formation of alpine mylonites and pseudotachylytes at the base of the silvretta nappe, eastern alps

Tecfonophysics, 204 (1992) 289-306

Elsevier Science Publishers B.V., Amsterdam

289

Formation of Alpine mylonites and pseudotachylytes at the base of the Silvretta nappe, Eastern Alps

Norbert Koch and Ludwig Masch

Mineralogisch-Petrographisches Institut a’er Universitiit Miinchen, Theresienstrasse 41, D-8000 Munich 2, Germany

(Received January 9,199l; accepted March 28,199l)

ABSTRACT

Koch, N. and Masch, L., 1992. Formation of Alpine mylonites and pseudotachylytes at the base of the Silvretta nappe, Eastern Alps. In: J.F. Magloughlin and J.G. Spray (Editors), Frictional Melting Processes and Products in Geological Materials. Tectonophysics, 204(spec. sect.): 289-306.

Optical microscopy and transmission electron microscopy (TEM) have been used to investigate deformation-induced microstructures in Alpine pseudotachylytes, mylonites and ultramylonites from the Upper Austro-Alpine Silvretta nappe at the southwestern margin of the Lower Engadine Window in the Eastern Alps of Switzerland and Austria. The earliest Alpine stage of deformation (Dl) produced mylonitic gneisses, with dislocation glide being the dominant deformation mechanism in quartz. In the second stage (D2a), mylonites were formed as a result of strain softening in quartz (manifest by dynamic recrystallization and a lattice preferred orientation). TEM observations are used to characterize the initial stages of migration and rotation recrystallization. Pseudotachylytes were generated during Dzb when strain hardening due to dislocation tangling was not outweighed by strain softening. Several generations of mylonites and pseudotachylytes are observed. A model for the propagation of seismic fractures, initiated in comparatively rigid amphibolites, into an otherwise ductile environment is proposed for the pseudotachylytes. As such they are formed as transient discontinuities in the mylonites. D, and Dla,b are related to the Eo-Alpine detachment of the Silvretta nappe. Subsequent Eo-Alpine metamorphism in the investigated area reached the stilpnomelane zone, and this coincided with the development of ultramylonites from both the mylonites and pseudotachylytes during D, and peak metamorphism. The increased temperature combined with fluid infiltration from the overridden Penninic strata favoured grain-boundary sliding (superplasticity) as the dominant deformation mechanism during Ds.

Introduction

Mylonites and pseudotachylytes are two different types of fault rock. While mylonites are the products of ductile deformation, pseudotachylytes are formed by brittle failure. Following the study of Sibson (19771, it is now widely accepted that pseudotachylytes are generated above and within the brittle-ductile transition zone at shallow levels of the crust down to about 10 km (Passchier, 1982; Maddock, 1986; Scholz, 1988), whereas the

Correspondence to: N. Koch, Mineralogisch-Petrographisches Institut der Universitlt Miinchen, Theresienstrasse 41, D-8008 Munich 2, Germany.

formation of mylonites extends to greater depths. The comparatively low ductile strength of quartz marks the beginning of the brittle-ductile transition zone, not only for quartzites but also for polymineralic quartz-bearing rocks (Scholz, 1988). The occurrence of pseudotachylytes within an otherwise ductile environment has been reported by, for example, Sibson (1980) in the Outer He- brides Thrust zone. The occurrence of pseudotachylytes at the base of the Silvretta nappe within the Eastern Alps of Switzerland and Austria was initially reported by Bearth (1933).

It is a widely accepted hypothesis that pseudotachylytes are formed by the fusion of wall rocks as a result of frictional melting during seismic failure (e.g., Maddock, 19861, although Wenk

0040-1951/92/$05.00 0 1992 - Elsevier Science Publishers B.V. All rights reserved

290 N.KOCHANDI.. MASC‘H

(1978) proposes an origin by extreme cataclasis. Aspects of the occurrence and origin of pseudotachylytes from the Sllvretta nappe have been studied in detail by Masch (1970, 1974, 1979). Evidence in favour of an origin by frictional melting includes preserved skeletal microlites of plagioclase and hornblende, spherulitic biotite and plagioclase embedded in a cryptocrystalline matrix and “finger-print” textures suggesting the partial melting of plagioclase relicts.

This study involved detailed field work and investigations of the microstructures exhibited by mylonites, pseudotachylytes and their host rocks, as well as extensive TEM work on dislocation sub-structures in quartz. The aim was to establish the sequence of formation for the pseudotachylytes and mylonites. From an understanding of the deformation-induced microstructures it has been possible to deduce the prevailing deformation mechanisms and the rheological behaviour of the rocks during certain stages of nappe movement.

Methods

Over 150 hand specimens (mostly oriented) from 26 localities were sampled. In this study, the plastically deformed gneisses are divided into mylonitic gneisses, mylonites and ultramylonites ac- cording to their proportion of deformed and/or recrystallized grains: < 33% for the mylonitic gneisses; 33-90% for the mylonites and > 90% for the ultramylonites. All transitions between mylonitic gneisses and ultramylonites can be observed. 56 TEM foils were prepared from se- lected thin sections by thinning on a dual-gun Technics MIM-IV ion mill. Care was taken that holes in the foils were produced at the sites of prominent optically visible microstructures (such as deformation lamellae, rims and cores of relict grains, serrated grain boundaries and recrystallized grains). TEM observations were made using JEOL JEM-1OOB and JEM-200A electron micro- scopes operated at 100 kV and 200 kV, respec- tively. Dislocation densities were measured using the method outlined by Ham (1961). Albite in the ultramylonitic matrix was identified using a Cameca SX 50 microprobe.

Regional geology

The internal structure of the Silvretta nappe which, in the investigated area, consists of amphibolites and ortho- and para-gneisses (Streckeisen, 1928; Bearth, 1932; Spaenhauer, 1932; Wenk, 1934) is the result of Palaeozoic and probably Variscan deformation (Schlingentektonik) under amphibolite facies conditions (Wenk, 1934; Grauert, 1966; Amann, 1985; Magetti and Galetti, 1988).

Eo-Alpine nappe movements have detached the Silvretta from its basement in the Southern Alps and, after north- to northwestward transport, have emplaced it over Penninic strata (Laubscher, 1983). The latter consist of calcare- ous flysch with intercalated ophiolites which are now exposed within the Lower Engadine Window (Fig. 1).

All gneisses located at the base of the Silvretta nappe (which comprise augen gneisses, biotite- plagioclase gneisses and garnet-staurolite gneisses) show a mylonitic overprint of varying intensity. This has produced mylonitic gneisses, mylonites and ultramylonites, with foliations oriented predominantly subparallel to the Variscan foliation in the investigated area. Where two sets of foliations are discernible, discordant type I and II S-C mylonites can be observed. Amphibolites show no evidence of plastic deformation on the mesoscopic scale, but are pervasively fractured (Masch, 1974).

