueschist facies; (2) eclogite facies at ca. 550-630°C and -t-20 kbar; (3) a second blueschist facies overprint (ca. 450°C, 10-15 kbar); and (4) exhumation to pressures of ca. 6 - 7 kbar at 500-550°C (upper greenschist to lower amphibolite facies). An eclogitic penetrative foliation (S~) and S-dipping stretching lineation (Ll) formed during the final stages of subduction of the eclogite-bearing unit. The P - T evolution of the eclogites during D~ documents the final increment of the prograde P - T path. Later blueschist facies metamorphism is contemporaneous to top-to-the-N emplacement of the eclogite-bearing unit onto continental basement units within a subduction zone. We discuss two models of eclogite emplacement. Most probably, emplacement is achieved by a thrust along the base of the eclogite-bearing unit. Alternatively, the P - T path of the eclogite facies rocks suggests an emplacement model similar to corner flow. Subsequent to eclogite and nappe emplacement, the nappe stack is entirely affected by a penetrative deformation (D2) that resulted in the development of a mylonitic foliation ($2) and an E-W-oriented stretching lineation (L2) within lower amphibolite to upper greenschist facies. This deformation is related to a crustal-scale detachment zone, that formed during the exhumation of the Penninic units. © 1998 Elsevier Science B.V. All rights reserved. Keywords: eclogites; microfabric; P-T-d path; Tauem Window; Eastern Alps

* Corresponding author. Present address: Institut ftir Geologie und Pal~iontologie, Karl-Franzens-Universitfit Graz, Heinrichstrage 26, A-8010 Graz, Austria. Tel.: +43-316-380 8725; Fax: +43-316-380 9870; E-mail: [email protected] 0040-1951/98/$19.00 © 1998 Elsevier Science B.V. All rights reserved. PII S 0 0 4 0 - 1 9 5 1 ( 9 7 ) 0 0 1 8 8 - 1

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W. Kurz et al. / Tectonophysics 285 (1998) 183-209

1. Introduction Many eclogites and eclogite facies rocks are generally interpreted as oceanic and/or continental crustal material buried to mantle depths. Such eclogite facies assemblages are recorded from the following geodynamic settings: (1) oceanic lithosphere subducted along an anomalous low geotherm (Ernst, 1971, 1973, 1975, 1988); (2) continental lithosphere subducted along an anomalous geotherm reaching epidote-amphibolite facies conditions subsequent to high-pressure metamorphism (e.g., Carswell, 1990); and (3) subcontinental mantle lithosphere (Carswell, 1990). Eclogite facies metamorphism is reached at lithospheric levels of up to 70-100 km depth and, depending on the subduction angle, at horizontal distances of up to 200 km and more from the oceanic trench (Peacock, 1993). So, several mechanisms may contribute to remove material in the hanging-wall of the eclogite facies rocks within an ancient subduction zone, and to emplace these rocks at shallow lithospheric levels. Different mechanisms of exhumation of high-pressure rocks produce different patterns of flow, and different paths for their exhumation, and, therefore, different P - T - d paths (e.g., Platt, 1993). Several mechanisms have been called for eclogite exhumation (for review, see Platt, 1986, 1993). These include: (1) continued underplating and overthickening of an accretionary wedge is compensated by extension at the surface in terms of normal faulting (Platt, 1986, 1993; Henry et al., 1993); this allows previously underplated material to rise to the surface and thus the exposure of the high pressure-low temperature assemblages; (2) shift from subduction of oceanic lithosphere to collision and attempted subduction of buoyant sialic crust; this results in rapid exhumation, uplift and denudation of the HP/LT assemblages; (3) the corner flow model developed by Cloos (1982, 1985) and Shreve and Cloos (1987), which is based on the mechanics of flow melanges that are constrained by a low-angle comer;

185

(4) the return of buoyant material to the surface at the cessation of subduction, carrying eclogite blocks that were too small to sink rapidly through it (England and Holland, 1979); (5) removal of material by wrench faulting causes the juxtaposition of rocks with different burial histories (Behrmann and Ratschbacher, 1989). The eclogites of the Tauem Window are very well investigated from the petrological and petrographical point of view (e.g., Holland, 1979a,b; Dachs, 1986, 1990; Frank et al., 1987a; Droop et al., 1990; Selverstone et al., 1992; Getty and Selverstone, 1994). The structural investigations concerning the deformation paths of eclogite facies rocks, combined with varying metamorphic conditions are rather scarce or concentrated more on surrounding rock assemblages than on the eclogites (e.g., Behrmann and Ratschbacher, 1989). Furthermore, models for the burial and exhumation history of these eclogites were either based on petrological data (e.g., Selverstone, 1985; Frank et al., 1987a,b; Selverstone et al., 1992), or simply on structural investigations (Behrmann and Ratschbacher, 1989). The scope of this study is to combine the textural and microfabric evolution of eclogites with P - T data. This gives access to a more detailed eclogitic and post-eclogitic deformation history from burial by subduction to subsequent exhumation.

2. Geological setting of the Tauern eclogites The Tauem Window (Fig. 1) exposes Penninic units in the foot-wall of the Austroalpine nappe complex, which forms the hanging-wall plate during Late Cretaceous and Tertiary plate collision. The Penninic units within the Tauern Window are separated into several nappes, that are characterized by typical lithofacial assemblages (Frisch, 1974, 1975; Tollmann, 1975, 1977; Frank et al., 1987a; Kurz et al., 1996) (Fig. 1). From the foot-wall to the hanging-wall, the nappe stack includes: (1) the Venediger Nappe and the Wolfendom Nappe comprising a pre-Variscan basement intruded

Fig. 1. (a) Tectonic map of the Eastern Alps with location of the Tauern Window (TW). (b) Tectonic map of the Tauern Window with structural position of the Eclogite Zone (A-A': position of cross-section in (c)). (c) Cross-section across the central part of the Tauern Window.

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W. Kur= et al./Tectonophysics 285 (1998) 183 209

by Variscan granitoids (the Zentralgneis) and a cover sequence of Jurassic metacarbonates (Hochstegen Marble Fm.), and Cretaceous metapelites and metapsammites (Kaserer Group); (2) the Storz and Riffl Nappes comprising Variscan and Alpidic polymetamorphic basement rocks covered by the late Palaeozoic (?) or Cretaceous (?) metapelites and graphitic quartzites of the MurtOrl Group; (3) the E c l o g i t e Z o n e which is restricted to the central southern Tauern Window and is characterized by a Mesozoic volcano-sedimentary sequence of a distal continental slope; (4) the Rote Wand-Modereck Nappe, formed by basement rocks of the Rote Wand-Modereck Lamellae that are covered by the Permian to Triassic quartzites and Triassic metacarbonates, Jurassic breccias, calcareous micaschists and metatuffs as well as Cretaceous metapelites and metapsammites; (5) the Glockner Nappe, comprising an oceanic basement (serpentinites and ultramafic rocks) and a partly incomplete ophiolithic sequence; (6) the Matrei Zone, interpreted to represent an accretionary wedge that is characterized by metamorphic flysch sediments (mainly calcareous and carbonate-free micaschists), breccias and olistolithes mainly of Austroalpine derivation (Frisch et al., 1987); and (7) the Klammkalk Zone, comprising calc-schists, massive marbles and minor developed green phyllites, forming a low-grade metamorphic equivalent to the 'Biindnerschiefer' within the Glockner Nappe. These units are surrounded by Lower Austroalpine units. In the central-southern part of the Tauern Window (Fig. lb, Fig. 2), a slice of eclogites and associated high-pressure eclogite facies metasediments is intercalated between pre-Alpine basement rocks and a basement-cover imbricate stack. These were overthrust by the ophiolitic Glockner Nappe (Miller et al., 1980; Frank et al., 1987a; Kurz et al., 1996). The Eclogite Zone comprises mafic eclogites of tholeiitic and mildly alkaline chemical composition (Miller, 1974, 1977, 1987). Protoliths are assumed to represent basalts of an intraplate character (H6ck and Miller, 1987). These rocks are associated with eclogite facies meta-tuffs and coarse-grained eclogite fa-

