Austroalpine boundary - CiteSeerX

Journal ofthe Geological Socieq, London, Vol. 146, 1989, pp. 835-849, 11 figs. Printed in Northern Ireland

Kinematics of the Alpine plate-margin: structural styles, strain and motion along the Penninic- Austroalpine boundary in the Swiss- Austrian Alps U . RING',L.RATSCHBACHER', W . FRISCH', D. B I E H L E R ' & M . KRALIK2 Institut fur Geologie, Universitat, 0-7400Tubingen, FRG 'Geotechnisches Institut, BVFA-Arsenal, A-l031 Wien, Austria Abstraet: A structural and kinematic analysis of parts of the Alpine suture zone, i.e. the Arosa Zone and the edges of the overlying Austroalpine and underlying Penninic units, is presented. The Arosa Zone is interpreted as part of an accretionary wedge formed by the overriding of the Austroalpine units, the northern part of the Adriatic plate, over the Penninic units. Structures, strain and stretching trajectoriesindicate that two main deformationevents followed eachother discontinuously. The earlierevent was a WSW- to NW-directedthrusting under noncoaxial deformation and led to melange-type structures. The second event, less penetrative in the Arosa Zone and with a downward increase in strain and noncoaxial flow, formed N- to NE-vergingfolds and imbricationzones. West-directeddisplacementcommenced in theEarly Cretaceousandthe motionchanged to a prevailing northerly direction during the Eocene, as revealed by sediments involved in thrusting and by radiometric dating of the mylonites. The kinematic history deduced from structural analysis agrees with the history obtained by recent work onplate margin sediments. Wecorrelate the anticlockwise rotation of the Adriatic plate relative to Europe between about 130 and 60 Ma with the west-directed motion in the Arosa Zone. A change to relative N-S motion during the Lower Tertiary led to the N to NE directed displacements. A direct relationshipthus exists between the motion of the Adriaticplate andthe resulting deformation along its northern margin during the early history of the Eastern Alps. Partitioning in brittle and ductile deformation and structural slicing under noncoaxial deformation is responsible for the melange-type structures.

1985). In the Arosa Zone, parts of the overriding plate are also involved in the accretion process. A similar case was reported by Frisch et al. (1987) from the Matrei zone of the Tauern Window (Fig. lb) which is considered to represent an equivalent to the Arosa Zone further to the E (Waibel & Frisch 1989). Using structural, strain and kinematic data, we will show four features. (1) The Arosa Zone is part of an imbricate thrust stack genetically dominated by the overriding of the Austroalpine units and not by accretionary stacking related to subduction; the thrusting led to the formation of zones of tectonic melange. (2) Deformation is heterogeneous on all scales, prevailingly noncoaxial, and partitioned into ductile and brittle processes. (3) A kinematic analysis is, nevertheless, possible; superposedevents are resolved in terms of progressive deformation which resulted in multiple translation paths of single nappes. The age of thrusting is constrained by sediments involved and radiometric data of mylonites. (4) Thedeformation history indicates thatthe tectonic pattern in theEastern Alps is not the result of simple orthogonalcompression.Structural and sedimentological evolution consistently demonstratethatdextral transpression dominated during the Cretaceous and changed to about orthogonal convergence in the Early Tertiary. The kinematic history of deformation at the PenninicAustroalpineplate margin thus reflects the CretaceousEarly Teriary motion path of the Adriatic plate relative to Europe. We will concentrate on the kinematic evolution of the Penninic-Austroalpine platemargin.Thisincludes,from bottom totop, Middle Penninic, South Penninic, and

Some current interests in structural geology on convergent plate margins concern (1) the relationship between large-scale plate motions and deformation along its margins (Choukroune et al. 1986), (2) the accretion process, melange formation, deformation mechanisms, and deformation partitioning in brittle and ductile processes in accretionary wedges (Cowan 1985; Platt 1986), and (3) the interaction between tectonic and sedimentary processes (Hsii 1982). Recent studies of the structural and sedimentary evolution of Austroalpine units and the Austroalpine-Penninic boundary in theEasternAlps (Fig. l a , b) indicate a probable link betweenplatemotions and displacements of nappes (Triimpy 1975; Ratschbacher et al. 1987). An accretionary wedge formed along the western and northern rims of the Austroalpineunits (Frisch et al. 1987; Winkler 1987). The Sedimentary evolution along this plate margin reflects changes in the overallconvergence(Weissert & Bernoulli 1985; Waibel & Frisch 1989). This article presents a kinematic analysis of the Arosa Zone and adjacent portions of the Austroalpine and Penninic units. The Arosa Zone is part of the main suture zone of the Alps (Fig. lb) along which the South Penninic ocean was subducted underneaththeAustroalpine continental margin (Ernst 1971; Dietrich 1976; Frisch 1979). At map-scale, this zone displays a block-in-matrix structure including blocks of differentgeotectonicenvironments. In some parts, deformation is enhanced leading to characteristic features of tectonic melanges. Melanges generally form by accretion anddeformation of theupper portion of subducting oceanic crust and its overlying sediments (Hsii 1971; Moore & Karig 1976, 1980; Raymond 1984; Cowan 835

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Fig. 1. (a) Tectonic sketch map of the Eastern Alps. (b) detail of the western Part of the Eastern Alps and lithotectonic units referred to in the text. 1, Austroalpine units; 2, Penninic units. (c) Palaeotectonic reconstruction of the Penninic ocean in the Jurassic, after Weissert & Bernoulli (1985).

Austroalpine nappes located atthe western end of the Eastern Alps and Penninic nappes in the lower Engadine Window (Figs l a , b and 2). Moreover, we will address some problems of accretionary wedge deformation.