The pseudotachylytes form discordant systems of veins (Masch, 1974; Schmutz, 1982) and Rb-Sr isotopic age determinations indicate that they are of Cretaceous age (Thiini, 1987). No cataclasites were seen to accompany the pseudotachylytes (cf. Macaudikre et al., 1985; Magloughlin, 1992).

The area1 distribution of pseudotachylytes and mylonites is shown in Figure 1. The pseudotachylytes are restricted to the Silvretta nappe and are most commonly observed between Val Tuoi and the Fimbertal, but can be found all along the southern and western margins of the Lower En- gadine Window for a distance ranging from a few metres to approximately 300 m above the thrust plane (Masch, 1974). Though generally more abundant in amphibolites than in gneisses, they

FORMATION OF ALPINE MYLONITES AND PSEUDOTACHYLYTES 291

5

km

A

Abundant/sporadii pseudotachylytes

Abundant/sporadic mybnites

Scarl nappe

dtztal napps

Sesvenna nappe

Siivretta nappe

Penninic nappss

Thrust fault

Mountain peak

Fig. 1. Simplified geology of the study area and its regional location within the Eastern Alps. The distribution of pseudotachylytes and mylonites is shown by the density of diagonally ruled lines. Pseudotachylytes are most abundant between the Fimbertal and VaI Tuoi. Mylonites are widespread south of the River Inn. Compiled from Bearth (19331, Cadisch et al. (19411, Masch (19741, Schmutz

(1982) and field work during this study. Tectonic units after Triimpy (1980).

occur in all types of rock within the Silvretta crystalline complex. No material from the Pen- ninic footwall has been incorporated into the pseudotachylytes.

Metamorphic grade

Detailed investigations have established the influence of Alpine metamorphism on the amphibolites and gneisses as well as in the mylonitic suite and pseudotachylytes. Alpine metamorphism of the stilpnomelane zone is documented by the growth of stilpnomelane in coexistence with chlorite-muscovite-K-feldspar-albite- quartz in gneisses. The paragenesis clino- zoisite/epidote-chlorite-albite-actinolite is sta- ble in the amphibolites. These assemblages par- tially replace the pre-Alpine amphibolite facies

parageneses: biotite-muscovite-K-feldspar- plagioclase-garnet-staurolite-aluminosilicate- quartz developed in the gneisses and plagioclase- hornblende-garnet-biotite in the amphibolites. The appearance of stilpnomelane and/or rarely zoisite in the fine-grained matrix of the Silvretta pseudotachylytes, as well as in enclosed rock fragments, is unequivocal evidence for an overprint by Alpine metamorphism belonging to the regional stilpnomelane zone.

The Alpine parageneses indicate an approximate upper temperature limit of 300400°C (Winkler, 19791, with the earlier results of Niggli (1970) favouring the lower temperature. Hurford et al. (1989) have compiled a regional distribution of maximum temperatures for Eo-Alpine metamorphism based on the closure temperatures of various minerals used for age determinations. Ac-

292 N. KOCH AND I. MASCH

cording to them the maximum temperature was about 300°C at the southern margin of the Lower Engadine Window. Only to the north, where Thijni (1982) and Amann (1985) report the iso- lated occurrence of Alpine chloritoid, did the temperatures reach approximately 400°C. In the samples investigated in this study no chloritoid was observed.

Optical microstructures

Amphibolites

Amphibolites make up the main body of the Silvretta nappe complex in the investigated area. They possess variable modal compositions and are commonly layered. Where present, quartz grains show increasing elongation with sub-grains oriented parallel to the c-axis as the basal thrust plane is approached. Recrystallization to a grain size of a few microns is only observed along heavily serrated grain boundaries. The development of sub-basal deformation lamellae increases with proximity to the pseudotachylyte veins. Hornblende and plagioclase are increasingly fractured: the latter being intensely altered by Alpine saussuritization. Biotite, though scarce, shows kink band development and alteration to chlorite.

Gneisses

Those gneisses showing no influence of Alpine deformation exhibit a widely spaced (cm) foliation emphasized by muscovite and biotite. Kink bands in micas are absent. Quartz grains are only weakly elongated with moderate undulose extinction and well developed sub-grain boundaries. Grain boundaries are not serrated and deformation lamellae were not observed. Alkali feldspars with perthitic intergrowths and plagioclase with albite twinning show neither ductile nor brittle deformation effects.

Mylonitic gneisses

The texture of the parent rock is still recognizable in hand specimen, but microshear zones, inclined either 30” to 40” or parallel to the main

Fig. 2. (a) Lattice-preferred orientation of c-axes in quartz ribbons from mylonite; n = 163; (b) contoured at 0.5, 2.5, 4.5

and 6.5% intervals.

foliation, are developed in certain parts of the mylonitic gneisses. The size of relict quartz grains ranges from 100 pm to 1 mm. They show intense undulose extinction (misorientation across one grain can exceed 1001, while sub-grains (parallel to prism planes) are not well developed. With increasing strain, quartz deforms to ribbon-shaped grains that show a type I cross girdle lattice- preferred orientation (Fig. 2). The grain boundaries of the ribbons are strongly serrated and it is here that the initiation of recrystallization can be observed (Fig. 3a) as well as along mostly sub- basal deformation lamellae. The size of the newly

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294 N. KOCH AND I. MAS(‘H

formed recrystallized grains is in the order of a few microns. Extremely fine-grained recrystallization (grain size l-5 pm) also occurs in the direct vicinity of crosscutting veins of pseudotachylyte.

Plagioclase is saussuritized and fractured, the cracks being filled mostly with quartz. Alkali feldspar is also fractured. Biotite and muscovite are sharply kinked with, in biotite, incipient chlo- ritization occurring along kink band boundaries. Stilpnomelane is typically found growing in small needles on muscovite or in quartz and alkali feldspar. In the mylonitic gneisses the stilpnomelane always occurs undeformed.

Mylonites

In contrast to the mylonitic gneisses a strong reduction in grain size can be observed in hand specimens of the mylonites. Almost all quartz grains have deformed to ribbons. Recrystalliza- tion starting at serrated grain boundaries has partly obliterated whole grains. The grain size of recrystallized grains is about 20 pm. Feldspar grains are progressively reduced in grain size by fracturing. Kinking in muscovite is intense and fine-grained recrystallized layers initiate at fish- shaped grains and flow around relict feldspar grains. Biotite is almost completely replaced by chlorite. Stilpnomelane is present in only a few samples. Though it is mostly undeformed in the mylonites, it grows syntectonically in certain mylonites showing very high strain effects. These mylonites are transitional with the ultramylonites.

Pseudotachylytes

Stilpnomelane flakes growing in the pseudotachylyte matrix, as well as in rock fragments enclosed in pseudotachylyte, are a clear indicator for an Alpine metamorphic overprint that partly replaces the original quenched assemblage. A detailed study of the matrix of the pseudotachylytes is found in Masch (1970, 1974) and an example of this is shown in Figure 3b. Stilpnome- lane is mostly undeformed (Fig. 3c), but on rare occasions it is found oriented parallel to the preferred orientation of small micas in the matrix (Fig. 3d).

