Ductile deformation and rheology of sub-continental mantle in a hot collisional orogeny: Example from the Bohemian Massif

Do

Va

b

c

a

ARRAA

KOPSCV

1

ilo(mlsc(

UT

0d

Journal of Geodynamics 56– 57 (2012) 108– 123

Contents lists available at ScienceDirect

Journal of Geodynamics

jo u r n al hom epage : ht tp : / /www.e lsev ier .com/ locate / jog

uctile deformation and rheology of sub-continental mantle in a hot collisionalrogeny: Example from the Bohemian Massif

ladimír Kusbacha,c,∗, Stanislav Ulrichb, Karel Schulmannc

Institute of Petrology and Structural Geology, Charles University in Prague, Albertov 6, 12843 Praha 2, Czech RepublicInstitute of Geophysics v.v.i., Academy of Sciences of the Czech Republic, Bocní II/1401, 14131 Praha 4, Czech RepublicInstitut de Physique du Globe de Strasbourg, IPGS – UMR 7516, CNRS et Universite de Strasbourg (EOST), 1 Rue Blessig, 67084 Strasbourg, France

r t i c l e i n f o

rticle history:eceived 30 January 2011eceived in revised form 6 June 2011ccepted 10 June 2011vailable online 23 July 2011

eywords:livine lattice preferred orientationeridotite mylonitetrainrust–mantle couplingariscan orogen

a b s t r a c t

Fabric patterns of strongly serpentinized peridotite were determined using eigenvector analysis andeigenvalue classification of lattice preferred orientation of olivine and orthopyroxene. This approachhas been applied to a rootless fold-shaped body of mylonitized spinel to garnet peridotite surroundedby fine-grained and partially retrogressed ky–kfs granulite. The EBSD data show either axial [0 1 0] or[1 0 0](0 k l) pattern, both characteristic for ‘dry’ slip systems. The former pattern occurs predominantlyalong the inner margin and southern limb while the latter is mainly developed in the hinge of the foldshaped body. Foliations and lineations deciphered from the LPO data suggest that the [1 0 0](0 k l) patternreflects constrictional deformation (prolate strain ellipsoid) in the hinge of the peridotite fold while theaxial [0 1 0] pattern reflects pure flattening (oblate strain ellipsoid) inherited from the period of emplace-ment of the peridotite sheet in the crust. Similarity in finite strain pattern of peridotite and surroundinggranulites indicates their common thermal and mechanical evolution during folding. The petrology andstructural data result in a model of burial of peridotite below thickened crustal root, its exhumationand folding. The burial stage is associated with prograde metamorphism resulting in a coarse-grainedmicrostructure and development of spinel and garnet zones. The emplacement of peridotite into lower
crustal granulites occurred along a shear zone associated with grain size reduction in both peridotite andgranulite and rapid cooling of mylonitized peridotite to the ambient temperatures of lower crust. Furtherascent to mid-crustal levels occurred within the vertical granulite channel. Final fold shape of the peri-dotite developed during subsequent indentation of the weak vertically anisotropic crust by the adjacentcontinental promontory. The degree of mechanical coupling between folded peridotite and granulite inmid-crustal levels is estimated using comparison of studied microstructures with experimental data.
. Introduction

Continental lithosphere consists of different layers of contrast-ng rheology: brittle upper continental crust, weak and ductileower continental crust and stronger sub-crustal mantle, with rhe-logy evolving with temperature, pressure, depth and compositionBrace and Kohlstedt, 1980; Ranalli and Murphy, 1987). Rheological

odels of the lithosphere are based on experimental rheologicalaws describing brittle and ductile behavior of rocks. Relative
trength of the different layers is evolving through time due tohanges in tectonic style and thermal evolution of the lithosphereThompson et al., 2001). Lithosphere rheology models can be
∗ Corresponding author at: Institute of Petrology and Structural Geology, Charlesniversity in Prague, Albertov 6, 12843 Praha 2, Czech Republic.el.: +420 603181316; fax: +420 22195 1533.

E-mail address: [email protected] (V. Kusbach).

264-3707/$ – see front matter © 2011 Elsevier Ltd. All rights reserved.oi:10.1016/j.jog.2011.06.004

© 2011 Elsevier Ltd. All rights reserved.

validated using depth of earthquakes hypocentres, which shouldbe in accordance with the thermal structure of the lithosphere andassumed the thickness and depth position of lithological layers.Earthquakes hypocentres located in the upper crust and in theuppermost part of lithospheric mantle indicate that these layerscan be relatively strong and brittle (Liang et al., 2008; Molnar andChen, 1983). Nevertheless, earthquakes located in the uppermostmantle, where Byerlee’s law cannot be extrapolated, are rare andtheir source mechanism is still debated (Monsalve et al., 2009).These observations are commonly interpreted using a “jelly sand-wich” model in which a weak lower crust is sandwiched betweenstrong elastic-brittle upper crust and an elastic-ductile uppermantle (Burov and Watts, 2006). Other seismological studies showan absence of earthquakes in the upper mantle underneath thick
orogens, which is interpreted using so called “crème brûlée” modelmade of strong seismogenic crust and a weak and “wet” uppermantle (Jackson, 2002). Aseismic behavior of upper mantle is alsosupported by presence of ductile shear zones recorded in spinel
dx.doi.org/10.1016/j.jog.2011.06.004

http://www.sciencedirect.com/science/journal/02643707

http://www.elsevier.com/locate/jog

mailto:[email protected]

dx.doi.org/10.1016/j.jog.2011.06.004

eodyn

p22

akgbpaadicTflsfTlrtd11

psSccaicp(soipa2oo2naWedabS

fHscmab

2

o

V. Kusbach et al. / Journal of G

eridotite massifs and xenolites (Drury et al., 1991; Tommasi et al.,000) deformed at depth greater than 30–40 km (Dijkstra et al.,004).

Deformation of the lithosphere and mechanical coupling of crustnd mantle are not only dependent on rheology but also on theinematic regime. For convergent plate boundaries there are twoeneral models of mechanical behavior of the crust and the mantleased on structural observations and mantle anisotropy geo-hysical studies: (1) vertical deformation coupling between crustnd mantle expressed by steep deformation fabrics in the crustnd underlying mantle associated with transcurrent/tranpressiveeformations (Vauchez et al., 1998), (2) clutch tectonic model

mplying transfer of ductile deformation from the mantle to therust along a weak lower crustal zone (Tikoff et al., 2002, 2004).he latter model is supported by a concept of weak lower crustowing over strong and mechanically passive upper mantle pre-erving “frozen” fabric from the period of the mantle lithosphereormation (Babuska and Plomerová, 2006; Babuska et al., 2008).he degree of coupling between the crustal and mantle parts of theithosphere is also a function of the balance among surface forceselated to plate tectonics, the gravity force related to lateral varia-ions in lithospheric thickness, and the buoyancy forces related toensity variations (Artyushkov, 1973; Chemenda et al., 1995; Ellis,996; Molnar and Lyon-Caen, 1988; Molnar et al., 1993; Ramberg,981).