Tectonic overview During Triassic and Jurassic times, a period of rifting and associated stretching affected the Austroalpine realm, which became part of the Adriatic microplate in the Cretaceous (Channel1 & Horvath 1976; Frisch 1979; Dercourt et al. 1986) and the upper plate during the Alpine orogeny. The rifting process prograded northwestward, and led tothe formation of oceanic crust in the South Penninic ocean in theLate Jurassic and Early Cretaceous(Dietrich 1976). Palaeotectonic reconstructions of the Jurassic and lowermost Cretaceous (Fig. lc) and analogous models of the contemporaneous opening of the Atlanticoceanindicate that extension was mainly controlled by strike-slip faults

(Bosellini 1981; Weissert & Bernoulli 1985; Waibel & Frisch 1989). Compressional motions began in the Austroalpine in Early Cretaceous times (f130 Ma; Frank et al. 1987 for the Eastern Alps). They progradefi :o the foreland, affecting the Penninic units during the Gztaceous and the Helvetic units during the Eocene (Frisch 1979; Trumpy 1980; Laubscher & Bernoulli 1982). Palaeomagnetic studies show that the Adriatic plate moved independently from the African plate and rotated counterclockwise with respect to Eurasia in the Cretaceous (130 to 60 Ma; Zijderveld et al. 1970; Dercourt et al. 1986). Thisrotation caused dextral NW-SE directed transpression and a dextral strike-slip component onand along its northern margin, the Austroalpine margin (Trumpy 1975; Laubscher & Bernoulli 1982; Ratschbacher et al. 1987). During the Tertiary, the AdriaticPlate was linked to Africa and moved northward. Displacement analyses based on microstructures in ductile mylonites of thrust and shear zones, and studies on thrustgeometries,

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Fig. 2. Locations cited in text and tectonic units along the western margin of the Eastern Alps (Graubiinden-Ratikon). Heavy dots, Tertiary Bergell intrusion; black, uppermost Penninic units: Arosa Zone and equivalent ophiolite-bearing tectonic units. Modified after Weissert & Bernoulli (1985). confirm transpressional W- to NW-directed motions in the Austroalpine and the South Penninic Nappes during early stages of the Alpine orogeny (Laubscher & Bernoulli 1982; Mazurek 1986; Schmid & Haas 1987; Ratschbacher 1986, 1987; Ratschbacher & Neubauer 1989). IntheEastern Alps, this early motion was followed byN-S compression (Ratschbacher 1986, 1987). This is corroborated by a study of Jurassic to Lower Eocene sediments in the Lower Engadine Window (Fig. l b ; Waibel & Frisch 1989). Cretaceous sediments there reflect a close proximity to their Austroalpinesource areaandthus indicate lateral, E-W directed relative motion. The highly irregular leading edge of the Austroalpine unit partitioned the deformation at the plate margin in strike slip and convergence.Subduction along the convergent sections accounts for Early Cretaceous & Bernoulli 1986). high-P metamorphic minerals (Winkler During the Tertiary, rapid accretion of sediments along the former strike-slip margins suggests a change to compressional movementdirectedapproximately N-S (Waibel & Frisch 1989).

Geological-tectonic settingand structural style The Austroalpine units of the study area are presented by the Silvretta and Err-Bernina thrust systems (Upperand

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Lower Austroalpine, respectively; Trumpy 1980). Both thrust systems consist of slices of Mesozoic rocks and crystalline basement. In the Graubunden-Ratikon area at the western edge of the Eastern Alps (Fig. 2), the Penninic unit is subdivided into South and Middle Penninic units. The South Penninic nappescontain pelagic limestone,shale,andradiolarite associated with ophiolites. Flysch is of Aptian to Turonian age and no sedimentsyounger than Coniacian are known from the South Penninic realm (Oberhauser 1983; Winkler 1987; Ludin 1987). The Middle Penninic,stacked into the Falknis-Sulzfluh thrust system, encompasses massive limestone, breccia, mar1 (CouchesRouges) and flysch. The youngest rocks encountered in the study area are PalaeoceneCouchesRouges (Oberhauser 1983). In the Lower Engadine Window (Fig. lb) which forms an eye-lid window (Boyer & Elliott 1982; Boyer 1987), the Penninic units are represented by a still not well understood stack of flysch and wildflysch (Oberhauser 1983; Waibel & Frisch 1989). The rocks of the Tasna zone (Fig. 2) are interpreted by Waibel & Frisch (1989) as being deposited close to the Austroalpine margin. The Tasna zone is generally regarded as a continuation of the Middle Penninic Zone of the Central and Western Alps (Triimpy 1980). The centre of the window consists of ‘Bundnerschiefer’ which is connected with the Rhenodanubian flysch troughstretching all along of theEastern Alps. The youngest thenorthernedge sediments in the Lower Engadine Window are Lower Eocene in age (Oberhauser 1983; Rudolph 1982). The Arosa Zone (Figs l b and 2 ) , sandwiched between the Austroalpine nappes on top andMiddle Penninic nappes below, is a highly tectonized ophiolite-bearing unit which is a few tens of metres to more than one thousandmetres thick. The occurrence of ophiolites has led most workers to consider the Arosa Zone as a South Penninic unit (Weissert & Bernoulli 1985; Triimpy 1975). TheArosaZone is exposed along the western edge of the Eastern Alps (Fig. 2). Its northern continuation is represented by the dismembered slices of the Walsertal Zone (Winkler 1987). Inthe south it is associated with the large ophiolitic assemblages of the Platta Nappe and the Malenco-Forno Zone (Fig. 2). The Arosa Zone reappears in the Gargellen and Lower Engadine Windows. Another equivalent of the Arosa Zone is seen in the Matrei Zone of theTauern Window (Frisch et al. 1987; Fig. lb). Idealized stratigraphic South, Middle Penninic and Lower columns of the Austroalpine units for the study area from Winkler (1987), Ludin (1987) and Waibel (1985) are shown in Fig. 3. The Arosa Zone contains a number of imbricated slices with blocks andsequences of bothPenninic and Austroalpine origin, which float in a matrix of mid-Cretaceous shale and flysch of South Penninic origin. Components thought to be of South Penninic derivation are Jurassic to Cretaceous ophiolitic rocks (serpentinite,ophicalcite,basalt), radiolarian chert, and pelagic limestone.Austroalpine slices and blocks, in part olistoliths, are composed of Variscan basement rocks, Permo-Scythian arkose and quartzite, Middle to Upper Triassic dolomite (mostly Upper Triassic Hauptdolomit), and Jurassicsandstone with intercalated scarp breccia, radiolarian chert and pelagic limestone. Arosa Zone close the to The location of the Austroalpine margin is indicated by the occurrence of breccias and numerousintercalated blocks and slices of Austroalpine origin, which gives an overall melange