Fig. 4. Lattice-preferred orientation data for quartz from mylonitic gneiss and rock fragments in pseudotachylyte. The angle between the girdle orientation of quartz c-axis and foliation remains the same in both host rock and enclosed

rotated rock fragments within the pseudotachylyte.

Rock fragments in the pseudotachylytes are weakly rounded and show microstructures identical to those of the neighbouring host rock. Where pseudotachylytes cut through mylonites showing a lattice preferred orientation (LPO) of quartz, the LPO of the quartz c-axis has been rotated ac- cording to the rotation history of the fragments (Fig. 4).

Masch (1974) has reported at least three generations of pseudotachylytes in the gneissose host rocks. The younger pseudotachylyte veins cut and displace older veins with almost no ductile deformation on the mesoscopic scale. Optical microscopy reveals the extremely fine-grained recrystallization of quartz directly adjacent to pseudotachylyte veins. This study has shown that some pseudotachylytes are strongly sheared and have transformed to ultramylonite bands (Figs. 5a, b). It is important to note that where pseudotachylyte occurs within mylonite, it cuts the mylonitic foliation, and is itself further deformed to ultramylonite (Fig. 5~).

Ultramylonites

Two types of ultramylonites can be distinguished: (1) type A ultramylonites that develop from mylonites by the continued reduction of grain size due to the complete recrystallization of quartz and muscovite (Fig. 5d-f), the different phases impeding each other in grain growth. Be-


(4

Fig. 5. (a) Sketch showing pscudotachylyte veins sheared to ultramylonite bands. Outlined box shows approximate location of micrograph in Fig. 5b. Scale bar 10 cm. (b) Micrograph of pseudotachylyte deformed to ultramylonite. Note the deformed rock

fragments in pseudotachylyte. The matrix of the ultramylonite behaved, in a more ductile manner than the neighbouring rock and this ductility contrast generated drag folds. Sampie MKS 147. Width of micrograph 6.3 mm. (c) Mylonite with quartz ribbons (1) cut by ~eudo~chyl~e (2) which is deformed to ul~~ylonite. The matrix of the p~udotachyl~e is mica-rich and flows round enclosed rock fragments. Sample MKS 125. Width of micrograph 2.5 mm. (d-0 Gradual transition from mylonite to ultramylonite: (d) muscovite is almost completely recrystallized and forms fine grained layers. Note weakly developed type II S-C fabric. Sample MKS 122. (e) A fine-grained matrix of quartz, albite and muscovite is generated. Small islands of relatively coarse grain size are still visible. Sample MKS 124-4. 0 The ultramylonite matrix shows a strong foliation. Layers with varying amounts of mica can be

distingu~hed. Size of clasts is reduced to < 0.5 mm. Sample MKS 124-4. Width of micrographs 5d-f is 6.3 mm.

296 N. KOCH AND I MASC‘H

(b)

Fig. 6. (a) Lattice preferred orientation of quartz c-axes in monomineralic quartz domains in ultramylonite; n = 84; (b) contoured at 0.5, 2.5, 4.5 and 6.5% intervals. The change from type I cross-girdle to oblique-girdle orientation is interpreted by Schmid and

Casey (1986) to result from deformation under plane strain with an increasing component of simple shear.

tween the mylonites and ultramylonites there ex- ists a transition zone with a width in the order of a few centimetres, and (2) type B uitramylonites developed from pseudotachylytes. These ultramylonites form sharp boundaries to the host rocks (Fig. 5a, b).

Both types of ultramylonite are well foliated and possess matrices consisting mainly of quartz, albite and muscovite, with grain sizes between 10

and 40 pm. The grain size in type B ultramylonites is generally less than that of type A ultra- mylunites. In type A ultramylonites, quartz also farms extremely elongated and completely recrystallized monomineralic domains (grain size about 50 pm), with quartz c-axes forming an oblique girdle to the foliation (Fig. 6). Relict clasts are observed in both types. Relict clasts in type B ultramylonites consist of fragments of host rock

Fig. 7. (a) Several generations of ultramylonite developed from pseudotachylyte: early pseudotachylyte is deformed to an ultramylonite band (I), which is cut by a later poeudotachylyte (21, that is in turn smeared out to ultramylonite (3). Sample MKS 149. Width of micrograph 6.3 mm. (b) Sketch of Fig. 7a illustrating the relations between different generations of ultramytonite and pseudotachylyte and suggesting interruption of ultramylonite formation by the generation of pseudotachylyte, which is in turn

further deformed to ultramylonite.

FORMATION OF ALPINE MYLONITES AND PSEUDOTACHYLYTES

enclosed in previously undeformed pseudotachylytes, whereas in type A ultramylonites the clasts are usually small relict grains of plagioclase. One sample suggests that the formation of ultramylonites was interrupted by the generation of pseudotachylyte (Fig. 7a, b).

TEM microstructures

Mylonitic gneiss and mylonites

Both the mylonitic gneisses and mylonites show identical TEM substructures. The density of free dislocations in relict quartz grains in all samples

297

investigated is high (8 x lo8 cme2 to > lo9 cmm2) and corresponds to strong undulose extinction. The dislocations occur straight as well as curved. Dislocation tangles (Fig. 8a) are typically found near high angle boundaries. Dislocation networks form low angle tilt boundaries and are commonly pinned to obstacles such as small grains of mica or infringing corners of grain boundaries (Fig. 8b). This compares well with the observations of Liddell et al. (1976) on unrecrystallized quartzites from the Bergell region of Switzerland deformed under greenschist facies conditions. In contrast to the relict grains, those recrystallized grains (grain size about 1 pm) that occur near strongly ser-

Fig. 8. TEM micrographs. (a) Tangled dislocations in a quartz grain adjacent to pseudotachylyte. On the left a thin planar zone of tangled dislocations corresponds to optically visible deformation lamellae. Sample MKS 10. Scale bar 1 pm. fb) Two sub-grain boundaries (tilt walls) pinned to a comer of an elongated quarts grain. Sample MKS 125. Scale bar 1 pm. (c) Recrystallized grains at serrated grain boundaries of quartz ribbons show highly variable dislocation densities. The boundaries of the grains with low dislocation densities show convex curvature where neighbouring grains possess a higher dislocation density. Note the absence of sub-grains in the deformed grain. Sample MKS 125. Scale bar 1 pm. (d) High-angle boundaries of recrystallized quartz grains

pinned to a small mica crystal (I). Note pinned low angle boundary (2). Sample MKS 125. Scale bar 1 pm.