There exist only rare structural and microstructural data fromeridotites allowing direct assessment of rheological behavior ofub-continental mantle (Precigout et al., 2007; Roux et al., 2008;oustelle et al., 2009) recorded during continental rifting andrustal thinning. However, there is no information about rheologi-al behavior of subcontinental mantle during continental collisionnd crustal thickening. The Bohemian Massif represents such annternal orogenic zone of double thickened Variscan orogen whichontains the high grade lower crustal Gföhl Unit marked by theresence of garnet and spinel peridotite bodies of different originMedaris et al., 1990) enclosed within large complexes of high pres-ure granulites (Carswell and O’Brien, 1993). Large exposures ofrogenic lower crust and upper mantle allow studying mechanicalnteractions between these two key lithologies at elevated tem-eratures and pressures corresponding to actual crustal thicknessnd temperature of the Tibetean or Andean orogens (Guy et al.,011; Lexa et al., 2011). However, up to the present time, there arenly few studies concerning internal deformation fabric of somef large peridotite bodies in the Bohemian Massif (Kamei et al.,010; Machek et al., 2009; Medaris et al., 2009) compared to aumber of structural and microstructural studies of HP granulitesnd to studies of similar lower crustal complexes (e.g. Caledonianestern Gneiss Region in Norway; Brueckner, 1998; Brueckner

t al., 2002). Lack of such studies in peridotites is due to highegree of serpentinisation of these rocks allowing only petrologicalnd geochemical analyses of few well preserved mineral assem-lages (Medaris et al., 1990, 2005; Schmädicke and Evans, 1997;chmädicke et al., 2010).

This work attempts to combine structural and microfabric datarom a relatively well preserved peridotite body and surroundingP granulites with existing petrological and geochronological con-

traints to propose a consistent model of rheological behavior ofrust and mantle in a deep orogenic root. Apart from the conceptualechanistic model these dataset results in the semi-quantitative

ssessment of the degree of rheological and mechanical couplingetween crust and mantle during orogeny.

. Geological setting

The Bohemian Massif represents the easternmost exposuref the European Variscan belt (Fig. 1a) (Edel and Weber, 1995).

amics 56– 57 (2012) 108– 123 109

The present day structure of the Bohemian Massif originated bythe southeastward subduction of the Saxothuringian OrdovicianOcean underneath the Tepla-Barrandian domain continentallithosphere during Upper Devonian (Franke, 2000). Closure ofthe Saxothuringian Ocean followed by underthrusting of theSaxothuringian lithosphere beneath the Tepla-Barrandian domainled to development of a double-thickened orogenic root rep-resented by the upper crustal Tepla-Barrandian domain andmiddle/lower crustal Moldanubian domain (Schulmann et al.,2009) (Fig. 1a). Thickening of the orogenic root domain wasfollowed during Lower Carboniferous by the opposite indentationof the continental promontory of the Brunia continent, triggeringextrusion and exhumation of the orogenic lower crust into themid-crustal depths (Schulmann et al., 2005). The studied area islocated in the Gföhl Unit of Moldanubian domain representingoutcrops of this orogenic lower crust at the eastern margin of theBohemian massif (Fig. 1b).

The Gföhl Unit consists of partially molten orthogneisses,amphibolites and migmatite gneisses including bodies of gran-ulites, eclogites and peridotites (Tollmann, 1982). Felsic granulitesof granite/rhyolite affinity (Fiala et al., 1987) are composedof quartz, garnet, kyanite, alkaline feldspar, plagioclase andrutile (Carswell and O’Brien, 1993). Typical felsic granulites areinterpreted as metamorphosed equivalents of Lower Palaeozoic,probably high-level granites underthrust and buried to form thedeepest part of the continental root during the Variscan collision(Janousek et al., 2004; Janousek and Holub, 2007; Lexa et al., 2011).Petrological studies of felsic and pyroxene bearing intermediategranulites show peak conditions of 800–1000 ◦C and 16–20 kbar(Cooke, 2000; Cooke and O’Brien, 2001; Racek et al., 2006, 2008;Stípská et al., 2004; Tajcmanová et al., 2006) and amphibolitesfacies retrogression at 750–700 ◦C and 9–4 kbar. The process of ver-tical movement of high-grade granulite complexes to mid crustallevels has been recently described and attributed to a vertical extru-sion of orogenic lower crust (Schulmann et al., 2005; Tajcmanováet al., 2006). The extrusion was later followed by horizontal chan-nel flow (Schulmann et al., 2008; Stípská et al., 2008) in middlecrustal level, coupled with retrogression of granulites and intensivemelting (Hasalová et al., 2008a).

The studied granulite–peridotite rock association occurs withinthe Námest’ Granulite Massif (Matejovská, 1967) at the bound-ary with the continental Brunia promontory to the east and tothe north (Fig. 1b and c). The granulite body is nearly com-pletely enveloped by a belt of pyroxene and/or garnet amphiboliteswith a tholeiitic geochemical signature (Sichtarová, 1981). Thisgranulite–amphibolite complex is further surrounded by garnetand sillimanite bearing migmatitic orthogneiss (Matejovská, 1975).Hasalová et al. (2008a,b) showed that the different orthogneisstypes can be considered as a continuous sequence ranging frombanded orthogneiss to nebulitic migmatites, developed by meltinfiltration in a middle-crustal channel.