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Fig. 3. Simplified columnar stratigraphic sectionsof the Jurassic to Tertiary of the Middle Penninic (Falknis-Sulzfluh), South Penninic (Arosa Zone) and Lower Austroalpine units at the western end of the Eastern Alps (Graubiinden-Ratikon) andof the Tasna unitof the Lower Engadine Window in the studied area; after Winkler (1987), Liidin (1987), and Waibel& Frisch (1989).

character to this zone. However, the Arosa Zone as a whole is not a typical melange as defined by Cowan (1985). Coherent units are mappable on a scale of 1 : 10 000 and characteristic scale-independent block-in-matrix assemblages are restricted to shear and thrustzones. At map-scale we distinguished two different block-in-matrix assemblages of tectonic origin in the Arosa-Davos area. In the lower part of the Arosa Zone, exotic blocks or slices measuring up to 2 km along strike float in a sheared serpentinite matrix. The blocks are of Austroalpineorigin. Closely spaced shear fractures in the serpentinite define a planar foliation which wraps around phacoids of basalt and pyroxenite. Deformation structures indicate brittle rock behaviour of the blocks and cataclastic flow in the serpentinite. In descriptive terms, this assemblage resemblesaserpentiniticmelange(Cowan 1985; Saleeby 1982; Gansser 1974). In the upper part of the Arosa Zone blocks of Austroalpinederivation of smaller size (up to 400111 in length) are embedded in a matrix of black shale and flysch. The internal structure of the flysch matrix isin partchaotic with disrupted competent beds floating in a matrix of penetratively cleaved shale. Blocks of highly competent rocks like greenstoneorHauptdolomit show a concentration of fracturing near their margins and along thrusts. Most of the fractures are extension veins. It is suggested that they formed during thrusting.

Methods of kinematic analysis Overprinting criteria, orientation and nature of structures as well as continuous tracing of penetrative structural elements on a regional scale have enabled us to distinguish two major deformation events which are referred to as D1 and D2. We reconstructed the strain history by analyzing the sequence of vein formation and accompanying fibre growth in brittle rocks, and by measuring pebble shapes in ductile rocks, following a strain gradient from high crustal levels to deep parts of the pile. Finite strain values were calculated

(Rf / @ from shapes of deformed breccias andpebbles method, Lisle 1985), quartz grains, coarse calcite grains, foraminifera(Fry-method, Fry 1979), and extension veins (Ramsay & Huber 1983). Fault slip data were used to constrain strain orientations inbrittle rocks (ophiolite, sandstone; right dihedron method, Angelier 1984). Field study has shown that the D1 event dominates in the upper portions of the pile, while the D2 event becomes increasingly more intense with depth. It was the major goal of our finite strain analysis to map this strain superposition and the strain gradient, and to determine quantitatively the geometry and intensity of the single events. Breccias which were stretchedduring D 1 were collected from the upper parts of the pile; those samples, which suffered no or minor D2 are thought to represent the D1 strain. The incremental Dl-D2 superposition and total strain were analysed either in more ductile rocks or in thedeep parts of the pile. Incremental strain was derived from mineral-fibre studies in extension veins in the XY plane of finite strain ( X > Y > Z , principal strains). Calcite and, less frequently, quartz fibres grew during the D1 and D2 events. Information on the rotational component of strain was obtained from shear criteria in the XZ plane of incremental strain: S-C fabrics (Berth6 et al. 1979), asymmetrical pull aparts (Hanmer 1986), asymmetric strain shadows, rotated remnants of fracturedbrittle minerals, sheared minerals (Simpson & Schmid 1983) and curvilinear fibres (Etchecopar & Malavieille 1987). Textures of fine-grained calcite mylonites of D1 and D2 shear zones were studied with a texture goniometer. Displacement dataare based on two assumptions: (1) strains are commonly large in the deformation zones of the pile (Choukroune et al. 1986); (2) simple shear was the dominant component of deformation at least during the main period of rock flow. We justify the latter assumption by combining the planestrainincrementsderived from strain analysis with the general rotational deformation. By

A D R I A T I C - E U R O P E A NP L A T EK I N E M A T I C S ,A L P S

sampling at close intervals and recognizing consistent small scale displacements, we reconstruct the kinematics on a regional scale.

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Penetrative structures penetrative A foliation, S1, is well developed in incompetent rocks and generally runs subparallel to bedding andthrusts. S1 is defined by fine-grained phyllosilicates, films of opaque minerals,elongated quartz and calcite grains, and truncated fossils. AnastomosingS-patterns are common, and a distinction between S- an4 C-surfaces (Berth6 et al. 1979) is possible in moderatelydeformed rocks (Fig. 4a, b). In highly sheared zones, only one set of surfaces interpreted as C-surfaces is recognized. These are marked by opaque minerals and fine-grained phyllosilicates, whereas S-planes are mainly defined by grain elongation. High-strain shear zones show the development of a third or of multiple sets of S-planes with extensional character (Fig.

Structures Slump folds are the only sedimentary deformation structures recognized. They are distinguished from tectonicfolds by highly variable fold style, lack of axial plane cleavage, and by subsequent S1 overprinting. A strongly overprinted pre-D1 foliation is preserved locally. It is of the spaced type (Powell 1979) and consists of a series of fine seams which are defined by an alignment of opaque minerals and phyllosilicates.