29x N. KOC‘H AND I MASC‘H

rated grain boundaries in mylonites show a high variation in dislocation density between neighbouring grains (lo7 cmm2 to 10’ cmm2). The grain boundaries of recrystallized grains with low dislocation densities are Curved outward if a neighbouring grain has a higher dislocation density (Fig. 8~). This shows that the growth of undeformed grains occurred at the expense of highly strained grains; the driving force of grain- boundary motion being a difference in stored strain energy. (Surface energy-driven motion of grain boundaries would have the reverse sense. Additionally, dislocation densities in neighbouring grains should be equal; Poirier, 1985.)

The observed microstructures are, therefore, interpreted as preserved primary structures of migration recrystallization. Further grain growth was impeded by the pinning of grain boundaries to small micas (Fig. 8d). In those small recrystallized grains, dislocation networks are usually absent and dislocation tangles were not observed.

Pseudotachylyte host rocks

The dislocation density of relict quartz grains is generally > 2 X lo9 cm-* (obtained mainly from gneisses). Tangled dislocations are common, especially in grains in the immediate vicinity of pseudotachylytes. Recrystallized grains that are found in proximity to pseudotachylytes show a uniformly high density of free dislocations (z=- 3 x lo9 cm-*>. They are typically filled with dislocation tangles. This contrasts with the observations of recrystallized grains from the investigated mylonites. A striking feature is the general absence of well-ordered dislocation networks in the recrystallized grains adjacent to pseudotachylytes.

Ultramylonites

Recrystallized grains from monomineralic quartz domains in ultramylonites have an average dislocation density of 5 x lo8 cm-* to lo9 crnp2, Dislocation tangles were not observed. Sub-grain boundaries consisting of well-ordered dislocation networks are found more commonly in ultramylonites than in the recrystallized grains of mylonites. Dislocation densities of quartz grains in

ultramylonitic matrices also tend to be lower than in the mylonites. Average values cannot be pre- sented since, in photographs, quartz grains could not reliably be distinguished from albite grains, and the number of unambiguously identified quartz grains was too small.

In all rock types, optically visible sub-basal deformation lamellae correspond to a sharply bounded planar zone of tangled dislocations. Re- covered zones and sub-grain walls as described by Christie and Ardell (1976) were not observed.

Deformation mechanisms

The deformation mechanisms that operated at the base of the Silvretta nappe during it’s detachment can be determined using the combined optical and TEM microstructural and textural studies discussed above. In addition, information about the rheological behaviour during deformation of the basal lithologies can be obtained. In the case of plastic deformation, three principal types of experimentally obtained stress/strain curves have been distinguished with increasing strain after an initial rise in stress (e.g., Poirier, 1985; Twllis and Tullis, 1986): (1) the applied stress increases (strain hardening); (2) the applied stress de- creases (strain softening) and (3) the stress remains constant (steady state flow). Steady state flow in this context is the result of a balance between strain hardening and strain softening processes. These can, under natural conditions, occur in cycles over a short period of time, thus resulting in overall steady state flow conditions.

Dislocation tangles formed due to the mwltipli- cation and interaction of dislocations which im- pede easy glide are a feature of strain hardening (Poirier, 1985). The onset of recrystallization (migration and rotation) is seen as a cause of strain softening Wrai et al., 1986) leading to a concentration of the deformation in a shear band or a shear zone. Experimental deformation of magnesium has shown that the softening effect due to dynamic recrystallization is most markedly developed at low temperatures, although the temperature has to be high enough to permit recrystallization (Poirier, 198%.

FORMATION OF ALPINE MYLONITES AND PSEUDOTACHYLY-IES 299

Mylonitic gneisses and mylonites

The dominant deformation mechanism in quartz is by dislocation glide and climb. This is revealed by optical microscopy (e.g., elongation of grains, lattice-preferred orientation) and TEM investigations (e.g., high dislocation densities, curved dislocations). The following dominant mechanisms of strain softening can be observed:

(1) Softening due to lattice-preferred orientation. The observed c-axis preferred orientation is a type I cross girdle changing to an oblique girdle. This type of LPO is interpreted by Schmid and Casey (1986) to be the result of glide mainly occurring on the rhomb, prism and basal planes under conditions of plane strain with increasing components of simple shear. The glide direction was not measured, but Schmid and Casey (1986) identified (a) as the dominant glide direction in

comparable samples; (2) Softening due to dynamic recrystallization.

Strongly elongated ribbons show primary recrystallization starting at strongly serrated grain boundaries (Fig. 3a). TEM observations show the migration of convex, high-angle boundaries (driven by stored strain energy) and grains with low dislocation density consuming grains with high dislocation density (Fig. 8~). Rotation recrystallization can be inferred from the observed misorientation (loo to 20”) between small recrystallized grains and relict grains (Bell and Etheridge, 1976). TEM observations show several dislocation networks (tilt walls) in relict grains pinned to a common obstacle. This is interpreted as the initial state of coalescence of sub-grain boundaries preceding the formation of a high-angle boundary (Fig. 8b). The f ormation of high-angle boundaries by the coalescence of sub-grain boundaries is seen as the mechanism of rotation recrystallization (White, 1977).

Pseudotachylytes

Field observations (e.g., mylonites cut by pseudotachylyte veins, Fig. 5~1, optical observations (rotated quartz LPO in rock fragments within pseudotachylytes, Fig. 4) and TEM investigations of quartz (dislocation multiplication in grains ad-

jacent to pseudotachylyte veins) indicate an initial phase of plastic deformation prior to the formation of the pseudotachylytes. In contrast to the formation of mylonites, however, where strain softening processes are observed (see above), features correlating with strain hardening are seen to dominate over strain softening mechanisms in the pseudotachylyte host rocks. Dislocations in relict grains in the immediate proximity of pseudotachylytes are strongly tangled and no well ordered dislocation networks are to be observed. This implies that a multiplication of dislocations occurred, with dislocations becoming tangled and effectively blocking each other. The fine-grained

recrystallized quartz grains found at the edge of pseudotachylyte veins also show uniformly high dislocation densities with dislocation tangles and no sub-grain walls. This also demonstrates the domination of strain hardening following initial softening by dynamic recrystallization. Therefore, prevailing strain hardening in all quartz grains near pseudotachylytes has blocked further plastic deformation after an initial plastic phase. The reasons for this are discussed later.

Though paleopiezometers based on the density of free dislocations (McCormick, 1977; Weathers et al., 1979) and the size of recrystallized grains (Twiss, 1977; Mercier et al., 1977) still face numerous problems, as discussed by White (1979), and have to be applied with care, they can be used to give estimates regarding the differential stress distribution in mylonite zones (Christie and Ord, 1980; Ord and Christie, 1984). In the case of the Silvretta nappe, differential stresses exceed- ing 2 kbar can be inferred from both the size of recrystallized grains and their dislocation densities (Masch and Koch, 1988). This value corresponds to the upper value for the fault strength of rocks in the upper crust (Paterson, 1978; Lachenbruch, 1980). Such high differential stresses can initiate seismic fracturing with the resulting frictional heat causing the generation of pseudotachylyte melt (Sibson, 1975).