The Mohelno peridotite belongs to numerous bodies of spinelto garnet peridotites enclosed within retrogressed granulite ofthe Námest’ Granulite Massif. These bodies form lenses rangingfrom meters to several kilometers in size (Urban, 1992), while theMohelno serpentinized peridotite exhibits a form of a large foldarc with decreasing thickness from ∼1000 m in the northern limbto ∼300 m in the southern one (Fig. 1c). Peridotites are stronglyserpentinized (50–100%) but their original pristine microstruc-ture is locally preserved. The original mineral assemblage of theperidotites is represented by olivine, orthopyroxene, clinopyrox-ene and spinel (Stage I of Medaris et al., 1990). Primary spinel
is overgrown by garnet (Stage II) followed by further breakdownto clinopyroxene and spinel (Stage III) and amphibole and spinelsymplectites (Stage IV) associated with fine-grained myloniticmicrostructure. Peridotites display a geochemical signature

110 V. Kusbach et al. / Journal of Geodynamics 56– 57 (2012) 108– 123

F the Eut udy a

cep(eAap1takfi1dihz1

ig. 1. Geological setting: (a) position of the Bohemian Massif in the framework of

he eastern margin of the Bohemian Massif; (c) a simplified geological map of the st

haracteristic for an asthenospheric origin, and because of pres-nce of pyroxenite veins, it was proposed that the original spineleridotite developed at temperatures close to dry pyrolite solidusKamei et al., 2010; Medaris et al., 1990). However, pressurestimates for this evolutionary stage have not been quantified.ccording to standard geothermobarometry the peak mineralssemblage of garnet peridotite occurred at 2.3–2.8 GPa and tem-erature ∼1200 ◦C (Medaris et al., 1990) or at 1.8–2.5 GPa at100–1250 ◦C (Kamei et al., 2010) while mylonitic microstruc-ure with secondary spinel reveals conditions of 0.8–1.5 GPat 800–950 ◦C. The clinopyroxene-spinel and amphibole-spinelelyphite assemblage around garnets are correlated with granuliteacies and amphibolite facies retrogression of the surround-ng granulites, respectively (Kamei et al., 2010; Medaris et al.,990, 2005). The contact zone between granulite and peri-otite along the inner margin of the Mohelno peridotite body

s characterized by the presence of various magmatic rocks asornblendite, biotitite and gabbrodiorite interpreted as a reactionone between hot peridotite and host granulites (Dobretsov et al.,984).

ropean Variscides (modified after Franke, 2000) and (b) simplified tectonic map ofrea according to the geological map 1:50,000, sheet of the Czech Geological Survey.

3. Structural geology of Mohelno peridotite and host rocks

The structural pattern of the Mohelno peridotite revealshomogeneous mylonitic fabric and pyroxenite layering whilethe surrounding granulite and amphibolite exhibit a polyphasestructural evolution (Fig. 2). Special attention was paid to the dis-tribution of garnet and coarse spinel within the Mohelno PeridotiteBody (MPB). Our study shows that the coarse-grained spinel varietyoccupies the majority of the peridotite body, while the garnet one isrimming the inner arc of the fold hinge with exception of one sam-ple located in the southeastern extremity of the peridotite body(Fig. 3e). In general, the peridotite mineral assemblage is rarelypreserved due to extensive serpentinization. Both coarse-grainedspinel and garnet bearing peridotite are mylonitized resulting indevelopment of fine-grained spinel bearing matrix and orthopy-roxene porphyroclasts (Fig. 3a). This mylonitic foliation is generally
steeply dipping and its orientation follows the fold shape of theperidotite body with the fold hinge steeply plunging to the West(Fig. 3e). The second fabric element is represented by milimetres toseveral centimetres thick bands of websterites to orthopyroxenites

V. Kusbach et al. / Journal of Geodynamics 56– 57 (2012) 108– 123 111

F study

g uctured erthr

(pti

GSeriappfi(oEGt

ig. 2. Structures in granulites observed in the field: (a) the structural map of the

rained foliations (S2–3) and lineations (L2–3) in granulites as well as S4 and L4 stromain preserving steep S2 mylonitic foliations affected by S4 flat fabric due to und

Fig. 3a) that form layers either parallel or oblique with respect torincipal mylonitic fabric. The pyroxenite layers are also myloni-ized and affected by late fractures which progressively disappearn the surrounding serpentinized matrix.

Four deformation fabrics were recognized within the Námest’ranulite Massif directly surrounding the Mohelno peridotite. Both1 and S2 fabrics are characterized by a granulite facies min-ral assemblage such as garnet, kyanite, perthitic feldspar andutile. Coarse-grained granulite facies S1 fabric is preserved onlyn low-strain domain within the internal part of the fold hingend is defined either by lithological layering or by a faint shapereferred orientation of mineral aggregates. Due to strong D2 trans-osition the orientation of S1 fabric cannot be determined in theeld. Reworking of the S1 fabric is correlated with evolution ofultra)mylonitic granulite S2 fabric associated with development
f isoclinal folds F2 (Fig. 3b). S2 foliation is striking N–S and–W in the northern and in the southern parts of the Námest’ranulite Massif, respectively (Fig. 2a). Here, the subvertical S2 is
urned together with the peridotite layer and bear mineral lineation

area shows trajectories of S2–3 foliation, contoured pole figures of mylonitic fine-s; (b) the idealized cross-section shows a Námest’ Granulite Massif as a low strain

usting of the Brunia basement promontory.

moderately plunging to the West (Fig. 2a). The rotation of theS2 foliation into the E–W direction is associated with the largescale F3 folding of the granulite layering and peridotite sheetresulting in formation of upright peridotite F3 folds with E–Wtrending axial plane and west plunging hinge. The large scale fold-ing is accompanied with a retrogression E–W trending S2 fabricsunder amphibolite facies conditions associated with partial melt-ing forming foliation parallel leucosomes. The retrogression alongfold limbs leads to development of steep composite S2–3 fabricmarked by alternations of granulites with migmatitic granuliticgneiss (Matejovská, 1975). The internal part of the peridotite foldhinge is affected by D3 shear zones which may be interpretedas axial cleavage of large scale F3 folds. Here, the granulites arestrongly sheared and retrogressed as shown by a syn-D3 growth ofsillimanite and biotite and elongated lenses of granitic melt. Far-
ther away from the peridotite body is D3 expressed also by thedevelopment of meter scale parasitic F3 folds.
Sub-vertical S2 and S3 fabrics are reworked by sub-horizontalS4 foliation, which is a common structure of the Námest’ Granulite


Fig. 3. Main structural elements in the close vicinity of the Mohelno peridotite body observed in the field. (a) Mylonitic foliation in peridotite is defined by elongatedporphyroclasts of orthopyroxenes and rework also websterite layers that occur parallel or oblique to the mylonitic foliation; (b) mylonitic fine-grained foliation S2 ingranulites show often isoclinal folds; (c) a granulite–peridotite interface along the inner fold with felsic partial melts and granulite wedging towards the peridotite; (d)localized flat S4 shear zones in granulites rework suvbertical S2–3 foliation; (e) a simplified geological map of the study area shows trajectories of the mylonitic foliation andwebsterite layering in peridotites, and S2–3 foliation in granulites. Occurrences of garnet (black squares) and sampling sites (black circles) are shown as well as foliationsm ation,

MNfftmiapSba

easured in the field; (f) contoured pole figures of the poles to porphyroclastic foli

assif. In the volumetrically more important western part of theámest’ Granulite Massif, the N–S and E–W striking steep S2–3

oliations are reworked by flat shear zones and recumbent openolds with N–S and E–W trending hinges, respectively (Fig. 3d). Inhe eastern extremity of the massif the degree of the D4 defor-

ation is more important leading to development of close tosoclinal E–W trending F4 folds associated with development of

pervasive S4 migmatitic foliation. Here, the S4 is marked by
referred orientation of biotite and coarse leucosome layers. The4 bears SW plunging mineral and stretching lineation markedy alignment of sillimanite and elongation of quartz-feldsparggregates (Fig. 2a).
websterite layers and S2 together with S3 foliations.