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fig. 4. (a) to (e) D1meso- and microstructures; (a) Small scale shear zone in Jurassic shales west of Davos, shear band gives same sense of shear; (b) Ductile-brittle S-C structures in black shales, Verajoch, Ratikon; (c) S-C structures and shearbands in serpentinite, N of Arosa; (a) Layer disruption by small scale shear zones, Verspalaflysch, Ratikon; ( e ) a-shaped pressure shadow around and extension fracturing within siliceous limestone clast of polymict flysch breccia, north of Klosters. Sense of shear is top WSW to WNW in all sketches. ( f ) to (g) D2 mesoand microstructures; (f) N-verging F2 fold cut by back limb thrust, flysch of the Walsertal Zone; rotation direction is top N to NNE; (g) a-shaped, quartz-filled pressure shadows behind syn-D2 pyrite in Palaeocene Couches Rouges; S1 parallel S2 outside the shadow, Schweizer Tor, Ratikon.

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Fig. 6. Stretching lineations due to D1 (closed arrow heads) and D2. Each symbol is the averageof 3 to 15 measurements. Arrow heads point down-plunge.

4c);they correspond to shear bands (e.g. White et al. 1980). Criteria to distinguish rotational and irrotational strain will be described and discussed in a later section. Brittle minerals are commonly cracked and pulled apart.Open cracks are filled with debris, chlorite, calcite andquartz fibres (Fig. 4e). No foliation is visible in most competent rocks such as the Hauptdolomit. A subhorizontalstretchinglineation,L1, lies in S and trends WSW-ENE to NW-SE (Figs 5a and6).The lineation is defined by oriented minerals, elongated and/or pulled apart pebbles, minerals and fossils, and the long axes of asymmetric strain shadows (Fig. 4e). First generation folds, F1, are tight to isoclinal, recumbent, and their fold axes (Bl) have variable orientations toL1 (Fig. 5a stereonet, Fig. 6). Younging criteriaindicate that isoclinal folding occurred in many places. We infer mediumto large-scale folding with sheared-out hinges. The angle between F1 axial planes and the C-surfaces decreases as the angle between B1 and L1 decreases. F1 folding shows a variation in style approaching D1 shear zones (Fig. Sa): In moderately deformed regions, F1 folds are tight (interlimb angles 20-50”), west-vergent folds and fall into class 1C of Ramsay (1967). The folds become isoclinal in shear zones and the limbs are intensely stretched and boudinaged. The hinge lines rotate toward L1

and are frequently curved. Here, F1 folds resemble sheath folds (Cobbold & Quinquis 1980). Conjugate sets of shear fractures, several generations of extension veins, and bedding parallel veins form a complex vein pattern in calcareous rocks (Fig. 7a,b).Repeated crack-seal deformation(Ramsay 1980) led to growth of several generations of fibres parallel tothe incremental extension axes in extension veins. Pressure solution (Durney 1976) of calcite and quartz was the dominant deformation mechanism during vein development.This is indicated by truncated grain boundaries and fossils decorated with seams of residual insoluble materialin the S-planes. Nearly all extension veins containmineral fibres oriented parallel to L1. In incompetent rocks, the veins may be folded indicating early formation during the D1 event. The axial planes of such folds are parallel to the external S-foliation. Massive, competent rocks like ultramafic and crystalline basement blocks of the Arosa Zone are thoroughly fractured. Abundant chlorite andserpentine fibres and synthetic Riedel shears (Petit 1987) constrain the displacements on single faults. The E-W to SE-NW trending extension directions obtained by microtectonic analysis (see examples in Fig. 8a) correspond to the trend of L1 in the ductilely deformed rocks. D1 is responsible for the melange formation in the Arosa

Fig. 5. (a) Composite diagram of D1 structures; left, variationof structural style with respect to shear zones. Right, t’/lw-plots (Ramsay1967) for folds outside(top) and within zones of high strain (bottom). Insets: stereographic projection (lower hemisphere) of F1 fold axes (stars), L1 (dots) and intersection lineation between beddingand S1 (squares). (b) Composite diagram of D2 structures; Left, variationof structural style toward tectonically deep, high strained D2 zones in the Arosa Zone and Middle Penninic nappes. Right, t’/cu-plotfor F2 folds from upper (top) and lower (bottom) partsof the pile. Symbols asin (a).

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Zone. S1 planes are refolded at the millimetre to kilometre scale. A spaced second cleavage, S2, crenulates S1 or runs parallel to it. There is a marked increase in intensity of S2 toward deeper parts of the pile. In theupperparts, especially in competent rock units (radiolarite), S2 is completely missing or is developed as a fanning,spaced cleavage in upright F2 hinges (Fig. 5b). Pencil structures result from S1-S2 interference. The angle between S2 and S1 as well as the spacing of S2 decrease as strain increases toward deeper levels (Fig. 5b). In theArosaZoneand especially in theTasnaZone of the LowerEngadine Window (Fig. lb) S2 is a penetrative crenulation cleavage. S2 and S1 are indistinguishable in D2 shear zones. A subhorizontalN-to NE-trending stretchinglineation, L2, accompanies S2 (Figs 5band 6). It is expressed by elongatedand orientated minerals,pressureshadows on pyrite (Fig. 4g), and calcite and quartz fibres. These fibres

Fe.7. (a) Complex, closelyspaced vein pattern in Verspalaflysch,Ratikon; (b) Typical, fibre filled vein system used for incremental strain analysis; calcareous tlysch sandstone, west of Arosa.

grew in extension veins which crosscut veins with E-W trending fibres. In somecases a continuous growth sequence from fibres oriented parallel to L1 to those parallel to L2 is visible: early fibres trend, in general, E-W and curve, rather abruptly but still continuously into a N-S orientation (Figs 9a, b). Pyrite never has pressure shadows inthe L1 direction and therefore must have grown during the D2 event. F2 shows a variation in morphology and axial trend with respect to zones of increasing shear strain (Fig. 5b). In low strain zones or high levels, there is a distinct maximum of subhorizontal B2 approximatelyperpendicular to L2. The folds are open folds (interlimb angles 50-130") and .close to parallel style (Ramsay 1967). They may beupright or slightly inclined towards the north. Approachingzones of high D2 strain toward deep units, B2 rotates toward L2, fold outlinesbecomeasymmetric and tight(interlimb angles 20-50") and tend to be moresimilar in style.