Ultramylonites

Apart from dislocation glide, it can be pre- sumed that the fine-grained matrix encountered

300 N.KOCH ANU I..MASCH

in ultramylonites has deformed by grain-boundary sliding (superplasticity). Superplastic behaviour in naturally deformed rocks was first reported by Bouillier and Gueguen (19751, followed by similar reports for virtually all rock types. The main condition that has to be met to facilitate superplastic flow by grain-boundary sliding is a small grain size. In order for metals to behave superplastically, it is known that the temperature must exceed roughly half the melting point (Haasen, 1974; Poirier, 1985). This does not seem to be necessary for materials deformed under geological strain rates, as can be concluded from the works of Behrmann (1985) and Schmid et al. (1981). The characteristic rheological feature of superplastically flowing aggregates should be an increased strain rate, a stress drop or a concentration of strain (Behrmann, 19851 resulting from an increase in ductility.

The observed high ductility of the ultramylonite matrices is an indicator for the reduced strength together with a concentration of strain (Fig. 5b). In addition, the larger grain size in the recrystallized monomineralic quartz domains (as compared to the size of recrystallized grains in mylonites), the lower dislocation densities and the absence of dislocation tangles point to a lower differential stress in the Silvretta ultramylonites.

The role of recovery

Recovery produces dislocation networks and sub-grains, while recrystallization is characterized by the appearance of new, strain-free grains and the migration of high-angle boundaries (Poirier, 1985). Metallurgists distinguish recovery and recrystallization processes that occur during (dynamic) and after (static) deformation (Haasen, 1974). Dynamic recovery is seen as a competitive mechanism with dynamic recrystallization to reduce stored strain energy (White, 1977).

Though recrystallization processes are readily seen in quartz from the Silvretta tectonites, the following observations point to the importance of recovery processes. Sub-grains are observed mostly in the recrystallized grains from quartz domains in ultramylonites, whereas in the small

recrystallized grains from mylonites they are rare. The occurrence of sub-grain boundaries pinned to obstacles in highly strained ribbon grains is seen as evidence for dynamic recovery, since they must have formed after the deformation and be- fore the onset of recrystallization (Fig. 8bl. The overall dominance of recrystallization over recovery compares well with the observations of Lid- dell et al. (1976). White (1977) suggests that high-stress conditions favour recrystallization processes. Static recovery caused by the thermal peak of Alpine metamorphism cannot have played an important role in the Silvretta tectonites, since the primary structures of initial recrystallization (e.g., recrystallized grains with highly differing dislocation densities), as well as the primary structures of strain hardening (e.g., dislocation tangles near pseudotachylyte veins), are not obliterated. Furthermore, only a few sub-grain walls in the small recrystallized grains formed in the mylonites could be detected.

Typical structures of static recovery should in- clude well-ordered sub-grain walls and sub-grains with only moderate dislocation densities (Poirier, 1985). The occurrence of ordered networks within relict grains can be attributed to pre-Alpine recovery processes and must not be taken as evidence for an Alpine static overprint. In grains directly adjacent to pseudotachylytes it is notable that there is no evidence of recovery. Only structures relating to strain hardening can be observed.

Causes for changes in the deformation regime

Brittle versus plastic deformation

Field observations (pseudotachylytes cutting mylonites, pseudotachylytes deformed to mylonites) at the base of the Silvretta nappe show that the pseudotachylytes and mylonites were both formed at the same depth. The common mineral assemblage imposed both on pseudotachylytes and mylonites (stilpnomelane zone) indicates that they were formed in the brittle-ductile transition zone (Sibson, 1986; Scholz, 1988). The compila- tion of maximum Alpine temperatures for the Silvretta nappe made by Hurford et al. (1989)


preclude the possibility of temperatures exceed- ing 300°C in the investigated area.

It has been shown above that not only my-

lonites but also pseudotachylytes have formed after an initial phase of plastic deformation. How is it to be explained that in one part of the Silvretta nappe strain hardening could have outweighed strain softening processes, while in other parts strain softening could lead to the development of well-defined mylonite zones?

Experimental work and investigations of naturally deformed quartzites have shown that the plastic behaviour of quartz is essentially con- trolled by three factors (e.g., Christie and Ardell, 1976; Poirier, 1985): (1) trace amounts of intracrystalline water; (2) temperature; and (3) strain rate. Trace amounts of intracrystalline water can drastically reduce the critical resolved shear stress of glide systems in quartz, as has been shown by Blacic (1975) and McLaren et al. (1989). In rocks containing mostly “wet” crystals, strain softening by lattice preferred orientation and dynamic recrystallization can thus substan- tially lower the required differential stress so that deformation is more easily facilitated. The deformation then usually concentrates in mylonite zones. In rocks with mostly “dry” crystals, plastic deformation is blocked by the tangling of dislocations, thus raising the required stress above a critical point that would finally lead to brittle failure accompanied by slip and the formation of pseudotachylytes. Little is known how pH,O in- fluences the content of intracrystalline H,O in quartz during deformation and recrystallization, but it seems reasonable to assume that deformation and recrystallization in a “wet” environment will result in grains with a higher content of intracrystalline H,O than deformation and recrystallization in a “dry” environment. Thus, strain hardening leading to seismic failure and the generation of pseudotachylytes is to be ex- pected in a “dry” environment. Note that Sibson (1986) postulates a low pH,O for the generation of pseudotachylytes, because of the heat balance required for frictional melting.

By studying the influence of the Alpine metamorphic gradient in the Gotthard profile on the deformation-induced microstructures of quartz,

Voll (1976) demonstrated that the temperature interval in which quartz deforms plastically with-

out recrystallization is relatively small (onset of plastic deformation at around 27O”C, onset of recrystallization at around 290-300°C). The strong plastic deformation accompanied by the onset of recrystallization with a very small grain size in the investigated samples of the Silvretta nappe shows that they were deformed at temperatures lying in this temperature interval (around 300°C). This statement compares well with the metamorphic assemblage, as well as the findings of Hurford et al. (1989). Very small variations in temperature, which could probably not be detected by differing metamorphic assemblages, could therefore drastically change the deformation behaviour of quartz. Slightly elevated temperatures would favour recrystallization as the dominant softening mechanism, while at lower temperatures multiplication of dislocations by plastic deformation without recrystallization would lead to the tangling of dislocations and strain hardening.

The high sensitivity to minute changes of temperature of the deformation behaviour of quartz at low temperatures is further confirmed by com- paring the samples of the Silvretta nappe to mylonites from the base of the lower Austro-Alpine Bernina nappe from the St. Moritz region. These lithologies still belong to the Alpine stilpnomelane zone, but must have experienced slightly higher temperatures because of their higher metamorphic grade. In samples investigated by us, quartz shows intense recrystallization (typical mortar structure) after moderate elongation. Also, the size of the recrystallized grains is significantly larger (around 100 pm) than in the mylonites from the Silvretta nappe.