4. Peridotite and granulite microstructures

According to experimental works, deformation microstructuresand grain size depend on temperature, stress and/or strain rateand can be used to decipher prevailing deformation mechanisms(Rutter and Brodie, 1992). These data may contribute to understandmechanical behavior of mantle and crust during different stages oftectonic evolution. In peridotite, grain size was measured in less
serpentinized domains using linear-intercept method. Individualmicrostructures of granulites were manually traced and mediangrain size was estimated using area-intercept method for individ-ual phases. It is noteworthy that serpentinization alters 50–100% of


Fig. 4. Microphotographs of the most important microstructures: (a) a coarse-grained opx + spl + ol porphyroclastic domain characterises the original mineral association int ned ms re in

m

tcFa

sdso(tafdm(Ivmts2

eeFdo

cqd(mKw

he majority of the peridotite body; (b) elongated garnet replace spinel in coarse-graiurrounded by serpentinized mylonitic microstructure; (d) coarse S1 microstructuicrostructure in granulite (mineral abbreviations after Kretz, 1983).

he rock volume and is expressed by presence of antigorite bandsoloured with a dispersed submicroscopic pigment of iron oxides.erromagnetic minerals, mainly magnetite and maghemite are alsobundant, and chlorite and talc are present as accessories.

The original coarse-grain microstructure is reported for both,pinel and garnet peridotite relics, preserved in less serpentinisedomains (Kamei et al., 2010; Medaris et al., 2005). Coarse-grainedpinel peridotite relics contain large orthopyroxene (1–15 mm),livine (1–4 mm), clinopyroxene (1–2 mm) and coarse spinel1–2 mm) (Fig. 4a). Orthopyroxene porphyroclasts locally showhin exsolution lamellae of clinopyroxene along mineral cleav-ge. In the coarse-grained garnet peridotite the garnet occurs inorm of large and elongated (up to several centimetres in size oniameter) crystals (Fig. 4b and c) surrounded by coarse-grainedatrix preserving stable grain boundaries with orthopyroxene

1–4 mm), olivine (1–3 mm) and clinopyroxene (1–2 mm) (Fig. 4b).n the whole MPB the coarse-grained original microstructure is con-erted into dynamically recrystallized fine-grained spinel-bearingylonitic matrix (mean grain size ∼60 �m). Many garnets found in

he fine-grained spinel-bearing matrix show inner clinopyroxene-pinel and outer amphibole-spinel kelyphytic rims (Kamei et al.,010; Medaris et al., 2005).

Millimeter to several centimeter-thick pyroxenites bandsxhibit a porphyroclastic microstructure defined by orthopyrox-ne porphyroclasts surrounded by polygonal orthopyroxene grains.ine-grained clinopyroxene is occasionally preserved withinynamically recrystallized orthopyroxenes with mean grain sizef ∼60 �m.

Rare granulite S1 microstructure is characterized byoarse-grained mineral assemblage (several mm in size) ofuartz + perthite + garnet + kyanite + biotite. The S1 fabric isefined by elongation of garnet grains and quartz ribbons
Fig. 4d). Perthitic feldspars (∼200 �m) show core and mantle
icrostructure rimmed by neoblasts of plagioclase (∼35 �m) and-feldspar (∼30 �m; Fig. 5). Large quartz grains (∼200 �m) occurithin platy ribbons and exhibit highly lobated grain boundaries

icrostructure along the inner margin of the body; (c) garnet-pyroxene porphyroclastgranulite, with core-mantle features; (e) ultra-mylonitic S2 microstructure; (f) S3

pointing to grain boundary migration recrystallization mechanism(Guillope and Poirier, 1979). Granulite from a close vicinity ofthe peridotite body show well developed mylonitic granuliteS2 fabric with stable mineral association of high pressure gran-ulites (qtz + kfs + plg + grt + ky + rt + bt). Recrystallized feldspars aswell as quartz are intermixed within equal-sized fine-grainedrecrystallized microstructure of quartz (∼59 �m), plagioclase(∼50 �m) and K-feldspar (∼43 �m) (Fig. 4e, Fig. 5). Garnet crystalsare rich in large quartz and perthitic feldspar inclusions withlobate grain boundaries. These inclusions reveal larger grain sizecompared to the fine-grained matrix. Similarly, kyanite crystalsform inclusions inside rims of coarser perthitic feldspars pointingout to their early growth. Retrogression of S2 by D3 deforma-tion is generally characterized by fabric coarsening and morespecifically by growth of plagioclase around unstable kyanite,formation of coarse-grained quartz ribbons and leucosome layersparallel to the S2 fabric. The S4 fabric is associated with a for-mation of amphibolite facies mineral assemblage represented byqtz + kfs + plg + bt ± grt ± ky ± sill ± ms ± rt ± ilm. Retrogression ismostly characterized by heterogeneous grain coarsening, hydra-tion of original mineral assemblage marked by development ofbiotite rich bands and leucosome layers parallel to S4 foliation(Fig. 4f), and growth of a new fine-grained garnet as reportedalso by Hasalová et al. (2008a). Kyanite is replaced by sillimaniteforming acicular aggregates or fibrous crystals aligned with biotiteschlierens replacing garnet. Locally, muscovite formed at theexpense of biotite. Both feldspars form coarser-grained matrixcomposed of K-feldspar (∼151 �m) and plagioclase (∼103 �m)that enclose thick ribbons consisting of coarse quartz grains(∼122 �m) with lobated grain boundaries (Figs. 4f and 5).