ties metagabbros. Eclogitic lenses and boudins of a few centimetres to 100 m of scale are intercalated with metasediments. Especially along tectonic contacts the eclogites are strongly boudinaged. The eclogites, in places, contain relics of pillow structures (Miller et al., 1980). They are often retrogressed to garnet-amphibolites and garnet-bearing greenschists due to overprint by Barrovian-type, greenschist to amphibolite facies metamorphism. The degree of retrogression varies laterally as well as vertically. The associated metasediments, like quartzites, paragneisses, garnet-micaschists, calcareous micaschists, and calcitic and dolomitic marbles, form a typical continental margin sequence. These metasediments experienced the same high-pressure metamorphism (Franz and Spear, 1983; Chopin and Schreyer, 1983; Dachs, 1986, 1990; Spear and Franz, 1986).

3. Metamorphic evolution The rocks exposed within the Eclogite Zone experienced a polyphase metamorphic evolution. Inclusions in garnets, which are sometimes interpreted as pseudomorphs after lawsonite, document a first stage of metamorphism at ca. 400°C (Miller, 1977, 1986; Frank et al., 1981, 1987a). Eclogite facies metamorphism is only observed clearly within the Eclogite Zone. The eclogite facies rocks were buried to a depth of at least 65 km (20 kbar, ±600°C; Holland, 1979a,b; Dachs, 1986, 1990; Frank et al., 1987a; Droop et al., 1990; Selverstone et al., 1992; Zimmermann et al., 1994; Getty and Selverstone, 1994). Within the Glockner Nappe eclogite facies assemblages are just locally observed (Sturm et al., 1996; Kurz et al., 1996). The entire nappe pile in the Tauern Window was subsequently affected by blueschist facies metamorphism. Within the Eclogite Zone pressures of 7 - 9 kbar and temperatures of ca. 450°C are estimated by Raith et al. (1980); 450°C, 10-15 kbar are estimated by Holland (1979b) and Zimmermann et al. (1994), but the P - T data are not well constrained due to the subsequent strong overprint by Barrovian-type metamorphism. Within the other tectonic units peak pressures of up to 1012 kbar have been evaluated (Holland and Richardson, 1979; Selverstone et al., 1984, 1992; Cliff et al., 1985; Droop, 1985; Holland and Ray, 1985;

W. Kurz et al./Tectonophysics 285 (1998) 183-209

Frank et al., 1987a; Behrmann and Ratschbacher, 1989; Behrmann, 1990; Selverstone, 1993). Finally, the entire nappe pile was affected by Barrovian-type upper greenschist to lower amphibolite facies metamorphism (e.g., Frank et al., 1987a; Selverstone, 1993). The high-pressure metamorphic assemblages within the Eclogite Zone were suspected to form a 'paired metamorphic belt' together with Cretaceous middle-pressure metamorphic sequences within Austroalpine units, according to Ernst (1971, 1975) (e.g., Frank, 1987; Frank et al., 1987a; Wallis et al., 1993). However, no direct evidence of a Cretaceous age for high-pressure metamorphism within the Penninic realm is documented by these authors. In contrast to supposed Cretaceous ages, phengite 4°Ar-39Ar mineral ages of ca. 36-32 Ma (Zimmermann et al., 1994) and 38 Ma (L. Ratschbacher, pers. commun., 1995) from the Eclogite Zone are interpreted to represent cooling ages subsequent to Eocene blueschist facies metamorphism (Zimmermann et al., 1994). The possibility of a younger age of high-pressure metamorphism is also elucidated by Inger and Cliff (1994). At the moment it is not clear if argon isotopic systems of high-pressure phengite survived later thermal metamorphism (ca. 500-550°C) in excess of the argon retention temperature of ca. 410°C (e.g., von Blanckenburg et al., 1989).

4. Eclogite fabrics 4.1. Mesoscale fabrics

Generally, two structural types of eclogites are discernable in the Eclogite Zone: (1) coarse-grained massive eclogites with a grainsize of up to 1 cm; (2) fine- to medium-grained, foliated eclogites and eclogite-mylonites. Several types of eclogites are discernable petrographically (the petrographic characterization follows Miller, 1974, 1977; Frank et al., 1981). ( 1) Coarse-grained eclogites displaying relict gabbroic textures; omphacite grains reach 1 cm in size (omphacitel) and are supposed to be pseudomorphs after augite; the omphacites are partly surrounded by garnets up to 0.5 mm in size. (2) Porphyroclastic eclogites with omphacitel of up to 0.5 cm and garnet of ca. 0.2-0.3 cm within a very fine-grained matrix (grain size <0.1 mm)

187

in places; deformed relics of omphacitel are surrounded by fine-grained omphacite2 crystals ('cataclastic eclogites' after Miller, 1977); these omphacite2 crystals represent recrystallized grains. (3) Kyanite and talc-beating banded eclogites. (4) Epidote-eclogites: strongly foliated eclogitic mylonites with a well developed mylonitic foliation and stretching lineation defined by elongate garnet, dynamically recrystallized omphacite2, zoisite, kyanite, epidote/zoisite, and glaucophane. Locally, a compositional layering of single grain thick quartz layers, garnet and omphacite is developed (Fig. 3a). (5) Very fine-grained, compositionally layered, banded eclogitic mylonites. Syn-eclogitic D1 deformation structures ($1; Lj) (Fig. 2a) are difficult to reconstruct because the eclogites are often retrogressed to garnet-amphibolites and garnet-bearing greenschists during exhumation. The degree of retrogression of the eclogites is irregularly distributed as well within the Eclogite Zone as within individual eclogite bodies. Where the stretching and/or mineral lineation L1 is defined by HP minerals (e.g., garnet, omphacite, kyanite, glaucophane), its orientation is variable in places. A relative maximum of S- to SW-dipping LI mineral lineations, especially of omphacite and glaucophane, is developed in the exposures of the Dorfertal, Timmeltal, and Frosnitztal (Fig. 2a). The exposures near the 'Raneburger See' exhibit LI lineations that are mainly defined by elongated garnet. Here the LI orientation data scatter along a N-trending great circle. Very often the eclogites are cross-cut by kyanitebearing veins and shear planes with kyanite slickensides. The variation of Ll orientation might be related to variable rotation of eclogite slices and boudins during later greenschist to amphibolite facies deformation (D2). The eclogitic foliation is transposed subparallel to the penetrative foliation $2 (Fig. 2c) that was developed within the upper greenschist to lower amphibolite facies metamorphic conditions. Eclogites that are retrogressed during D2 (type 4 after Miller, 1974, 1977) are characterized by a very well developed penetrative mylonitic foliation $2 (Fig. 2), defined by the subparallel arrangement of actinolitic hornblende, epidote, chlorite and locally biotite, and an also well-developed, E-trending, subhorizontal stretching lineation L2 (Fig. 2), defined