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strain Incremental strain The first mineral fibres which appear in extension veins imply WSW-ENE trendingextension andare only well developed in the Austroalpine unit (Fig. 9a, b). The Ratikon area shows some early N-S to NW-SE oriented fibres (Fig. 8a). E-W to SE-NW oriented fibres follow in time;they arethedominant fibre direction andare well developed in the Arosa Zone and in the Austroalpine unit, but are less dominant in the underlying Middle Penninic nappes. After formation of these fibres, the extension direction changed to N-S or NE-SW. Fibres parallel to these directions are well developed in the Arosa Zone and the Middle Penninic nappes and are scarce in the Austroalpine unit (Fig. sa). A plot of average incremental strain intensity against average incremental orientations of all extension vein data shows the correspondence of L1 and L2 with the orientation maxima (Fig. 9a, b).

Finite strain

Fig.8. (a) Structural map of Ratikon (see Fig. 2), stretching lineations (lines), finite strains (ellipses) and incremental strains ( X Y sections, arrows points toward younger increments) and shear criteria (bold arrows, heads pointin shear direction). Note inversion of D1 related shear directionby subsequent folding. Stereoplots give compression and extension fields (latter shaded) from fault plane analysis (X,best-fit extension direction).(b) Fold axes (B) on map (arrows give plunge) and stereoplots of S1 and B (equal area, lower hemisphere).

Nonpenetrative structures Open third generation folds are superimposed onD2 structures. Their axes scatteraround N-S to NW-SE directions. Their wavelengths range from 1m up to 200 m. Locally, third generation fibres trend E-W to NE-SW. A fourthdeformation event(D4) causes locally developed, large scale, upright, gentle folds. B4 weakly clusters around the E-W direction. E-W and N-S trending strike-slip faults in the Ratikon area (Fig. 8) cut all other structures.

Pebble strains, reflecting ductile strain, were compared with strain data obtained from extension veins. Fig. 9c illustrates the corresponding shape fabrics in the XY plane of finite strain. In the tectonically upper parts of the Arosa Zone, L1 nearly parallels the X axis of the finite strain ellipsoid and S1 is subparallel to the X Y plane (Fig. 9c). Strain intensity is low and its geometry is close to plane strain (Fig. 10a); these strains are interpreted to be due to D1 only. Toward deeper parts of the Arosa Zone, X strain axes occupy intermediate positions between the directions of L1 and L2 (Fig. 9c), but with a maximum still around L1. In addition, strain geometry becomes more prolate or oblate (Fig. 10a).We interpret this by the superposition of D2 upon D1. D2 with an approximately N-S trending maximum stretching overprints D1 with an E-W trending maximum stretching. We interpret the finite straingeometries using a model similar to that of Ratschbacher & Oertel (1987). S2 crenulates S1 in F2 hinges (Fig. lob). Taking the foliations as X Y planes, the nearly orthogonal stretching lineations L1 and L2 as X directions of the related strains, and D1 strain intensity as always exceeding that of D2, the resultant strain isin the prolate field (Fig. lob, path 2 of Fig. 10a). The situation is different in the limbs of F2 where S2 roughly parallels S1. In this case, the orientation of the least extension is nearly parallel during both deformation events and thus a flattening geometry results (Fig. lob, path 1 of Fig, 10a). The oblate fabrics require a coaxial component in theD2 strain history, because in the case of progressive non-coaxial strain superposition with a common shear plane any change in the kinematic framework would lead to total plane strain. It is not feasible to assess possible volume changes during deformation. Tectonic consolidation may account for up to 50% reduction in volume in modernaccretionary wedges (Moore & Karig 1976), however most of our rocks must have been widely consolidated at the onset of D1 as shown by sharp, straightboundaries of extension veins which break across grains in the matrix(Fisher & Byrne 1987).

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unit circle

1-5 illustrates the progressive sequence

Fig. 9. (a) Incremental strain in Austroalpine and Penninic units at the western end of the Eastern Alps (Graubiinden-Ratikon; arrows: origin, sample location; head, toward younger increments; lengthof increments, proportional to stretch) read from the progressive sequence of vein formation (e.g. inset upper right) and graphical presentation of incremental strain changes with orientation (left inset, average of all plotted data). Theei maxima correspond toL1 and L2. (b) Incremental strain in the Lower Engadine Window; for explanation of symbols and diagram see (a). (c) Finite strain ( X Y projection, X 2 Y 2 Z principal axes); P , pebble strain measurements; E, data from extension veins. Inset: stereographic projectionsof L1, L2 and X of finite strain.

Deformation regime

A

b

path 1

Mesoscopic, west-translating D1 shear zones are locally developed (Fig. 4a). High strain is indicated by mesoscopic sheath folds, double-verging folds and rootless hinges with axes almost parallel to L1. We suggest dominant simple shear deformation for these shear zones. Mesoscopic S-C surfaces occur in moderately deformed rocks (Fig. 4b) andareinterpreted as the effects of a non-coaxial deformation history according to Berth6 et al. (1979). Further strain concentrations are expressed by shear bands. In general, only the synthetic set is developed (Fig. 4c, d) consistently indicating top W to NW shear. Asymmetric pinch-and-swell structures occur in competent layers, such as quartzitic and dolomitic layers in phyllites or carbonatic schists. The internal foliation of the swells curve atthe ends into single sets of asymmetrical extensional shears, oriented less than 40" to the external foliation (Sl). The sense of rotation indicated by the shear bands of the single swells (synthetic) is opposite to that (antithetic). The overall rotation is top to the west. Asymmetric strain shadows and recrystallized tails around clasts and minerals (Fig. 4e) also indicate noncoaxial Fig. 10. (a) Finite strain read from outlines of deformed objects plotted on a three axis diagram and factorized in two possible strain paths (curved arrows). (b) Interpretation of strain paths in terms of D1-D2 superposition. X1, Y1, Z1 and X2, Y2, 22, principal strain axes during D1 and D2,respectively. Assumptions: Foliations = XY planes, stretching lineations = X,L1 orthogonal to L2, D1 strain > D2 strain, irrotational deformation components; ( X > Y > 2, principal axes of incremental strains). In F2 hinges, where S1 has a high angle to S2, the superposition results in prolate strain geometry; on F2 limbs, where S1 and S2 are subparallel, flattening strain results.