5. Lattice preferred orientation and deformation
mechanisms
LPO of olivine and orthopyroxene in peridotite allows deter-mining the orientation of the main fabric preserved from


F icrostrs

str2RP

5

nabNweicpsoimcbom

sa(t(strptt

ig. 5. Histograms showing grain size distribution of quartz and plagioclase in S1 mtatistical data in the top right corner.

erpentinization. It also provide information on active slip sys-ems of olivine and orthopyroxene that depend on (a) deformationegime (Tommasi et al., 1999), (b) temperature (Demouchy et al.,009; Tommasi et al., 2000), (c) pressure (Couvy et al., 2004;aterron et al., 2009) and probably (d) water content (Chopra andaterson, 1984; Jung and Karato, 2001; Mackwell et al., 1985).

.1. Methods

Lattice preferred orientation of olivine was measured on a scan-ing electron microscope CamScan4 at the Institute of Petrologynd Structural Geology, Charles University in Prague by electronack-scattered diffraction (EBSD) using HKL technologies CHAN-EL 5 software (Schmidt and Olesen, 1989). Diffraction patternsere acquired at a working distance of 40 mm and using an accel-

rating voltage of 17 kV. The whole procedure (pattern acquisition,mage freezing, band detection, indexing and result backup) wasarried out on the studied samples manually. Additionally latticereferred orientation of olivine and orthopyroxene were mea-ured on a TESCAN scanning microscope at the EOST, Universityf Strasbourg using EDAX OIM software. LPO data were obtainedn automatic mode along with chemical elements distribution to

inimize mis-indexing. Measured raw data were manually pro-essed to orientation data sets based on one point per grain. Inoth approaches each individual grain is represented by only onerientation measurement, and equal weighting to each orientationeasurement has been given for the contouring of the pole figures.Due to strong serpentinization, the LPO was measured in thin

ections prepared in three ways: (a) normal to porphyroclastic foli-tion and parallel to stretching lineation both defined in the field“matching F&L” samples in Fig. 6), (b) normal to porphyroclas-ic foliation defined in the field and parallel to its dip direction“matching F” samples in Fig. 6), (c) in case of strongly serpentinizedamples, normal to a randomly chosen fracture plane and parallelo its dip direction (“not matching” samples in Fig. 6). All LPOs were
otated and presented in non-polar, lower hemisphere equal arearojections in the geographic coordinates (North is located in theop of the pole figure). Foliation plane and lineation measured inhe field are presented in the pole figure as solid line and triangle,
ucture, and quartz, plagioclase and K-feldspar in S2/S3 microstructure. Tables with

respectively. The different patterns of olivine LPOs are illustratedin three pole figures of the main axes [1 0 0], [0 1 0] and [0 0 1].

In order to characterize the occurrence of the given type ofolivine LPO in a more objective way, a PGR diagram was car-ried out according to Vollmer (1990). He proposed the eigenvalueclassification for orientation data in order to quantify the typeof distribution (point, girdle and random). The magnitude ofthree eigenvalues for every presented crystallographic direction(�1 ≥ �2 ≥ �3 with the normalization �1 + �2 + �3 = 1) are used todefine three fabric indices, point maximum (P = �1 − �2), girdle(G = 2(�2 − �3)) and random (R = 3�3). These indices range from 0to 1 and have the property that P + G + R = 1. Both analyses as wellas projection of pole figures have been carried out by software fromthe shareware package (Mainprice, 1990); ftp://www.gm.univ-montp2.fr/mainprice//CareWare Unicef Programs/).

Additionally strength of the fabric in every sample wasexpressed in misorientation index (M-index, Skemer et al., 2005).The M-index is based on the distribution of uncorrelated misorien-tation angles. Value of the M-index corresponds to the differencebetween the observed distribution of uncorrelated misorientationangles and that predicted for a random fabric. It has values from 0(for random LPO fabric) to 1 (for a single crystal fabric).

5.2. Results

Lattice preferred orientation of olivine is not uniform in themeasured samples. The eigenvalue classification showed a strongcomponent of random distribution in all samples, but also allowto identify two types of LPO patterns (Fig. 6). The first type showsgirdle distribution of the [1 0 0] axes with the sub-maximum insidethe girdle and broad point maximum of the [0 1 0] axes provid-ing the strongest maximum. The [0 0 1] axes show generally weakpreferred orientation, but some samples show a weak point maxi-mum (Fig. 6). In the PGR diagram the [1 0 0] axis plots in betweenG and R corners, indicating girdle dominated distribution of this
axis, while the [0 1 0] axis plots between P and R apexes suggest-ing a point type distribution (Fig. 8b). This LPO is characteristic forthe AG-type (Mainprice, 2007) or axial [0 1 0] pattern accordingto Tommasi et al. (2000), also called [0 1 0]-fiber pattern (Bunge,
ftp://www.gm.univ-montp2.fr/mainprice//CareWare_Unicef_Programs/


Fig. 6. Measured olivine LPO is presented in contoured pole figures of [1 0 0], [0 1 0] and [0 0 1] crystal directions in geographic coordinates. The left column shows LPOwith axial [0 1 0] pattern, while right column present LPO with [1 0 0](0 k l) pattern. ‘Matching F&L’ group of LPOs correspond to samples with clearly defined foliation andstretching lineation in the field. ‘Matching F’ (mylonitic foliation recognised) and ‘not matching’ (neither foliation nor lineation recognised) group show LPOs that belongto more intensely serpentinized samples. Full line and dashed line half-circle correspond to the trace of the mylonitic foliation measured in the field and defined from theeigenvector analysis, respectively. Full triangles and full circles show positions of the stretching lineation measured in the field and calculated from the eigenvector analysis,respectively. Pole figures are in equal area projection and lower hemisphere, and contoured at interval 0.5 times of uniform distribution. N is a number of measured grainsand P, G and R values are calculated from the eigenvalue analysis.


Fig. 7. LPO of olivine and orthopyroxene measured on ‘not matching’ samples show parallelism of the olivine [1 0 0] with orthopyroxene [0 0 1] maxima positions whichr es are1 easu

1fmlsttptKpa

fposvtofepb(tto

met

6

t

epresent independent control of the stretching lineation determination. Pole figur.0 times of uniform distribution in olivine and opx, respectively. N is a number of m

982). It is developed preferentially along the inner margin of theold-shaped serpentinized peridotite body. Locally, it has also been

easured in a few samples from the central part of the southernimb of the peridotite fold (Fig. 8a). The second LPO type shows thetrongest point maximum of [1 0 0] and girdle distribution for bothhe [0 1 0] and the weakest [0 0 1] axes (Fig. 6). In the PGR diagramhe [1 0 0] axes plot between P and R corners while the [0 1 0] axeslot in between G and R apexes. Such a type of texture correspondso [1 0 0](0 k l) pattern (Tommasi et al., 2000) or D-type (Jung andarato, 2001; Karato et al., 1980; Mainprice, 2007) or [1 0 0]-fiberattern (Bunge, 1982) and is typically developed in the interior andlong the outer parts of the fold shape peridotite body (Fig. 8).