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W. Kurz et al./Tectonophysics 285 (1998) 183-209

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primarily by actinolitic hornblende and plagioclase (An < 10%). $2 and L2 form the penetrative fabrics in the foot-wall (Venediger Nappe, Riffl Nappe) and hanging-wall (Rote Wand-Modereck Nappe, Glockner Nappe) tectonic units, too. The penetrative foliation is oriented subparallel to principal tectonic contacts with country rocks. Ol structures are transposed subparallel to $2 resulting in the development of a composite foliation S1,2. In places, the eclogitic foliation is gently folded subperpendicular to L2, but tight and isoclinal folds are developed continuously with their axes subparallel to the L2 stretching lineation (Fig. 2). In metapelites and metacarbonates, the S orientation is close to that of the shear planes (C). It is transected by C' and less by C" shear bands at scales of centimetres, decimetres and metres. These sense-of-shear indicators (shear bands and extensional crenulation cleavage according to the definition of Platt and Vissers, 1980), document top-to-the-W displacement. Symmetrically boudinaged eclogite layers (Fig. 3b) locally document pure shear deformation as well as strain partitioning during D2. The whole lithological sequence is affected by E-trending isoclinal folds B3 and open to tight E-W-striking B4 folds with subhorizontal axes. This is documented in girdle distributions of the poles to the penetrative foliation S1,2 (Fig. 2a,

c). A final deformation is documented by km-scale (map-scale) N-trending open folds.

4.2. Microfabrics The eclogites of the central-southern Tauern Window exhibit textures and microfabrics that document the pre-eclogitic, eclogitic and post-eclogitic metamorphic and deformational evolution of these rocks. A summary of the mineral assemblages that are characteristic for the distinct phases of deformation and metamorphism is given in Table 1. Pre-eclogite facies assemblages are only documented by garnet inclusions. These are mainly quartz, chlorite, biotite, epidote/zoisite, +phengite, and locally glaucophane in garnet cores. Omphacite, kyanite, quartz, paragonite, phengite, zoisite, glaucophane and talc occur closer to the garnet rims and already form eclogite facies assemblages. The coarse-grained massive eclogites (type I; type 1 after Miller, 1974, 1977) contain many coarsegrained omphacitel crystals that are characterized by strong undulatory extinction, the development of subgrains and polygonalization. They are often surrounded by dynamically recrystallized omphacite2 grains of 4-0.5 mm in size (Fig. 4a). In some domains, the large omphacite crystals are surrounded

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W. Kur: et al./Tectonophysics 285 (1998) 183-209

Fig. 3. Outcrop-scale structures from the Eclogite Zone. (a) Fine- to medium-grained eclogitic mylonite with tectonic layering of omphacite and garnet; 100 m W of WeiBspitze (3300 m), Timmeltal. (b) Eclogite boudins in a matrix of marble mylonite, calcareous micaschists and metapelites; eastern shore of Eissee (2550 m), Timmeltal.

W. Kurz et al./Tectonophysics 285 (1998) 183-209

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Table 1 S u m m a r y o f m i n e r a l a s s e m b l a g e s that are c h a r a c t e r i s t i c f o r t h e d i s t i n c t p h a s e s o f d e f o r m a t i o n a n d m e t a m o r p h i s m Inclusions pre-HP and pre-Di

HP inclusions

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Garnet Omphacite I Omphacite2

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Phengite I

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post-Dr

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Paragonite

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Rutile

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Kyanite

x

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Barroisite 1 Barroisite2

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Taramite

x

x

x

x

Hornblende Zoisite

x

D2; u p p e r greenschist facies

post-D2

x

x x

HP cracks

x x x

Phengite2

Chloritoid

Blueschist facies

x x

x

x

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Epidote

x

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Actinolite Chlorite

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Biotite

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Albite (An
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Sphene Symplectite

by an aggregate of fine-grained omphacite2, glaucophane and also paragonite. Omphacite2 formed with respect to increasing secondary grain size reduction. Within other domains, omphacitel crystals turn into symplectite along their margins, which consists of fine-grained aggregates of diopsidic clinopyroxene and albite. It is hard to evaluate whether these symplectite rims are former recrystallized omphacite2 crystals. The microfabrics of mylonitic banded and foliated eclogites (type II; types 2, 3, 5 after Miller, 1974, 1977) are characterized by a shape preferred orientation of omphacite2, and a strong crystallographic preferred orientation of omphacite2 which is already indicated by the regular extinction of the grains. Kyanite, glaucophane/barroisitel, zoisite and rutile are also arranged with their long axis subparallel to the eclogite facies penetrative foliation S i (Fig. 4b). Coarse-grained phengitel was developed synkinematically. S l is assimilated to the shear plane (C) of an S-C fabric (Berth6 et al., 1979). Elongate garnets (Fig. 4b) are partly surrounded by asymmetric pressure shadows that are filled by recrystallized

x

x x

x

omphacite, glaucophane/barroisitel, epidote/zoisite, and phengite, and a-shaped garnet porphyroclasts are rather scarce. These document top-to-the-N to -NE ductile shearing (Fig. 4c) during Dj. Garnet is very often cracked (Fig. 4d) or boudinaged with cracks oriented subperpendicularly to $1. The cracks are filled by HP assemblages of kyanite, zoisite, omphacite, and quartz, and furthermore by glaucophane, phengite and carbonate (e.g., Thomas and Franz, 1989). Samples that are affected by blueschist facies metamorphism display reaction zones between garnet and omphacite. These zones primarily consist of barroisitic to glaucophanitic hornblende (barroisite2). In places the elongated omphacite2 crystals that are aligned subparallel to the eclogitic penetrative foliation, are either surrounded by rims of subcalcic amphibole or barroisite2 (Table 3). At more advanced stages omphacite is completely replaced by barroisite2, and by subcalcic amphibole (Miller, 1974). Due to the instability of omphacite we interpret these amphiboles (barroisite2) to have formed at conditions that are different from the conditions