845

ADRIATIC-EUROPEAN KINEMATICS, ALPS PLATE

deformation (Simpson & Schmid 1983). In the X Z sections, fibres in D1 extensionveinshavesigmoidal shapesand indicate rotation of the principalstrainaxes. Incontrast, competent, poorly or unfoliated rocks (Hauptdolomit) show straightfibres in the X Z sections. The lack of aplanar material anisotropy inhibited noncoaxial flow and favoured coaxial accumulation of strain (Lister & Williams 1983). Tectonicmelange formation occurred by (1) heterogeneouslayer-parallelextension of rheologically competent, oncecontinuouslayers(e.g.flysch-sandstones)producing angular fragments; (2) rigid body rotation of individual blocks with asenseofshear consistent with the overall rotation; (3) disruption by S-C surfaces and multiple sets of shear bands. Strain,indicatingoverallplanestrain geometryduring D1,combined with rotation suggests that simple shear contributed significantly to D1 in incompetent rocks. Hence, L1 appears to be the direction of motion (Mattauer et al. 1981). The sense of vorticity is generally top to the W to N W . Rotation is less clearlyexpressed in D2 thanD1 structures.Intheupperparts of thesequence,structures indicative of irrotationaldeformationarefrequent.Inthe X Z sections, fibres in D2 extension veins are straight and F2 folds are upright and of parallel style (Fig. 5b). The intensity deep levels with of F2 folding, however, increases toward high D2 strain. Occasionally, F2 folding is disrupted by back limb thrusts (Fig. 4f) which have N- to NE-directed displacement. S-C surfaces are developed in regions with strong S2 fabrics; the angle between the two S2 sets is always 230". Together with asymmetricstrainshadowsaround syn-D2 pyrites in the Couches Rouges of theSulzfluh nappe (Fig. 4g), they indicate the existence of rotational a component of deformation with N- NE-directed to displacement. Nearly coaxial deformation and low strain at high levels and anincrease of rotationaldeformationtogether with strain intensity toward deep levels are thus indicated. Shear is parallel to L2 and its sense is top to the N to NE.

Calcite textures Three incomplete X-ray calcite a = (2110) pole figures of D1(samples 1 from the Arosa Zone, sample 2 from the Sulzfluh nappe) and D2 (sample 3 from the Sulzfluh nappe) mylonites are shown in Fig. 11. All measurements (5 samples) suffer from poor statistics. Although the mylonites are fine grained, the grain size limit of the goniometer is approached. The plots show a-axis girdles,typical for low-T deformation of calcite polycrystals (Wenk et al. 1987; Schmid et al. 1987). The girdles are orientated subparallel to each other and they are consistently inclined (4-13") to the foliation plane ( X Y of finite strain). This suggests that the strain path has a non-coaxial component with a top W to N sense of shear, respectively (Schmid et al. 1981; Wenk et al. 1987).

Metamorphism and crystallization history

+

The assemblage epidote actinolite + chlorite is typical for greenstones of Klosters (Fig. 2, see also Peters 1963, Dietrich et al. 1974). Epidote and actinolite grew parallel to L1. Mica beards and phyllosilicate alignment in the S1 plane indicatethatrecystallizationandgrowth of phyllosilicates occurred during S1 development. Recrystallization of quartz as well as littlecrystallinity values (Ludin 1987) indicate temperatures of low grade metamorphic conditions. Further tothenorth,between Klostersand Tschagguns (Fig.2), quartz is notrecrystallized,however,subgrainformation and polygonization is common. Together with illite crystallinity values (Liidin1987),thisindicates very low grade metamorphism. Fluid inclusion studies suggest pressures of between 2.5-5.2 kbar in the area around Davos and Arosa (Fig. 2), and 2.3-4.2 kbar between Klosters and Tschagguns for metamorphism accompanying D1.Frey et al. 1980) estimated minimum temperatures of 250°C and minimum pressures of 2.3-4.1 kbar in the Middle Penninic Falknis nappeimmediatelybelowtheArosaZonesome 10 km WSW of Tschagguns (Fig. 2). Leimser &L Purtscheller

I

Fig. 11. Incomplete a = (2110) X-ray pole figures of carbonate mylonites; equal area projections. Contour intervals: 1,2: 1.0, 1.23, 1.46 m.r.d.; 3: 1.0, 1.18, 1.35m.r.d.. 1, 2: D1 mylonites, 3: D2 mylonite. Sense of rotational component in deformation indicated by half-arrows.

846

U . RING E T AL.

(1980) reported temperatures up to 350 “C and pressures of 4-5 kbar in the Lower Engadine Window. Incipient recrystallization and growth of phyllosilicates are found only whereD2deformation was intense,i.e. mainly toward deep parts of the Arosa Zone, and temperatures lower than thoseattained during D1are inferred. Furthermore, due to the growth and synkinematic recrystallization of the main parageneses along D1 structures, we conclude thatD1 marks thetemperature climax of metamorphism in the Arosa Zone.