Both factors (trace amounts of water and temperature variations) can explain why, in some regions of the Silvretta nappe, pseudotachylytes are dominant while in other areas mylonites are more widespread. However, the coexistence of mylonite and pseudotachylyte as observed in some specimens needs another explanation, and for these cases a dynamic model is proposed.

Sibson (1980) has described the occurrence of pseudotachylytes within low-grade mylonites from the Outer Hebrides Thrust and explained them

302 N. KOCH AND L. MAS(‘H

as transient discontinuities resulting from a rise a zone that is undergoing plastic deformation. of stress and strain rate around rigid blocks in an During the initiation of mylonitization in gneis- otherwise ductile environment. Scholz (1988) pro- sose lithologies, the stress in the more rigid blocks poses the possibility of the propagation of earth- of amphibolites rises to a critical point. Seismic quakes initiated in the brittle zone down to fracturing (accompanied by the generation of greater depths of dominantly ductile deformation pseudotachylyte) then imposes a sudden drastic in strike-slip faults. The model of Scholz (1988) is increase of strain rate on the neighbouring rocks. unacceptable in the case of the Silvretta nappe, This results in strain hardening which leads to since nappe movement took place along dip-slip fracture because of the inability of the deforming faults. The existence of pseudotachylytes within crystals to accommodate the defects at this rate mylonites and the pseudotachylytes themselves by softening mechanisms. The zone of seismic being smeared out to ultramylonites, can be ex- failure, therefore, propagates into an otherwise plained by a synthesis of the models of Sibson ductile environment. Stress relaxation after seis- (1980) and Scholz (1988). Pseudotachylytes are mic failure allows further plastic deformation at more commonly found in amphibolites which lack reasonable geological strain rates with the pseu-

dominant plastic deformation. The amphibolites dotachylytes becoming involved in lower strain can be interpreted as relatively rigid obstacles in rate deformation processes (Fig. 5~). Because the

TABLE 1

Dominant transmission optical microscopy (TOM) and TEMmicrostructures in mylonitic gneisses, mylonites, pseudotachylyte host rocks and ultramylonites

D, mylonitic gneiss

DZ, mylonite

Dz, pseudotachylyte (host rock)

D3

ultramylonite

TOM microstructures

Kink bands in micas. Elongation of quartz to rib-

bon-shaped grains. Only minor recrystalliza-

tion at serrated grain

boundaries and along deformation lamellae.

Stilpnomelane unde-

formed.

TEM microstructures

High dislocation densities.

Dislocation tangles.

Structures of recovery present but not dominant.

Recrystallization of muscovite. All quartz grains deformed to ribbons. No well-devel-

oped subgrains. Recrystallization starting at grain boundaries annihi- lates relict grains.

LPO of ribbon quartz: Type I cross girdle.

Stilpnomelane undeformed.

High dislocation densities

in relict grains

Strongly varying dislocation density in recrystallized grains at serrated grain

boundaries.

Migration and rotation recrystallization.

Structures of recovety present but not dominant.

Additionally very fine-

grained recrystallization in relict grains cut by pseudotachylyte veins.

Stilpnomelane mostly undeformed in PST.

Dislocation densities in relict grains adjacent to PST very high, numerous dislocation tangles.

Dislocation density in recrystallized grains uni-

formly very high, numerous dislocation tangles.

No recovery.

Proportion of mylonitic matrix > 90% (quartz, muscovite and albite) grain size 10 to 40 pm

LPO of monomineralic domains of completely recrystallized quartz (grain size 50 pm): oblique girdle.

Stilpnomelane syntectonic.

Dislocation densities in matrix quartz grains and recrystallized grains in quartz domains slightly lower than in mylonites.

No dislocation tangles.

Dynamic recovery in quartz domains.


continuing deformation obliterates traces of brit-

tle failure, it can be shown why pseudotachylytes are not more commonly found in mylonites, a view that is also taken up by Sibson (1980).

Transition to superplastic behaviour

The formation of type A ultramylonites from mylonites results from the continuous recrystallization of quartz and muscovite, with high stresses favouring small grain sizes. This is because the different phases are impeding each other in grain growth so that grain boundary sliding becomes effective. The formation of type B ultramylonites from pseudotachylytes is favoured by the extremely fine-grained pseudotachylyte matrix which is rich in phyllosilicates and/or inosilicates and quartz. This matrix could even have been glassy after solidification. Both possibilities pro- vide favourable conditions for superplastic behaviour. Glasses are known from metallurgy to be easily deformable by super-plasticity (Haasen, 1974). Further factors favouring grain-boundary sliding are: (1) a slight rise in temperature as Alpine metamorphism reaches its thermal peak and (2) an infiltration of fluid phases (documented by the development of OH-bearing phases) as the Silvretta nappe is thrust over wet Penninic strata.

Correlation of deformation with Alpine tectonics

Table 1 and Figure 9 present an overview of the observed microstructures, deformation phases and their correlation with Alpine tectonics and metamorphism.

The first Alpine deformation phase (Dl) in the Silvretta nappe involves penetrative plastic deformation that generates the mylonitic gneisses. The overall strain is low, as is documented by the still recognizable pre-Alpine texture. D 1 produces elongated quartz grains and kink bands in micas. Mylonites (D2J and pseudotachylytes CD,) are both formed during a subsequent phase. While D,, mylonite formation is interpreted to be a continuous development of D,, the formation of D,, pseudotachylytes is a cyclic event. This is indicated by the presence of up to three generations of pseudotachylyte that can be distinguished in one sample (Masch, 1974). The formation of pseudotachylytes is synchronous with the formation of mylonites CD,,). This is documented by the appearance of discordant pseudotachylyte veins cutting through mylonites and with the pseudotachylytes becoming involved in subsequent deformation. The Eo-Alpine metamorphism (stilpnomelane) is post-kinematic to both D, and D,. The formation of the ultramylonites during D, is the last ductile deformation phase

mylonitic 01

P seudotachylytes

D2b

qneisses _____++_. . . . . .

\ mylonites D2a

ultramvlonites ‘.

thermal climax of

Fig. 9. Relationships between mylonitic gneisses, mylonites, pseudotachylytes and ultramylonites, Alpidic tectonic events and Eo-alpine metamorphism. Phases D,, D, and D, are interpreted to be a continuous sequence. D, (formation of mylonites) is interrupted by the cyclic formation of pseudotachylyte (D,,). D, and D, can be correlated with the detachment of the Silvretta nappe. D, (formation of ultramylonites) coincides with the peak of the Eo-Alpine metamorphism (stilpnomelane zone) in the

investigated area. Pseudotachylyte generation can also persist into the early stages of D,.