W. Kur= et al./Tectonophysics 285 (1998) 183-209

of omphacite formation at the pressure peak. In places, omphacitel is partly replaced by paragonite and epidote; kyanite in paragenesis with omphacite is unstable, probably due to a reaction of the type: omphacite ÷ kyanite = paragonite + epidote In distinct domains the DI fabrics have not been transposed during D2. Here the S I penetrative fofiation is overgrown by zoisite/epidote, amphibole, chlorite and phengite during the thermal peak of upper greenschist to lower amphibolite facies retrogression. The microfabrics of eclogites that are strongly overprinted by upper greenschist to amphibolite facies metamorphism (type 4 after Miller, 1974, 1977) are characterized by symplectite, which consists of fine-grained aggregates of diopsidic pyroxene and albite. These eclogites show a shape preferred orientation of actinolitic hornblende, epidote/zoisite, and white mica that form a penetrative foliation $2. Phengitel is instable; in some domains it turns into finegrained phengite2 along its edges. Strain shadows are arranged symmetrically around garnet; within some domains, asymmetric strain shadows document a top-to-the-W movement (Kurz et al., 1996). The pressure shadows are filled with green hornblendes, epidote/zoisite, quartz; chlorites are subordinate constituents. The garnets show inclusions of epidote, amphibole and white mica. The penetrative, partly symplectitic foliation is again overgrown by another generation of epidote and white mica. Some eclogites only show distinct domains of symplectitic shear zones. The mineral assemblages within these shear zones are similar to the assemblages described above and document decreasing pressure conditions. As a rule, garnet is missing within these shear zones. Rare cataclastic shear zones cross-cut the penetrative mylonitic foliation (Fig. 4e) showing a top-tothe-W sense of shear. In places, conjugate sets of cataclastic shear zones are developed that are asymmetrically arranged with respect to the penetrative foliation. They indicate a top-to-the-E sense of shear being associated with extensional cracks that are mainly filled with amphibole and chlorite. The later are oriented subperpendicularly to the penetrative foliation. Garnets, pyroxenes and epidotes are cracked, too. The cracks are locally oriented obliquely to the

193

foliation at angles between 60 ° and 85 °, indicating a W-directed sense of shear.

5. Thermobarometry 5.1. MethodsJor mineral chemistry and thermobarometric calculations The mineral chemistries of several mineral phases were analyzed with an JEOL JXA-8600 electron microprobe at the Department of Geology and Palaeontology of the University of Salzburg (measuring conditions: 15 kV acceleration voltage; 40 nA beam current; beam focused to 1 /zm, and 5 # m for mica analyses, with internal ZAF correction of the raw data). The detailed zonation profiles of garnets have been evaluated by measuring between 150 and 300 points along a defined line across a single grain. Structural formulae of minerals have been evaluated using unpublished working sheets for Microsoft Excel (version 5.0). We preferably chose samples documenting clear criteria of eclogite facies deformation, especially a penetrative foliation and stretching lineation defined by eclogite facies minerals, in order to constrain the relationships between deformational and petrological features. Fe 3+ concentrations of pyroxenes and amphiboles have been evaluated after Droop (1987). Representative results of microprobe analyses are listed in Tables 2 4 ; structural formulae and Kd-values that have been used for P - T calculations are given separately in Tables 5-10. Various geothermometers and geobarometers have been applied to estimate P - T conditions during distinct stages of the textural evolution of eclogites. We tried to use independent methods in order to prove the reliability of several results. Thermobarometric calculations have been carried out using the program packages of THERMOCALC version 2.2b3 (Holland and Powell, 1985, 1990; Powell and Holland, 1985, 1988, 1994), and PT-MAFIC version 2.0 (Soto, 1993; Soto and Soto, 1995), that includes several geothermometric and geobarometric calibrations. Geobarometric calculations are based on the experimental calibrations by Massone and Schreyer (1987, 1989) for phengite, and Holland (1980, 1983) for omphacite. For the jadeite + quartzgeobarometer temperatures between 450 and 650°C have been assumed; 600°C, which resulted from

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Table 4 Results of microprobe analyses (oxide weight% and structural formulae) for rock-forming eclogite facies minerals (continued) Oxide %

Talc WK539

Talc WK541

Talc WK561

Epidote core WK539

Epidote rim WK539

SiO2 TiO2 A1203 Cr203 Fe203 FeO MnO MgO NiO CaO Na20 K20 F CI H20+ Total No. of oxygens

62.00 0.00 0.53

60.20 0.01 0.58

61.20 0.03 0.35

38.40 0.13 29.20

38.10 0. I 1 27.80

3.60 total 0.01 29.10

3.32 total 0.03 27.60

3.37 total 0.00 28.50

Structural formulae Si Ti A] Cr Fe3. Fe2+ Mn Mg Ni Ca Na K F CI OH Total

5.30 total 0.01 0.14

7.20 total 0.02 0.15

Kyanite WK526

Dolomitc WK541

37.00 0.05 63.40

0.01 0.0 I 0.00

0.38 0.01 0.00

4.40 0.04 18.00

0.06 0.09 0.04

0.15 0.12 0.02

0.09 0.02 0.01

24.00 0.00 0.01

23.90 0.01 0.01

0.02 0.00 0.01

28.40 0.01 0.05

95.43 22

92.03 22

93.57 22

97.19 12.50

97.30 12.50

100.87 5

50.92 6

7.95 0.0(} 0.05 iv 0.03 vi

7.99 0.00 0.01 iv 0.09 vi

7.99 0.00 0.01 iv 0.05 vi

3.02 0.01 1.53 iv 1.20 vi

3.03 0.01 1.53 iv 1.20 vi

0.99 0.0(} 2.02

0.00 0.00 0.00

0.39 0.00 5.39

0.37 0.00 5.46

0.37 0.00 5.55

0.35 0.00 0.02

0.48 0.00 0.02

0.01 0.00 0.00

0.36 0.00 2.64

0.01 0.02 0.01

0.02 0.03 0.00

0.01 0.01 0.00

2.02 0.00 0.00

2.03 0.00 0.00

0.00 0.00 0.00

2.99 0.00 0.01

14.02

13.98

14.02

8.14

8.19

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6.00

g a r n e t - p y r o x e n e t h e r m o m e t r y , has been a s s u m e d for the p h e n g i t e barometer. W h i l e the j a d e i t e + quartz b a r o m e t e r o n l y gives m i n i m u m - p r e s s u r e estimates, the p h e n g i t e b a r o m e t e r o f M a s s o n e and S c h r e y e r (1989) gives absolute values, if p h e n g i t e is c o - e x i s t ing with kyanite and talc. Here, the p h e n g i t e b a r o m eter o f M a s s o n e and S c h r e y e r (1987) gives mini m u m - p r e s s u r e estimates. For absolute values this calibration is o n l y valid for the p r e s e n c e o f biotite and K-feldspar. G a r n e t - c l i n o p y r o x e n e t h e r m o m e t r i c data have b e e n e v a l u a t e d using the latest calibration o f A r a n o v i c h and Pattison (1995) and B e r m a n et

al. (1995). We used pairs o f o m p h a c i t e inclusions and a d j a c e n t garnet, a m p h i b o l e inclusions and adjacent garnet, and pairs o f garnet rims and o m p h a c i t e , r e s p e c t i v e l y a m p h i b o l e , and phengitic m u s c o v i t e in the matrix for g e o t h e r m o m e t r i c calculations. We ass u m e d pressures of 20 kbar f o r e c l o g i t e facies ass e m b l a g e s ( a c c o r d i n g to Holland, 1979a,b), and 10 kbar for blueschist facies a s s e m b l a g e s (according to Holland, 1979b; Z i m m e r m a n n et al., 1994). For g a r n e t - m u s c o v i t e g e o t h e r m o m e t r i c calculations different activity m o d e l s for garnet ( H o d g e s and Spear. 1982; G a n g u l y and Saxena, 1984; Hoinkes, 1986)

IE Kurz et al./Tectonophysics 285 (1998) 183-209

197

Table 5 Results of garnet-clinopyroxene geothermometric calculations according to the calibration by Berman et al. (1995) Garnet-clinopyroxene

WK527 (omphl) WK538 (core) WK5412 (cpx in) WK563 (core) WK526 rim WK538 rim WK539 rim WK540 WK541 WK5412 WK563

WK742

Temperature (°C) for 20 kbar

In Kd

562 564 528 520 508 602 540 620 598 583 558 597 589 576 570 577 569 612 531 602 623 602 577

1.9655 1.9501 2.0382 2.0918 2.3704 1.6480 1.8885 1.6054 1.7079 1.8957 1.9232 1.8362 1.6084 1.7019 1.8545 1.8110 1.8287 1.7883 1.9411 1.7515 1.6536 1.7154 1.7576

have been applied. The representative results of thermobarometric calculations are listed in Tables 5-10.