Kinematic interpretation Thrusting processes occurred during D1 and D2. We interpret occasionally preservedearlierfabrics as an early stage of D1. We propose the following displacement history: (1) Thrustingcommenced in the highest unit, namely the Austroalpineunit.Motion was top to the WSW. The same event is also observed in some outcrops of the Arosa Zone, but the strain is small andmakes uponlya small percentage of the total strain. These early processes are not recognized in the underlying Penninic nappes. The significance of earliest N-S to NE-SW extensionsin the Ratikon area, whether local complexities or regional events, cannot be assessed. (2) During progressive deformation, thrusting was successively transferredto lower parts of the sequence (piggyback tectonics;Boyer & Elliott 1982). Shear was directed to the west to northwest. This event coincided with the highest amount of ductile deformation and the peak of regional metamorphism. We attribute most of the displacement in the Arosa Zone and in the Austroalpine to this event.The underlying MiddlePenninic nappes were incorporated in the W- to NW-directeddisplacement at a latestage of this deformation.Subsequentdeformation occurred under falling P-T conditions. D1 (3) Achange in shear directionoccurredbetween and D2. The incrementalstrain analysis demonstratesthe progressive history of D1 and D2 strains. Thetop N- to NE-directeddisplacementduring the D2 phase was more effective in lower levels of thethrust pile than in higher ones. We suggest that the Penninic nappes N and NW of the Arosa Zone accommodated most of their displacement of displacement in the during this episode.Theamount Arosa Zone was smaller during D2 than during D1. The late nonpenetrative phases are only of local importance. These phases occurred at very low temperature and may indicate uplift of an already thickened thrust pile.

Age of deformation Isotopic K-Ar data of syn- to post-tectonic alkali amphiboles from the Platta nappe (Fig. 2) yielded ages between 110-70 Ma (Deutsch 1983) and 90-70Ma(Philipp 1982). These data suggest an onset of D1 in Early Cretaceous time, most probablybefore the Albian.Preliminary data from Rb-Sr investigation on synkinematically formed, fine grained mica ( c 2 p m ) in a D1shearzone show ages of 2100Ma. Similar ageswereobtained fromthesame size fraction separated from weakly deformed country rocks see (Kralik et al. [l9871 for technique used). A D2 shear zone yielded ages of 45 Ma and 115Ma. The time offlysch sedimentation in theArosaZone ranges from Aptian to EarlyConiacian, i.e. c. 110-80Ma

(Dietrich 1976; Oberhauser 1983; Winkler & Bernoulli 1986; Winkler 1987; Liidin 1987; Harland et al. 1982). The flysch deposition is considered to reflect the active continental margin stage and thus indicates that imbrication was active since the later part of the Early Cretaceous. Hintsas tothe timing of the change from dominant W-directed to dominant N-directed motion come from the Middle Penninic Sulzfluh nappe in the Ratikon. Sedimentation there continued until the Palaeocene, as is demonstrated by the marly CouchesRougesdeposits.Palaeocene Couches Rouges were still affected by the W- to NW-directed elongation (arrows 1 and 2 in Fig. 9a). However, incremental strain data reveal only small amounts of strain. The change to S over Nmotion must have occurred after the Palaeocene in the Middle Penninic of the GraubiindenJRatikon area, as can be shown by both sedimentologic and radiometric data. Flysch of Palaeocene to lowermost Eocene age in the Lower Engadine Window (Oberhauser 1983) was still affected by W-directed elongations (arrows 1 to 6, Fig. 9b) and indicates thatthe change in the movement direction took place most probably less than 50 Ma ago.

Discussion Nature of the Arosa Zone An accretionary wedge has to be envisaged as a model for the formation of theArosaZone in which ‘offscraping’ from the hanging wall (the Austroalpine unit) and accretion of oceanic material (the Penninic nappe) occurred simultaneously. Imbrication started in the Austroalpine unit (Ratschbacher et al. 1987), then Austroalpine nappes (Silvretta nappeetc., Fig. lb) werebrought into contact with Penninic units. Downwardpropagationled tothe imbrication of Austroalpine and Penninic units, and finally to the imbrication of the Penninic Zone itself. Therefore, imbrication was kinematically dominated by the overriding plate.

Active continental margin Relative E over W motion initiated in the Early Cretaceous which is consistently supported by radiometric data and the sedimentary record. Radiometric ages of high-P metamorphic assemblages in the Eastern Alps are scarce, but data from the Austroalpine and South Penninic units of the Western Alps clusterin the range of 130-95 Ma (Oberhansli et al. 1985; Carpena 1985; Chopin & MoniC 1984; Hunziker 1974). The deposition of detritalhigh-P minerals in probably Turonian-Early Coniacian flysches of the Arosa Zone (Winkler & Bernoulli 1986) requires rapid uplift as well as an onset of subduction as early as the Lower Cretaceous. The palaeomagnetic data for the motion of the Adriatic plate (Dercourt et al. 1986), our kinematic analysis and sedimentological studies(Waibel & Frisch 1989)all show that partitioning of deformation into strike-slip and convergent motion along the Penninic-Austroalpine margin occurred during the Cretaceous and Lower Tertiary. Hesse (1981) and Hsii (1972) concluded from the duration of sedimentation in the Rhenodanubian and other flysch basins (Early Cretaceous toEocene) that motion between the Penninic and the Austroalpine domains was mainly lateral.