In several sites the macroscopic lineation or both lineation andoliation were not known (Fig. 6, matching F and not matching sam-les) and therefore the eigenvalue analysis was used to identify therientation of these fabric elements. Because, the [1 0 0](0 1 0) slipystem is the most active in both LPO types the maximum eigen-ector orientations for the [1 0 0] and [0 1 0] axes were attributedo the stretching lineation and the pole of foliation, respectively. Inrder to confirm the orientation of lineation and foliation deducedrom olivine LPO the lattice preferred orientation of orthopyrox-ne was measured in several samples. Orthopyroxene LPO showsoint maximum in [1 0 0] and [0 1 0] direction, and girdle distri-ution of the [0 0 1] axes with the sub-maximum inside the girdleFig. 7) suggesting main activity of the [0 0 1](0 1 0) and minor con-ribution of the [0 0 1](1 0 0) slip systems. These pole figures showhat the orthopyroxene LPO confirms the foliation and lineationrientations deduced from olivine LPO measurements.

M-index calculation shows small strength of the olivine LPO inajority of samples providing values less than 0.1. Several samples

xhibit a higher LPO strength, however there is no correlation withhe type of the LPO (Fig. 8a).

. Discussion

Data presented in this work show the procedure for reconstruc-ion of fabric features such as foliation and lineation using olivine

in equal area projection and lower hemisphere, and contoured at interval 0.5 andred grains and P, G and R values are calculated from the eigenvalue analysis.

LPO data in strongly serpentinized peridotite. A preservation of thelattice preferred orientation in study samples shows that serpen-tinization postdate development of the LPO, and its randomizationis more likely related to dynamic recrystallization of olivine dur-ing development of porphyroclastic microstructure than to a lateserpentinization related event.

EBSD measurements of remnant olivine and pyroxene togetherwith the eigenvector analysis and the eigenvalue classification(Vollmer, 1990) helped to constrain the active slip system andprovide information about LPO symmetry. The studied sampleswith porphyroclastic microstructure show either [1 0 0](0 k l) LPOor axial [0 1 0] LPO pattern, implying that they have been deformedunder relatively high stress and ‘dry’ conditions (Jung and Karato,2001; Mainprice, 2007). However, origin of the olivine LPO is stillnot clearly established. The olivine [1 0 0](0 k l) pattern correspondsto the most common olivine slip patterns and is expected to occurat medium temperatures ca. 1000 ◦C (Carter and Ave’Lallemant,1970). It is commonly reported from peridotites with a coarse-grained microstructure (e.g. Ildefonse et al., 1995; Soustelle et al.,2010) and from high shear strain experiments (Bystricky et al.,2000). In terms of deformation regime this pattern of olivine LPOwas numerically modeled as a result of transtensional deformation(Tommasi et al., 1999). The axial [0 1 0] LPO pattern can either indi-cate contribution of the [0 0 1] glide in olivine (Demouchy et al.,2009; Tommasi et al., 2000), or it can be caused by flow componentorthogonal to shear plane in transpressional regime. This patternwas numerically simulated for dynamically recrystallized polycrys-talline olivine in transpressional regime (Tommasi et al., 1999).Recently, the latter pattern was measured in refractory harzbur-gites refertilized in mantle conditions, and explained as a resultof combination of recovery and static recrystallization that maymodify strain-induced axial [1 0 0] LPO pattern (Tommasi et al.,2008). The same patterns were also observed in Kerguelen xeno-
liths (Bascou et al., 2008), and in this case the least metasomatizedperidotites (harzburgites) display a axial [0 1 0] LPO pattern whilethe most metasomatized (dunites) display either an orthorhombicsymmetry or a [1 0 0](0 k l) pattern.


Fig. 8. (a) Orientations of foliation and lineation within the Mohelno serpentinized peridotite body according to the eigenvector analysis of olivine LPO. LPO revealed thatthe axial [0 1 0] pattern occur mainly along the inner margin and the southern limb of the fold, while the [1 0 0](0 k l) pattern occur in the rest of the body. Values of M-indexfor each measured sample are presented in the map with contours of the MPB; (b) point–girdle–random (PGR) diagrams for main crystallographic directions. Each symboli e LPO

l

6t

pAbitotapppqttsaptc

mb

n individual diagram represents one measured sample. Circles characterize olivinimb (white circles). Square symbols correspond to the [1 0 0](0 k l) pattern.

.1. Significance of the two types of olivine fabrics: a record of thehermal history of a quenched peridotite?

Our results confirm previously reported typology of olivine LPOatterns from the Mohelno Peridotite Body (Kamei et al., 2010).ccording to these authors, the existence of the two LPO types cane correlated to the thermal history of a hot peridotite emplaced

nto the colder granulites. Because of the slow cooling and con-inuous deformation in the interior of the peridotite body, theriginal [1 0 0](0 1 0) to axial [0 1 0] pattern (suggested as a high-emperature fabric) in the spinel peridotite was converted to thexial [1 0 0] pattern (suggested as a lower-temperature type). It wasroposed that the high-temperature [1 0 0](0 1 0) to axial [0 1 0]attern was preserved only at the garnet bearing margin of theeridotite body where cooling was rapid and the texture wasuenched. According to this model the reduction of grain sizehat occurred during later deformation partly obliterated the twoypes of previously developed fabric patterns in both garnet andpinel peridotites. The initial rapid cooling at high temperaturesssociated with deformation probably occurred after the mantleeridotite was emplaced into the granulites, which implies thathe spinel- to garnet-peridotite transformation took place in the
ontinental crust (Kamei et al., 2010).
This interpretation is mainly based on the petrological argu-ents and idea that garnet is not a primary mantle mineral,

ut crystallizes at the margin of the body due to differential

with axial [0 1 0] pattern either from the northern limb (black circles) or southern