5.2. Mineral chemistry 5.2.1. Omphacite Microprobe analyses of omphacite document clear differences between omphacitel (Jd33_38) and o m p h a c i t e 2 (Jd47-50) (Table 2). In places, omphacite is zoned showing Jd34_37 in the core and Jd47-50 at its margins (Table 2) (cf. Miller, 1977; Frank et al., 1981). Omphacites that are aligned with their long axes parallel to the mylonitic foliation show these zonation patterns too (Fig. 5). Omphacite inclusions in garnet also show chemical compositions of omphacite 1 and only subordinate omphacite2 compositions near the garnet rims. 5.2.2. Amphiboles Amphiboles of the eclogite facies assemblage show barroisitic chemical compositions (Table 3).

Fig. 5. Back-scattered electron image of elongated omphacite with discontinuous concentric chemical zonation, aligned subparallel to the penetrative foliation; bright cores show lower Na20 contents than dark rims (see Table 1 for chemical composition; sample WK526).

Barroisite crystals that occur as inclusions in garnet, and barroisite that had replaced omphacite (barroisite2), show similar chemical compositions. Additional garnet inclusions are several types of taramite and alumino-hornblende.

5.2.3. Mica Phengite inclusions in garnet display similar chemical compositions as phengitel in the matrix (Si max = 3.47 per formula unit (pfu)), while phengite2 displays Si max values of 3.15 pfu (Dachs, 1986) (Table 3). Accordingly, paragonite inclusions in garnet are similar in chemical composition to paragonite within the matrix (Table 3). 5.2.4. Garnet All the investigated garnets display (more or less) roughly the same zonation patterns (Fig. 6a), but in detail many garnet grains are inhomogeneously zoned. Euhedral garnets are zoned concentrically. Mn and Ca concentrations show a maximum in the

198

W. Kurz et al./Tectonophysics 285 (1998) 183-209

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W. Kurz et al./ Tectonophysics 285 (1998) 183-209

199

Table 6 Parameters of pairs of garnet and omphacite used for garnet-clinopyroxene geothermometric calculations given in Table 5 Garnet-omphacite pairs Omphacite(struct. formulae) WK526rim

WK527

WK527

WK538 WK538 WK539 WK540 WK540 WK541 WK541

Si AI Fe Mg Ca Na X(Na)

1.993 0.499 0.120 0.410 0.472 0.506 0.491

1.984 0.372 0.100 0.651 0.612 0.372 0.356

1.982 0.372 0.100 0.560 0.615 0.371 0.354

1.991 0.498 0.180 0.342 0.411 0.563 0.495

1.999 0.553 0.130 0.343 0.396 0.579 0.556

1.992 0.522 0.140 0.370 0.422 0.450 0.515

2.011 0.482 0.150 0.387 0.472 0.524 0.497

1.994 0.446 0.180 0.404 0.459 0.519 0.446

1.991 0.487 0.120 0.413 0.481 0.503 0.479

1.998 0.484 0.140 0.399 0.462 0.519 0.482

Garnet X(Mg) X(Ca) X(Fe) Temperature (°C) (20 kbar)

0.286 0.203 0.435 602

0.334 0.199 0.425 562

0.337 0.194 0.423 564

0.136 0.205 0.580 528

0.204 0.229 0.511 540

0.259 0.210 0.488 620

0.231 0.206 0.494 598

0.177 0.221 0.525 583

0.253 0.231 0.503 558

0.224 0.259 0.493 597

Garnet-omphacite pairs WK5412 inclusion WK5412 WK5412 WK563 WK563 WK563 WK563 WK742 WK742 WK742 Omphacite Si Al Fe Mg Ca Na X(Na)

2.000 0.462 0.140 0.431 0.474 0.484 0.468

1.996 0.497 0.110 0.418 0.463 0.509 0.497

1.995 0.463 0.110 0.459 0.496 0.473 0.501

1.988 0.336 0.200 0.511 0.561 0.408 0.326

1.980 0.473 0.120 0.452 0.513 0.464 0.452

1.989 0.470 0.110 0.447 0.510 0.472 0.459

2.013 0.453 0.110 0.455 0.492 0.473 0.468

1.991 0.458 0.190 0.382 0.424 0.553 0.449

2.000 0.519 0.130 0.384 0.433 0.534 0.519

1.996 0.479 0.160 0.391 0.443 0.530 0.478

Garnet X(Mg) X(Ca) X(Fe) Temperature (°C) (20 kbar)

0.187 0.228 0.492 520

0.334 0.160 0.439 589

0.334 0.160 0.439 576

0.138 0.202 0.578 508

0.273 0.215 0.463 570

0.289 0.218 0.435 577

0.289 0.218 0.435 569

0.193 0.211 0.574 612

0.226 0.211 0.533 531

0.226 0.211 0.533 603

core, with a s i m u l t a n e o u s m i n i m u m o f M g (Fig. 6a). W i t h i n the inner zones, M n and C a c o n t e n t s d e c r e a s e in a b e l l - s h a p e d curve; C a c o n t e n t s then slightly increase to the rims, w h i l e M g c o n t e n t s c o n t i n u o u s l y increase towards the rims. Fe b e h a v e s antipathetic to Mn. T h e s e e l e m e n t s d o c u m e n t an increase f r o m the core to the rims f o l l o w e d by two m a x i m a . N e a r the garnet r i m s Fe contents decrease. T h e i n h o m o g e n e o u s garnets are c h a r a c t e r i z e d by irregular radial patterns near the core, w h i c h do not pass through to the rims o f the grain. T h e e l e m e n t c o n c e n t r a t i o n within these z o n e s is v e r y similar to the e l e m e n t c o n c e n t r a t i o n along the r i m s o f the garnet grain (Fig. 6b). In detail, this results in a h i g h l y irregular z o n a t i o n (Fig. 6). This is also v e r y w e l l d o c u m e n t e d by X - r a y m a p s o f single garnet crystals (Fig. 6b).