N. KOCH AND I MASCH

03 1 I Fig. 10. (a) Schematic representation of pseudotachylyte oc- currences. Correlatable amphibolite horizons (up to several metres thick) are shown stippled. (b) Extensional features in an otherwise compressional regime of nappe thrusting are explained by Laubscher (1983) to be the result of a down step of the thrust plane in the direction of motion of the upper

plate.

and, in the investigated area, occurred during the peak of Alpine metamorphism. As such, D,, D,, and D, are not interpreted as distinct phases, but rather as a continuous development with increasing strain. This development is interrupted during mylonite formation by cycles of seismic fracturing and slip leading to the generation of the pseudotachylytes. These seismic interruptions mainly occurred during mylonite formation, but also ex- tended into the early stages of D3.

Geometric analysis of the Silvretta pseudotachylytes has shown that they were formed mostly under extensional conditions (Schmutz, 1982; and Fig. 10 in this work), though some compressional features are also found. The general extensional setting of the pseudotachylytes in an overall compressional regime (plate convergence, nappe thrusting) is explained by Laubscher (1983) as a result of the detachment of the Silvretta nappe as an “erogenic lid” (the “traineau Ccraseur”, Ter- mier, 1903) from its basement. Voids for pseudotachylyte intrusion were apparently developed where the thrust plane stepped down in the direction of motion of the upper plate (Fig. 10).

Phases D,, D,, and D,, are, therefore, correlated with the detachment of the Silvretta nappe, while the formation of ultramylonites (D,) and the thermal peak of Eo-alpine metamorphism are correlated with an early stage of transport and the onset of fluid infiltration from the underlying Penninic strata. During subsequent nappe thrusting the lithologies at the base of the Silvretta nappe were passively transported, whereas the underlying Penninic sedimentary and ophiolitic rocks acted as the main slip horizon (Laubscher, 1983).

Isotopic dating (Hurford et al., 1989) in the Silvretta nappe gives a minimum age of 90 Ma (Late Cretaceous) for the Eo-Alpine metamorphism which, in consequence, must be the minimum age for D, to D,. By this time the Silvretta nappe could not have reached its present posi- tion, because sedimentation up to the Oligocene is documented in the Lower Engadine Window (Oberhauser, 1983).

Acknowledgements

This work formed part of a Ph.D. thesis and was supported by the Deutsche Forschungsge- meinschaft for twelve months under grant number MA 686 14-2. We thank John Spray, Mark Swanson and an unknown reviewer for critically appraising the manuscript.

References

Amann, A., 1985. Zur Metamorphose des niirdlichen Sil- vretta-Kristallins. Ph.D. thesis, Innsbruck Univ. (unpubl.).

Bearth, P., 1932. Geologie und Petrographie der Keschgruppe. Schweiz Mineral. Petrogr. Mitt., 12: 256-279.

Bearth, P., 1933. uber Gangmylonite der Silvretta. Schweiz Mineral. Petrogr. Mitt., 13: 347-348.

Behrmann, J.H., 1985. Crystal plasticity and superplasticity in quartzite: a natural example. Tectonophysics, 115: lOl- 129.

Bell, T.H. and Etheridge, M.A., 1976. The deformation and recrystallization of quartz in a mylonite zone, Central Australia. Tectonophysics, 32: 235-267.

Blacic, J.D., 1975. Plastic deformation mechanisms in quartz, the effect of water. Tectonophysics, 27: 271-294.

Bouillier, A.M. and Gueguen, Y., 1975. S-P mylonites: origin of some mylonites by superplastic flow. Contrib. Mineral.

Petrol., 50: 93-104.


Cad&h, J., Bearth, P. and Spaenhauer, F., 1941. Erhiuterun-

gen zum geologischen Atlas der Schweiz. 1:25,000, Blatt

420, Ardez. Christie, J.M. and Ardell, A.J., 1976. Deformation structures

in minerals. In: H.-R. Wenk (Editor), Electron Microscopy of Minerals. Springer, New York, pp. 374-403.

Christie, J.M. and Ord, A., 1980. Flow stress from microstructures of mylonites: example and current assessment. J. Geophys. Res., 85: 6253-6262.

Grauert, B., 1966. Rb-Sr age determinations on orthogneisses of the Silvretta (Switzerland). Earth Planet. Sci. Lett., 1:

139-147. Haasen, P., 1974. Physikalische Metallkunde. Springer, New

York, 379 pp. Ham, R.K., 1961. The determination of dislocation densities

in thin films. Philos. Mag., 6: 1183-1184. Hurford, A.J., Flisch, M. and Jiiger, E., 1989. Unravelling the

thermo-tectonic evolution of the Alps: a contribution from fission track analysis and mica dating. In: M.P. Coward, D. Dietrich and R.G. Park (Editors), Alpine Tectonics. Geol.

Sot. London, Spec. Publ., 45: 369-398. Lachenbruch, A.H., 1980. Frictional heating, fluid pressure

and the resistance to fault motion. J. Geophys. Res., 85:

6097-6112. Laubscher, H.P., 1983. Detachment, shear and compression in

the Central Alps. Geol. Sot. Am., Mem., 158: 191-211. Liddell, N.A., Phakey, P.P. and Wenk, H.-R., 1976. The

microstructure of some naturally deformed quartzites. In: H.-R. Wenk (Editor), Electron Microscopy of Minerals. Springer-Verlag, New York, pp. 419-427.

Macaudiere, J., Brown, W.L. and Ohnenstetter, D., 1985. Microcrystalline textures resulting from rapid crystallization in a pseudotachylite melt in a meta-anorthosite. Con- trib. Mineral. Petrol., 89: 39-51.

Maddock, R.H., 1986. Partial melting of Iithic porphyroclasts in fault-generated pseudotachylytes. Neues Jahrb. Min- eral., Abh., 155: 1-14.

Magetti, M. and Galetti, G., 1988. Evolution of the Silvretta eclogites: metamorphic and magmatic events. Schweiz Mineral. Petrogr. Mitt., 68: 467-484.

Magloughlin, J.F., 1992. Microstructural and chemical changes associated with cataclasis and friction melting at shallow crustal levels: the cataclasite-pseudotachylyte connection.

In: J.F. Magloughlin and J.G. Spray (Editors), Frictional Melting Processess and Products in Geological Materials. Tectonophysics, 204: 243-260.

Masch, L., 1970. Die Pseudotachylite der Silvretta. Eine Un- tersuchung ihrer Deformation und Aufschmelzung. Ph.D. thesis, Munich Univ. (unpubl.).

Masch, L., 1974. Untersuchung der Aufschmelzung und De- formation der Pseudotachylite der Silvretta (Gsterreich, Schweiz). Neues Jahrb. Mineral., Monatsh., 11: 485-509.