cooling between the interior and the margin during emplace-ment of the peridotite sheet in the granulite rocks (Kamei et al.,2010; Medaris et al., 1990). According to the published chromiumcontent of primary spinel (Ycr = 0.121–0.193; Kamei et al., 2010),such an interpretation requires hypothetical crystallization of gar-net between 1200 and 1100 ◦C during an isobaric cooling atpressures of ∼2.3 GPa (Fig. 10a). However, such extreme PT con-ditions were never reported from Bohemian Massif granulites andsignificantly lower temperatures of 800–900 ◦C were commonlycalculated (Franek et al., 2011; Racek et al., 2008; Stípská et al.,2004; Tajcmanová et al., 2006). In addition, our detailed map-ping showed that garnet occurs only along the inner margin ofthe peridotite megafold and not all around the whole peridotitebody as it required by the quenching model. Also, ‘the quenching’of the high temperature LPO and microstructure in the body marginand its continuous alteration associated with decreasing temper-ature towards the centre of the peridotite does not fit with theregional grain size distribution of recrystallized pyroxenes reportedby Kamei et al. (2010). According to model, the larger grain sizesof pyroxenes now in the body centre should theoretically occurin the quenched body margin that preserves higher temperaturemicrostructure. On the basis of all these arguments, we do not
agree with the quenching model of LPO development and in thefollowing pages, we provide arguments favoring deformation his-tory that explains origin and regional distribution of both olivineLPO types.


F and lil nulite

6

soPotptgLmttt

dcaathotcttttdcivLiit

intracontinental rifting. The spinel preserved in garnet porphy-roblasts is interpreted as a relic of this stage (Medaris et al.,1990). Lithospheric thinning was followed by the developmentof a 70 km thick orogenic root (Schulmann et al., 2005, 2009).

Fig. 10. Pressure–temperature diagram showing results of previous petrologicalstudies and two contrasting PTt paths (grey dashed lines) for differential cooling ofcentral and marginal part of the Mohelno peridotite during ascent (Medaris et al.,

ig. 9. Lower hemisphere, equal area stereographic projections of poles to foliationsimb and the hinge zone; (c) LPO data from the fold southern limb and adjacent gra

.2. Relationship between field structure and olivine LPO

Field observations carried out in the Námest’ granulite bodyhow well preserved S2 steep mylonitic fabric concordant withlivine and orthopyroxene foliation of the fold-shaped Mohelnoeridotite Body (Fig. 3). In order to compare internal strain patternsf the peridotite with those of surrounding granulite, the orien-ations of the constructed olivine foliations and lineations wererojected in the map (Fig. 8a) and into the stereographic projec-ions (Fig. 9). The structural map shows that the olivine foliationsenerally follow the shape of the peridotite body for both types ofPO. The axial [0 1 0] type distribution is dominant along the innerargin of the body and along the southern limb of the large fold

hereby coinciding only partly with occurrence of garnet. In con-rast, the [1 0 0](0 k l) pattern is developed in the northern limb andhe hinge of the MPB fold.

The poles to foliation of all LPO data (Fig. 9a) are dominantlyistributed along a great circle and the pole of the great circle coin-ides with the megafold hinge (ˇ-axis of Sander, 1930; Fig. 3). Inddition, the lineations obtained from [1 0 0](0 k l) pattern occupy

similar position in the stereographic diagram (Fig. 9b). In con-rast, lineations from the axial [0 1 0] are rotated along the foldinge (Fig. 9c). This fabric pattern can be interpreted as a resultf large scale folding of pre-existing olivine fabric and its rota-ion around fold hinge. Our data show that the high M-indexoincides with prolate fabric in the hinge and northern limb ofhe MPB fold compared to oblate fabric and weak M-index inhe southern limb and inner part of the northern limb. This pat-ern can be explained by a two phase structural evolution. Duringhe first stage the whole peridotite body acquired weak oblateeformation related to emplacement of mantle sheet into lowerrustal granulites (Fig. 11b). Subsequently, the vertical N–S strik-ng sheet becomes folded due to horizontal N–S shortening andertical extrusion (Fig. 11c) as proposed by Franek et al. (2011) and
exa et al. (2011). The vertical extrusion is associated with fold-ng of the peridotite sheet and favors constrictional deformationn fold hinges, while folded limbs can preserve oblate fabrics. Ifhe vertical extrusion channel has circular or elliptical section, the
neations calculated from (a) all olivine LPO data; (b) LPO data from the fold northern.

constrictional flow is producing folds with vertical hinges due tolateral shortening of the vertical anisotropy (Kratinová et al., 2006;Weijermars, 1993).

6.3. Thermal and mechanical interactions between mantle andcrust

Medaris et al. (2005) proposed that the Mohelno peridotitecorresponds to the upper mantle depleted during Late Devonian

1990). A novel single PTt path is proposed (black dashed line) that reflects initialburial of the depleted peridotite to the garnet stability field, its further emplacementinto granulites and its rapid cooling. Garnet-spinel isoplets are after (1) Klemme andO’Neill (2000) and (2) Walter et al. (2002). Grt–spl isoplets shifted due to Y(Cr) inspinel after O’Neill (1981).


Fig. 11. Tectonic evolution of the Mohelno peridotite body: (a) rheological profile through thickened orogenic crust, spinel/garnet isograde according to Klemme and O’Neill( d empO n in tht idotite

Tsmaetp2stowetsl

dtfb(ot(dcbK(sa(

S

2000) and O’Neill (1981), (b) 3D model of imbrications of lithospheric mantle anblate ellipsoid is shown indicating development of first crystal plastic deformatio

o S2 fabrics, and development of F3 folds with steep fold hinges affecting both per

he crustal thickening model fits well with petrological studies ofpinel and garnet peridotites which reveal that both coarse-grainedicrostructure equilibrated at similar temperatures 1100–1200 ◦C

nd pressures of around 2.3 GPa (Fig. 10; Kamei et al., 2010; Medarist al., 1990). Such PT values could be regarded as peak PT condi-ions of the Mohelno peridotite, and the mineral zoning may reflectsosition of the spinel-garnet isograde in the mantle (Kamei et al.,010; Klemme and O’Neill, 2000; O’Neill, 1981). In such a setting thepinel–garnet transition is located only 5–10 km under the doublehickened Moho corresponding to a depth of 80 km (Fig. 11a). Rhe-logy of the lithosphere during both thinning and thickening stagesas calculated according to Thompson et al. (2001) and Schulmann

t al. (2002) and suggests that the warm and thin lithospheric man-le was inherited from the thinning stage. This tectonic setting isupposed to be the necessary prerequisite for further tectonic evo-ution (Fig. 11a).