T h e cracks p r e d o m i n a n t l y o c c u r in the garnet cores and do not penetrate to the garnet rims. T h e y postdate the garnet inclusions. T h e s e irregular radial patterns m a y r e p r e s e n t f o r m e r cracks w h i c h have b e e n a n n e a l e d and sealed during the final stages o f garnet growth. 5.3. T h e r m o b a r o m e t r i c c a l c u l a t i o n s

In the E c l o g i t e Z o n e o f the Tauern W i n d o w early inclusions in garnet (quartz, chlorite, biotite, e p i d o t e / z o i s i t e ) indicate a first m e t a m o r p h i c e v e n t at g r e e n s c h i s t to blueschist facies conditions (Miller, 1977). O m p h a c i t e l crystals (Jd33-38) (sample W K 5 2 7 ) f o r m e d during the s u b s e q u e n t first stage o f H P e c l o g i t e facies m e t a m o r p h i s m at

200

IV.. Kurz et a l . / Tectonophysics 285 (1998) 183-209

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Fig. 7. Temperature-pressure-time path including several phases of deformation for the Eclogite Zone, based on data of our ow investigation and P-T-paths published by Zimmermann et al. (1994).

minimum pressures of ca. 11-12 kbar according to the experimental calibrations of Holland (1980, 1983). Based on the paragenesis garnet-omphacitekyanite-paragonite-phengite temperatures of 532 + 38°C and pressures of 20.2 4- 1 kbar (Table 10) were reached during the formation of omphacite 1 (Fig. 7). These conditions have been evaluated from the following independent end-member reactions: 3 muscovite -1- pyrope .1. 4 quartz = 3 celodonite .1. 4 kyanite

3 muscovite + 6 diopside = 3 celadonite + 2 grossular + pyrope 3 celadonite + 6 cats = 3 muscovite -t- 2 grossular + pyrope 4 paragonite + 3 diopside = pyrope + 4jadeite + 3 cats + 4 quartz + 4 H20 3 hedenbergite .1. 3 cats = 2 grossular + almandine

201

W. Kurz et a l . / Tectonophysics 285 (1998) 183-209

Table 7 Results of garnet-amphibole geothermometric calculations for eclogite (E) and blueschist (B) facies assemblages

j a d e i t e component (Jd47-50). This documents slightly higher m i n i m u m pressure conditions of 14-15 kbar than documented by o m p h a c i t e l . The g a r n e t clinopyroxene (omphacite2), g a r n e t - a m p h i b o l e (barroisite2), and g a r n e t - m u s c o v i t e thermometric evaluations (Tables 5 - 8 ) indicate slightly but significantly higher temperatures (570-630°C) during the formation o f omphacite2 (Fig. 7). Barometric calculations gave somewhat higher pressure conditions ( 1 9 - 2 3 kbar) during the formation of omphacite2 recrystallized grains. Si contents of synkinematic p h e n g i t e l crystals, that are part o f the p e a k pressure assemblage (Si max = 3.47 pfu), document m i n i m u m pressures o f ca. 10-13 kbar, according to the calibrations by Massone and Schreyer (1987) (Dachs, 1990). Pressures o f ca. 2 0 - 2 2 kbar (according to calibrations by Massone and Schreyer, 1989) are evaluated for phengite that is coexisting with talc and kyanite (Table 9; Fig. 7). Subsequent blueschist facies conditions are only documented by the replacement o f omphacite by barroisite2. Miller (1974) described oriented, single crystal replacement o f o m p h a c i t e l by subcalcic amphibole. The P - T conditions are less constrained. G a r n e t - a m p h i b o l e (barroisite2 along margins of garnet) thermometric evaluations gave ca. 475-535°C according to calibrations o f G r a h a m and Powell (1984) and Perchuk et al. (1985) (Table 6).

Garnet-amphibole Temperature (°C)

In Kd

Graham and Powell (1984)/Perchuk et al. (1985) WK538 inclusion; B WK539; B WK540 inclusion; E WK540 inclusion; E WK540 inclusion; E WK541 inclusion; B WK541; E WK541; B WK742; E WK742; E WK742; E

384/309 532/474 639/589 594/545 613/560 534/454 607/535 542/475 606/558 655/606 657/620

3.384 2.124 1.529 1.732 1.659 2.243 1.784 2.117 1.675 1.456 1.393

(albite = j a d e i t e + quartz) (paragonite = j a d e i t e + kyanite + H 2 0 ) The temperature estimates are in accordance with g a r n e t - c l i n o p y r o x e n e ( o m p h a c i t e l ) (530-560°C) (Tables 5 and 6) and g a r n e t - a m p h i b o l e in garnet (545°C) (Table 7) thermometric evaluations. Omphacite2, that formed due to dynamic recrystallization o f o m p h a c i t e l , shows a significantly higher Table 8 Results of garnet-muscovite geothermometric calculations Temperature (°C) In Kd (Hynes and Forest, 1988) WK539

510 552 547 551 WK540 530 499 526 539 543 WK541 539 WK742 628 607 584 624 563

1.992 1.729 1.757 1.729 1.863 2.069 1.891 1.811 1.783 1.806 1.310 1.416 1.541 1.328 1.659

Temperature (°C) (Hodges and Spear, 1982)

In Kd

Temperature (°C) In Kd (Ganguly and Saxena, 1984)

Temperature (°C) In Kd (Hoinkes, 1986)

Temperature (pressure correction)

515 556 551 556 535 505 531 544 548 543 632 612 589 629 568

2.069 1.801 1.829 1.801 1.931 2.141 1.959 1.878 1.849 1.883 1.367 1.441 1.567 1.385 1.723

518 549 546 549 532 508 529 539 542 540 606 591 573 608 559

541 579 575 580 560 531 556 568 572 568

1.752 1.485 1.513 1.485 1.619 1.827 1.648 1.567 1.539 1.561

583 624 620 624 603 573 599 611 616 612

611

1.290

592

1.408

2.022 1.759 1.787 1.759 1.896 2.103 1.924 1.844 1.815 1.830 1.335 1.465 1.567 1.353 1.685

202

W. Kurz et al./Tectonophysics 285 (1998) 183-209

Table 9 Results of phengite geobarometric calculations Pressure (kbar) for 600°C Massone and Schreyer(1987) WK526 WK527 WK538 WK539 WK540 WK541 WK742

Massone and Schreyer (1989)

10.75 12.00 11.75 11.00

21.87

12.75 10.00 9.75 10.75

20.00 19.54 21.40 21.87 22.33 22.80

11.00

11.25 11.50

The conditions of subsequent lower amphibolite to upper greenschist facies metamorphism are extensively documented by, for example, Franz and Spear (1983), Dachs (1986, 1990), Spear and Franz (1986), Frank et al. (1987a). 6. D i s c u s s i o n

The textural and microfabric evolution of eclogites combined with P - T data allow the interpretation of the eclogitic and post-eclogitic deformation histories from burial by subduction to subsequent exhumation. Metasedimentary rocks, which are very often associated with eclogites, are more or less

completely overprinted during metamorphism and deformation subsequent to their burial history. For the exhumation history of eclogites and eclogite facies rocks of the central-southern Tauern Window many controversial models have been developed. According to Platt (1986), Behrmann and Ratschbacher (1989) assumed that top-to-the-N to -NE plate convergence is followed by E - W extension. They established the model of an overthickened accretionary complex which collapsed, leading to the exhumation of the HP metamorphic assemblages. This model implies, because of recorded pressures of at least 20 kbar in the eclogites and surrounding metasediments, an accretionary wedge of at least 70 km thickness. In this model the pressure peak would have been followed by rapid heating during decompression shortly after the collapse, which is not evidenced by published data and our own P - T data (Fig. 7). The fact that eclogite facies metamorphism is overprinted by blueschist facies (HP/LT) metamorphism and the long-term preservation of high-pressure/low-temperature assemblages requires a mechanism of continued refrigeration by a cold subducted lithospheric slab (Ernst and Dal Piaz, 1978; Rubie, 1984; Ernst, 1988), in order to avoid re-equilibration during prolonged periods of static heating during exhumation and tectonic uplift. England and Holland (1979) argued for buoyant uprise of eclogites incorporated in low-density matrix rocks of the accretionary wedge. This model argues for steeply dipping faults bounding the buoyant eclogite

Table 10 Results of geothermometric and geobarometric calculations THERMOCALC WK526 WK527 only omphl WK538 WK539 (average pressure) WK539 WK541 WK561 WK563 WK563/2

Temperature (°C)

s.d. temp

Pressure (kbar)

s.d. pressure

fit

cor.