Masch, L., 1979. Deformation and fusion of two fault rocks in relation to their depth of formation: the hyalomylonites of

Langtang (Himalaya) and the pseudotachylites of the Sil- vretta Nappe (Eastern Alps). Proc. 8th Conf. on Analysis

of Actual Fault Zones in Bedrock. U.S. Geological Survey,

Menlo Park, pp. 528-533. Masch, L. and Koch, N., 1988. Deformationsstrukturen der

Mylonite der Zone von Erbendorf-Vohenstrauss (ZEV) und Vergleich mit Basismyloniten der alpidischen Sil- vretta-Decke (Graubiinden). Erlanger Geol. Abh., 116:

81-100. McCormick, J.W., 1977. Transmission electron microscopy of

experimentally deformed synthetic quartz. Ph.D thesis,

Univ. Calif., Los Angeles (unpubl.). McLaren, A.C., Fitzgerald, J.D. and Gerretsen, J., 1989. Dis-

location nucleation and multiplication in synthetic quartz: relevance to water weakening. Phys. Chem. Miner., 16:

465-482. Mercier, J.-C., Anderson, D.A. and Carter, N.L., 1977. Stress

in the lithosphere: inference from steady state flow of rocks. Pure Appl. Geophys., 115: 199-226.

Niggli, E., 1970. Alpine Metamorphose und alpine Gebirgsbil- dung. Fortschr. Mineral., 47: 16-26.

Oberhauser, R., 1983. Mikrofossilfunde im Norwestteil des

Unterengadiner Fensters sowie im Verspalaflysch des Ratikon. Jahrb. Geol. Bundesaust. (Austria), 126: 71-93.

Ord, A. and Christie, J.M., 1984. Flow stresses from mi-

crostructures in mylonitic quartzites of the Moine Thrust zone, Assynt area, Scotland. J. Struct. Geol., 6: 639-654.

Passchier, C.W., 1982. Mylonitic deformation in the Saint- Barthtlemy massif, French Pyrenees, with emphasis on the genetic relationship between ultramylonite and pseudotachylite. GUA (Municiple Univ. Amsterdam) Pap. Geol., l(16): 1-173.

Paterson, M.S., 1978. Experimental Rock Deformation: the Brittle Field. Springer, New York, pp. 21-42.

Poirier, J.-P., 1985. Creep of Crystals. High Temperature Deformation Processes in Metals, Ceramics and Minerals. Cambridge University Press, Cambridge, U.K., 260 pp.

Schmid, SM. and Casey, M., 1986. Complete fabric analysis of some commonly observed quartz c-axis patterns. In: B.E. Hobbs and H.C. Heard (Editors), Mineral and Rock Deformation: Laboratory Studies. Geophys. Monogr., Am. Geophys. Union, 36: 263-286.

Schmid, SM., Casey, M. and Starkey, J., 1981. The microfab- ric of calcite tectonites from Helvetic nappes (Swiss Alps). In: K.R. McCIay and and N.J. Price (Editors), Thrust and

Nappe Tectonics. Geol. Sot. London, Spec. Publ., 9: 151- 158.

Schmutz, U., 1982. Ordnung und Chaos in den Pseudo- tachyliten der Silvretta am SW-Rand des Unterengadiner Fensters. Unpubl. manuscript, Base].

Scholz, C.H., 1988. The brittle-plastic transition and the depth. of seismic faulting. Geol. Rundsch., 77: 319-328.

Sibson, R.H., 1975. Generation of pseudotachylyte by ancient seismic faulting. Geophys. J. R. Astron. Sot., 43: 775-794.

Sibson, R.H., 1977. Fault rocks and fault mechanisms. J. Geol. Sot. London, 133: 191-213.

Sibson, R.H., 1980. Transient discontinuities in ductile shear zones, J. Struct. Geol., 2: 1655171.

306

Sibson, R.H., 1986. Earthquakes and rock deformation in

crustal fault zones. Annu. Rev. Earth Planet. Sci., 14:

149-175.

Spaenhauer, F., 1932. Petrographic und Geologic der Gra-

letsch-Vadret-Sarsurs Gruppe (Graubiindenb Schweiz

Mineral. Petrogr. Mitt., 12: 27-146.

Streckeisen, A., 1928. Geologie und Petrographie der

Fliielagruppe (Graubiinden). Schweiz Mineral. Petrogr.

Mitt., 8: 87-239.

Termier, P., 1903. Les nappes de recouvrement des Alpes

Orientales et la synthese des Alpes. Bull. Sot. GCol. Fr., 4:

711-765.

Thoni, M., 1982. Der Einfluss der kretazischen Metamor-

phose im Westabschnitt der ostalpinen Einheit: Interpre-

tation geochronologischer Daten. Mitt. Ges. Geol. Berg-

baustud. Gsterr., 28: 17-34.

Thorn, M., 1987. Rb-Sr isotopic evidence for mid-Upper

Cretaceous detachment and thrusting of the Austro-Al-

pine Nappes. Terra Cognita, 7: 72 (abstr.).

Triimpy, R., 1980. Geology of Switzerland: A Guide Book.

Wepf, Basel, Part A: 104 pp; Part B: 232 pp.

Tullis, T.E. and Tullis, J., 1986. Experimental rock deforma-

tion techniques. In: B.E. Hobbs and H.C. Heard (Editors),

Mineral and Rock Deformation: Laboratory Studies. Geo-

phys. Monogr, Am. Geophys. Union, 36: 297-324.

Twiss, R.J., 1977. Theory and applicability of a recrystallized

grain size paleopiezometer. Pure Appl. Geophys., 115:

227-244.

Urai, J.L., Means. W.D. and Lister, G.S., 1986. Dynamic

recrystallization of minerals. In: B.E. Hobbs and H.C.

Heard (Editors), Mineral and Rock Deformation: Labora-

tory Studies. Geophys. Monogr., Am. Geophys. tJnion, 36:

161-199.

Voll. G., 1976. Recrystallization of quartz, biotite and feldspars

from Erstfeld of the Leventina Nappe, Swiss Alps, and its

geological significance. Schweiz Mineral. Petrogr. Mitt..

56: 64 I-647.

Weathers, M.S., Cooper, R.F., Kohlstedt, D.L. and Bird, J.M.,

1979. Differential stress determined from deformation-in-

duced microstructures of the Moine Thrust zone. J. Geo-

phys. Res., 84: 7495-75097.

Wenk, E., 1934. Beitrge zur Petrographie und Geologic des

Silvrettakristallins. Schweiz Mineral. Petrogr. Mitt., 14:

196-278.

Wenk, H.-R., 1978. Are pseudotachylites products of fracture

or fusion? Geology, 6: 507-511.

White, S., 1977. Geological significance of recovery and re-

crystallization processes in quartz. Tectonophysics, 39:

143-170.

White, S., 1979. Difficulties associated with paleostress esti-

mates. Bull. Mineral., 102: 210-215.

Winkler, H.G.F., 1979. Petrogenesis of Metamorphic Rocks.

Springer-Verlag, New York, 5th edn., 348 pp.

formation of alpine mylonites and pseudotachylytes at the base of the silvretta nappe, eastern alps

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