ADRIATIC-EUROPEAN KINEMATICS, ALPS PLATE

This inference is also confirmed by the study of Waibel & Frisch (1989) of the Tasna Series in the Lower Engadine Window. Here sedimentationcontinuously spans the time between Upper Liassic to Palaeocene and indicate proximity tothe Austroalpineunit. The palaeomagnetic, structural and sedimentologicalevidence for long-lasting, transpressional motions in the early history of the Eastern Alps is compelling. At least 20 km of E over W thrusting are proposed by Spitz & Dyhrenfurth (1914) and Schmid (1973) at the western end of theEastern Alps.We suggest the rotation of the Adriatic plate was the cause for this tangential motion,an ideaalreadyexpressed by Triimpy (1975), Frisch (1979), Laubscher & Bernoulli (1982), Platt (1986), and Ratschbacher et al. (1987). N(NE)-S(SW) compression (D2 event) began during the Eocene in the study area. Imbrication of the Middle Penninic nappes in theRatikonarea occurred afterthe deposition of the Palaeocene Couches Rouges and is mainly due to D2 dated around 45 Ma. The youngest fossils in the Lower Engadine Window are Lower Eoceneand were found in sediments portraying the highest rates of syndepositional deformation. Waibel & Frisch (1989) interpreted these sediments as reflecting the tectonic activity which led to the overriding of the Penninic units exposed in the Lower Engadine Window by the Austroalpine nappes during S over N motion. Eventually tectonic activity progressed totheforeland, namely theRhenodanubian flysch basin and the Helvetic realm (Frisch 1979).

Brittle versus ductile deformation Partitioning between brittle and ductile deformation on all scales is observed in the Arosa Zone. Dilatational fracturing in blocks of Austroalpine basment, flysch sandstones, ultramafic rocks and massive limestones is incontrast to ductile processes in shale, calcareous flysch and high strain shear zones. Therathersharp margins of the fibre-filled extension veins attestto a high degree of lithification of sediments the atonset of deformation. The ductile deformation was accomplished by a variety of mechanisms including dislocation creep (resulting in cataclastic fracture textures) and predominantly diffusional processes (mainly pressure solution). Growth of new phyllosilicate grains also contributed.Pressuresolutionsuppliedmaterials forthe fibrous fillings of dilatational veins. Vein formationand dilatational fracturing as a result of layer-parallel extension also contributed to the formation of zones of melange. The progressive disruptioninthesezonesportraysdifferent mechanical properties of various lithologies under very low and low grade metamorphic conditions encountered in the Arosa Zone.

Rotational versus irrotational deformation in accretionary wedges and melange genesis The west-directed stacking in the accretionary prism is responsible for the formation of zones of tectonic melange containing South Penninic and Austroalpine elements. Although some mixing is doubtlessly of sedimentary origin (e.g. olistoliths),melangeformation is dueto tectonic processes. This appears to be accomplished by development of anastomosing patterns of micro- to mesoscale shear zones forming web structures (Byrne 1984). These zones contain a variety of structures (‘shear bands’) indicating rotation and

847

extension subparallel to the sheardirection. In multilayered, anisotropic sequences (e.g. flysch, strongly foliated mylonites),thesestructuresrevealheterogeneousdeformation, which ultimately may haveresultedin a block-in-matrix structure (Needham 1987). The competence contrast, as in well-layered rocks of flysch deposits, promoted a strain regime which was locally complex. This phenomenon may be attributed to the process of flow partitioning. According to Lister & Williams (1983) incompetent layers with a strong planar anisotropy are able to take up much of the vorticity of bulk flow at high deformation rates, allowing the interleaving domains to deform nearly coaxially. Platt & Behrmann (1986) compared such a system to a deck of plastic cards in which slip occurs on the card surfaces. These surfaces represent the high strain zones in naturally deforming rocks. The cards themselves represent low strain areas which deform coaxially. We think the fragmentation induced both by partitioning of deformation in ductile and brittle processes and by interaction of sets of noncoaxial deformation structures, is responsible for melange formation in the Arosa Zone. The melange zones of the Arosa Zone may be classified astype I melanges according to Cowan (1985), however we stress progressive noncoaxial flow as being the main controlling factor for the bulk deformation.

Conclusions The overall deformationalong the Penninic/Austroalpine boundary zone is divided into discrete phases which caused distinct structures to form. The direction of tectonic transportchanged in response tothe changing external strain regime. The first event, starting no later than Aptian to Albian, was a WSW- to NW-directed thrusting within an accretionary prism. It was caused by westward thrusting of the Austroalpine unit onto the Penninic units. Thisevent essentially caused the melange structure of the Arosa Zone to form. It was progressively replaced by horizontal N-S to NE-SW compression in the Eocene. Our structural data combined with palaeomagnetic work support the following geotectonic model: thenorthern margin of the Adriatic plate had anirregular shape inherited from the Jurassic-Lower Cretaceous rifting stage.Anticlockwise rotation of the Adriatic plate during the Cretaceous and Lowermost Tertiary caused dextral strikeslip motion along the about E-W trending segments of the margin, and thrusting toward the west along its approximately NE-SW-oriented segments and within the Austroalpine and Penninic units. The Cretaceous to Lower Tertiary kinematics of the Eastern Alps is a history of transpression. During the Tertiary, the Adriaticplate moved northward and caused a change in the direction of motion to orthogonal compression which, from then on, governed the evolution of theeastern Alpine orogenic system. The change in external strain orientation during Alpine orogeny in theEastern Alps is recorded by sedimentological and structural data. We suggest that, in the early history of the Eastern Alps, a directrelationshipbetween plate motion and deformation at the plate margins existed (Ring et al. 1988). The formation of the Arosa Zone is attributed to ‘piggy back’ thrusting in an overall rotational deformation regime. Bothdeformationpartitioning in coaxial and noncoaxial, and in brittle and ductiledeformationcontributed tothe formation of melange zones.

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D. Bernoulli, H. J. Weissert, S. Schmid (ETH Zurich), J. Behrmann (Giessen), P. Liidin, W. Winkler (Basel), B. Stockhert (Bochum), S. Nievoll (Graz) and W. Obermiller (Tiibingen) contributed by discussions and/or technical assistance. H. J. Bunge made texture measurements at TU Clausthal possible. S. Borchert and T. Flottmann are thanked for a native speaker’s upgrading of the English text. U. Ringacknowledgesfunding by theState of Baden-Wiirttemberg and W. Frisch and L. Ratschbacher additionally by the Deutsche Forschungsgemeinschaft.

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Received 11 April 1988; revised typescript accepted 2 December 1988.