636 532 544 550 575 600 625 573 528 593 575 592

54 38 49

20.5 20.2 22.9 22.8 23.2 23.6 23.9 21.3 20.7 23.7 22.5 15.6

1.30 1.00 1.10 2.53 2.93 3.37 3.84 2.4 1.50 6.2 3.9 6.1

0.93 0.42 0.71 1.10 1.30 1.50 1.80 2.28 1.22 3.97 2.65 1.18

0.813 -0.628 -0.849

92 58 73 47 36

--0.787 0.783 0.085 0.074 0.704

w. Kurz et al./Tectonophysics 285 (1998) 183-209

slices, but no clear evidence for such an eclogite emplacement exists. Genser et al. (1996) modelled the metamorphic evolution of the Penninic units within the Tauern Window showing the need of a subduction mechanism to explain the shift from eclogite to subsequent blueschist facies conditions. Our investigations have shown that the eclogites of the southern-central Tauern Window document as well the final increment of deformation and P - T evolution along the prograde P - T path, as the subsequent exhumation history along the retrograde section of the P - T path. Based on inclusions in zoned garnets, Miller (1977) has documented increasing P T conditions during garnet growth. A first generation of inclusions documents greenschist to blueschist facies conditions, while subsequent inclusions of omphacite crystals show the same chemical composition as omphacitel. Ductile deformation of eclogites is indicated by typical mylonitic features, like the formation of subgrains in omphacitel, and the formation of omphacite2 due to subgrain rotation recrystallization (White, 1977; Gottstein and Mecking, 1985). These microstructures are typical indicators of dislocation creep of omphacite (e.g. Buatier et al., 1991; Philippot and Van Roermund, 1992). The formation of the penetrative foliation S 1 was already beginning during growth of omphacitel along the prograde P - T path (Fig. 7), but has continued during the formation of omphacite2 at the pressure and temperature peak. This is indicated by the zonation patterns of omphacite crystals that are aligned parallel to the mylonitic foliation. We interpret the formation of these fabrics to be related to the subduction of this unit, which is documented by increasing minimum pressures, resp. increasing jadeite contents of omphacite, during the prograde evolution, constant or slightly increasing absolute pressures during the formation of omphacitel and omphacite2, and a slight temperature increase (Fig. 7). The prograde temperature evolution is documented by the chemical zonation profiles of garnet, too (Mg increasing, Fe decreasing from the core to the rim). The formation of omphacite2 at peak P - T conditions indicates the cessation of the subduction of the Eclogite Zone at depths of 60-70 km within the subduction zone. From the well preserved zonation patterns of omphacite we conclude that the homogenisation of omphacite was prevented by volume diffusion, either

203

due to a high subduction velocity or due to low temperatures. Omphacite is assumed as the controlling mineral of the rheology of the oceanic crust at depths greater than 40 km in a subduction zone. As data on the experimental deformation behaviour of omphacite do not exist, diopside is very often used as the rheological equivalent for wet omphacite (Philippot and Van Roermund, 1992; Mancktelow, 1993). It was documented by Philippot and Van Roermund (1992) that a sharp mechanical discontinuity from frictional sliding to dislocation creep is developed at 44-48 km depth, assuming a reasonable geothermal gradient of 12°C km -l within subduction zones, and a shear strain rate of l 0 -12 s - l . At this depth pressures of 12-14 kbar are reached. This would imply that dislocation creep and subgrain rotation recrystallization of omphacite, respectively clinopyroxene (Champness et al., 1974; Avb Lallemant, 1978; van Roermund and Lardeaux, 1991; Godard and van Roermund, 1995), started at these structural levels. This is consistent with our geobarometric results. The oriented single crystal replacement of omphacite by subcalcic barroisitic amphibole indicates that deformation was prolongued along the cooling path within the same kinematic boundary conditions. This might indicate the emplacement of the eclogites onto basement units and contact with cool lithosphere. High-pressure assemblages within cracks and fissures (Thomas and Franz, 1989) indicate locally elevated fluid pressures during the emplacement of these eclogites. The conditions of blueschist facies metamorphism within the Eclogite Zone are difficult to evaluate because of lack of appropriate parageneses due to subsequent overprint of greenschist facies metamorphism. Pressures of 10-15 kbar at 300 °450°C are reported by Holland (1979a,b), and 10-15 kbar at 400°-450°C by Zimmermann et al. (1994) for the blueschist facies metamorphism within the Eclogite Zone. Blueschist facies metamorphism is documented within the tectonic hanging-wall of the Eclogite Zone, too, for example by pseudomorphs after lawsonite (Fry, 1973; H6ck, 1974; Dachs, 1990), and relics of glaucophane and jadeitic pyroxene (Holland and Ray, 1985); temperatures of 400°-450°C and minimum pressures of 8 kbar are estimated for this stage of blueschist facies metamorphism within the Glockner Nappe. Phengite chemical compositions indicate minimum pressures of 10 kbar within the Ma-

204

W. Kurz et al./Tectonophysics 285 (1998) 183-209

trei Zone (Dachs, 1990). Within the foot-wall units (Venediger Nappe; Storz-Riffi Nappe) pressures of 10-11 kbar are reported by Selverstone et al. (1984), and Selverstone (1993). These geobarometric data indicate a 4-10 kbar pressure-gap between the peak pressures in the Eclogite Zone (20 kbar) and the foot-wall units (10 kbar). This would imply the excision of a ca. 30-35-km-thick rock section, or the emplacement of the Eclogite Zone by thrusting. The recorded P - T path suggests that the juxtaposition of these units was achieved in eclogite to blueschist facies conditions, while heating was considerably delayed along a colder P - T path. Such conditions are only known within active subduction zones. Therefore, we suggest that the emplacement of the Eclogite Zone was achieved along a S-dipping thrust along the base of this unit (Fig. 2b), by top-to-the-N transport. The cofacial nature of the HP assemblages in metasediments and metabasic rocks also suggests that the Eclogite Zone behaved as a coherent unit during eclogite facies metamorphism and its emplacement (Spear and Franz, 1986; Droop et al., 1990). This is supported by the apparent stratigraphic continuity of continental margin sequences on a scale of tens of metres to kilometres (Raith et al., 1980; Kurz et al., 1996). According to Hynes et al. (1996), eclogire emplacement has resulted from partial subduction of the continental margin, which, because of its high flexural rigidity, produced a rapid change in the trajectory of the descending slab. Therefore, the wedges overlying the subduction zone, and the subducting slab itself, experience substantial extra stresses that are adequate to effect significant deformation of the wedge and the subducting plate, and the detachment of the eclogite bearing unit (Fig. 8a). Top-to-the-N to top-to-the-NE sense of shear is recorded within the foot-wall and the hanging-wall units, too. Within these units, this deformation (Dl) is related to the subsequent formation and emplacement of the Penninic nappe imbricate stack in the hanging-wall of the Eclogite Zone (Kurz et al., 1996). Alternatively, the P - T path of the eclogite facies rocks is compatible with an emplacement model according to corner flow (Cloos, 1982, 1985; Shreve and Cloos, 1987) (Fig. 8b). The driving force is the movement of the descending lithospheric plate. When the flow is forced to occur in a low-angle corner, forced convection occurs, which can carry blocks of blueschists and eclog-

ites back to greenschist facies conditions. However, the P - T path might be influenced by various parameters such as subduction velocity, internal heat production, and shear stress. Following Cloos (1982), flow needs not always to produce a chaotic melange. For example, a higher thermal regime may lower the ductility contrast between different rock types. In this case, the original layering may not be lost and upwelled material may appear as a structurally complex zone characterized by isoclinal folding, as is the case in the Eclogite Zone, too (Fig. 2). This mechanism of emplacement might explain the locally irregular orientation distributions of eclogite facies stretching lineations (Fig. 2a), which, however, can be explained to result also from subsequent deformation too. The Eclogite Zone was overthrust subsequently by the Penninic nappe imbricate stack. Garnets seem to have been broken during this deformational phase at cooler conditions, as documented by the inhomogeneous radial element distribution patterns. We interpret these patterns as cracks which have been healed during subsequent amphibolite facies conditions at elevated temperatures. Later top-to-the-W (D2) shear is related to decompression along the exhumation path, and furthermore, documents a change of the kinematic boundary conditions. It developed at or slightly prior to upper greenschist to amphibolite facies metamorphic overprint (ca. 550°C, 6-7 kbar) (e.g., Dachs, 1990) of the Eclogite Zone. The foot-wall and hanging-wall units were affected by this deformation and metamorphism, too (Kurz et al., 1996). Epidote-amphibolite facies conditions were reached due to thermal equilibration subsequent to the subduction and accretion of continental lithosphere. This resulted in heating subsequent to blueschist facies metamorphism (Fig. 7). Most of the garnets do not reflect this development, and re-equilibration with the amphibolite to greenschist assemblages is limited, as documented by Dachs (1990). Only the garnet edges and annealed cracks reflect the amphibolite to greenschist facies evolution. However, amphibolite to greenschist facies metamorphic overprint is limited in the garnets and the garnets seem to be in disequilibrium with the amphibolite to greenschist assemblages. Spear and Franz (1986) documented final rock alterations at 200-350°C and pressures <3 kbar,

W. Kurz et al./Tectonophysics 285 (1998) 183-209

300 0

(km)

200

205

100

stress ctories

100

(km)

a

200

b Fig. 8. Models for eclogite emplacement within the Tauern Window. (a) Wedges overlying the subduction zone, and the subducting continental margin, experienced substantial extra stresses that are adequate to effect significant deformation of the wedge and the subducting plate (after Hynes et al., 1996). (b) Model documenting the flow paths of the corner flow model (following Cloos, 1982, 1985) and the emplacementof eclogite facies rocks onto basement units (modifiedfrom Spear, 1993. which might coincide with the formation the documented cataclastic shear zones. Semiductile and brittle normal faults which are documented from the western and eastern margins of the Tauern Window (Selverstone, 1988; Genser and Neubauer, 1989; Selverstone et al., 1995), and which have also been detected by our own investigations in the structural depression of the central Tauern Window, transect

the whole structural sequence and are related to the doming of the basement units within the Tauern Window.

7. Conclusions (1) The structural record of the eclogites within the Tauern Window is related to the contraction his-

W. Kur= et al./Tectonophysics 285 (1998) 183-209

206

tory. The early structures, like eclogite mylonites, document a contractional regime and vertical crustal thickening possibly related to a large-scale subduction zone. The P - T - d data document the final increment of the prograde evolution of the eclogites. (2) Emplacement of the eclogites onto basement units most probably was achieved by thrusting while heating was considerably delayed along a cold P-T path. An alternative interpretation is an emplacement similar to a corner flow model. (3) HP eclogite facies metamorphism occurred before emplacement of the eclogite facies rocks onto the foot-wall tectonic units; emplacement occurred possibly at blueschist facies conditions. (4) The youngest greenschist facies structures document deformation during decompression, subsequent to the cessation of subduction.

WK538

WK539

WK540

WK541

WK541/2

WK557

WK561 WK563

Acknowledgements WK742

We appreciate discussions with Johann Genser, Robert Handler, Gerhard Amann and with HansPeter Steyrer and their company during many wild and dangerous field trips. Further thanks to Robert Sturm for offering his programs to evaluate mineral structural formulae. Thanks to Lothar Ratschbacher who gave access to some unpublished geochronological data from the Eclogite Zone and for many exchanges of ideas. We appreciate the help during microprobe analytical work by Dan Topa, and Volker H6ck for running the microprobe laboratory. We appreciate the formal reviews by Karel Schulmann and two unknown colleagues. The study was funded by a grant (P9918-GEO) of the Austrian Research Foundation. Appendix A. Locations of eclogite mylonite samples "r" and "h" are the coordinates of the sample locations, based on the topographic map of Austria (OEK 1;50,000) [r = "rechts" (right); h = "hoch" (up)]. WK526

WK527

OEK 1 : 50,000, sheet 152 (Matrei), Dorfertal, Venediger H6henweg, ca. 750 m SE of Johanneshtitte, 2 375 m; r 374,600, h 213,500. OEK 1 : 50,000, sheet 152 (Matrei), Dorfertal, Venediger Hghenweg, ca. 500 m SE of Johanneshtitte, 2250 m; r 374,500, h 213,850.

FROI

FRO2

OEK 1:50,000, sheet 152 (Matrei), northwestern shore of Eissee, 2680 m, Timmeltal; r 377,970, h 214,450. OEK 1:50,000, sheet 152 (Matrei), northwestern shore of Eissee, 2700 m, Timmeltal; r 377,950, h 214,520. OEK 1:50,000, sheet 152 (Matrei), western shore of Eissee, 2670 m, Timmeltal: r 377,920, h 214,410. OEK 1 : 50,000, sheet 152 (Matrei); Timmeltal, 1900 m SW of Wei6 Spitze, 2500 m; r 377,430, h 214,200. OEK 1:50,000, sheet 152 (Matrei); Timmeltal, 1900 m SW of Weig Spitze, 2500 m; r 377,480, h 214,200. OEK 1:50,000, sheet 152 (Matrei); Frosnitztal, 250 m north of Steinsteig, 2050 m; r 383,050, h 215,500. OEK 1:50,000, sheet 152 (Matrei); Frosnitztal, Steinsteig, 2072 m; r 383, 080, h 215,240. OEK I : 50,000, sheet 152 (Matrei); Frosnitztal, 100 m southeast of Steinsteig, 2055 m; r 383,125, h 215,150. OEK 1:50,000, sheet 152 (Matrei); 350 m west of point 2273 (Raneburger See); re 386,740, h 215,625. OEK 1 : 50,000, sheet 152 (Matrei); Frosnitztal, 1000 m SSE of Steinsteig, 1920 m; r 383,380, h 214,170. OEK 1 : 50,000, sheet 152 (Matrei); Frosnitztal, 150 m west of Steinsteig, 2070 m; r 382,850, h 215,280.

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