Tectonic evolution continues by mylonitization of the peri-otites and the enclosed pyroxenite layers that could be linkedo the process of development of an intra-mantllic shear zoneollowed by imbrications of peridotite and granulite at theottom of the orogenic root during ongoing E–W shorteningFig. 11b). This shortening event is responsible for the formationf N–S trending crustal megafolds and gravity driven over-urns transporting lower crustal and mantle fragments upwardsLexa et al., 2011; Schulmann et al., 2008, 2009). The con-itions of mantle shearing during its emplacement into therust are recorded in clinopyroxene-spinel II kelyphite assem-lage around garnets, which may be correlated to kyanite,-feldspar and garnet assemblage in surrounding granulites

Kamei et al., 2010; Medaris et al., 1990, 2005). Progres-ively, both peridotite and granulite assemblages equilibrated
t temperatures and pressures of 800–900 ◦C and 1.5–0.8 GPaFigs. 10 and 11).
Field structural observations as well as EBSD study show that2 fabric in granulites and porphyroclastic fabric in peridotites

lacement of the peridotite sheet into the folded lower crust during the stage D2.e cooled peridotite. (c) Model of indentation of the Brunia promontory orthogonal

and granulite. Prolate ellipsoid shown for fold hinge deformation.

were actively folded during large scale F3 folding of the MPBassociated with partial melting of granulite and D3 amphibolitefacies retrogression. This event is correlated to the development ofamphibole-spinel coronas around garnet estimated to 700 ◦C and0.5–0.8 GPa (Fig. 10a; stage IV of Medaris et al., 1990). A similartransition from the granulite fine-grained mylonitic fabric (S2) tothe coarser-grained fabric (S3) associated to the partial meltinghas been described in other granulites in the Bohemian Massif asa result of decompression after rapid ascent of hot granulites tothe mid-crustal levels (Franek et al., 2006, 2011). The F3 foldingoccurred at the regional scale and is associated with indentation ofthe thickened orogenic root by the Brunian promontory (Fig. 11c) asproposed by Schulmann et al. (2008). According to these authors,the S4 fabric records the later subhorizontal channel flow whichtransported passively the peridotite and granulite fold in partiallymolten rocks above the Brunia basement.

Emplacement of the Mohelno peridotite body along the high-stress shear zone into the granulite followed by ascent to themid-crustal levels and active folding fits well also with bilin-ear cooling histories calculated from garnet composition. Medariset al. (1990) showed that the garnet core temperatures in excessof 1100 ◦C required extremely rapid cooling to quench the lowalmandine composition in the garnet interior. This event reflectsemplacement of the hot peridotite into the colder granulitic lowercrust during early stages of D2 event. The existence of garnet rimzoning requires a significant decrease of the cooling rate during theisothermal decompression associated with the vertical ascent ofgranulite and peridotite to the mid crustal levels during late stagesof D2 and D3.

6.4. Rheological constraints for mechanical behavior of peridotite
and granulite
An attempt is made to assess the rheological behavior of theMohelno peridotite body and adjacent granulites for the P-T-D


F cturea

htodcpaonozcgnWfmdaagt

psserdmat

ig. 12. Diagram illustrating the evolution of the peridotite and granulite microstrure proposed for different evolutionary stages of peridotite shearing and folding.

istory discussed above. The coarse-grained peridotite microstruc-ure cannot be successfully assessed because of insufficient numberf measurable grains of olivine and pyroxene. However, existingata from mantle xenoliths show the presence of LPO, which isommonly interpreted as a result of dislocation creep at high tem-eratures and slow strain rates (e.g. Ildefonse et al., 1995; Soustellend Tommasi, 2010). Dynamic recrystallization of study peridotiteccurred very likely at elevated stress conditions leading to sig-ificant grain size reduction down to 50 �m in diameter from anriginal coarse-grained microstructure in the intra-mantllic shearone. Microstructural analyses of other peridotite mylonites indi-ate that shear localization results from the combined effects ofrain size reduction, grain boundary sliding and second phase pin-ing during deformation (Kennedy et al., 2002; Toy et al., 2010;arren and Hirth, 2006). Microstructural analyses of the mylonite

abric in study as well as other peridotite bodies suggest that defor-ation of aggregates with a grain size of 10–100 �m occurred by

islocation-accommodated grain boundary sliding which producesn LPO. Grain boundary pinning, due to the mixing of pyroxenesnd spinel among olivine grains during dislocation accommodatedrain boundary sliding may result in permanent grain size reduc-ion (Jin et al., 1998; Vissers et al., 1995; Warren and Hirth, 2006).

According to our model (Figs. 11b and 12) the axial [0 1 0]attern originated at mantle P-T conditions during intra-mantllichearing leading to crystal plasticity characteristic for transpres-ional regime (Tommasi et al., 1999). Subsequent thrust relatedmplacement of peridotite sheet into the crust and associatedapid quenching led very likely to important hardening of the peri-
otite and freezing of intra-mantllic fabric. The D2 fine-grainedicrostructure of all phases in the host granulite suggests an
ccommodation of most of the strain during lower crustal ascentransporting passively peridotites (Fig. 12).

s for different tectonic stages along proposed PT path. Active slip systems in olivine

Our model also shows that the lineations calculated from the[1 0 0](0 k l) patterns coincide with the MPB fold hinge, whichsuggest that this LPO pattern was acquired during deformation atmid-crustal conditions estimated at ∼750 ◦C. This is in agreementwith experimental studies (Demouchy et al., 2009) as well asnatural mylonites (e.g. Warren and Hirth, 2006), which showthat the lithospheric mantle can be deformed plastically at rel-atively low temperatures (600–800 ◦C). It is during the late D3event when vertically anisotropic crust and mantle multilayeroriented at high angle to the advancing promontory (Fig. 11c)is folded thereby producing high stress concentrations in foldhinge regions (Schmalholz and Podladchikov, 1999) (Fig. 4). Inthis scenario, the folding of relatively strong peridotite layermay reactivate inherited fine-grained microstructure in hingezone and produce new LPO pattern. For such a microstructure, alow stress is enough to activate further plastic deformation viadislocation-accommodated grain boundary sliding that producesLPO (Jin et al., 1998; Vissers et al., 1995; Warren and Hirth, 2006).The host granulite structures like parasitic folds and axial cleavagezones indicate that the host rocks were significantly weaker andaccommodated folding of the peridotite layer.

Acknowledgements

This work was supported by Czech National Grant Agency(GACR) grant no. 205/09/0539, and the Research Intent of theGeophysical Institute, Academy of Sciences of the Czech Repub-lic AV0Z30120515, which are strongly acknowledged. Stays of V.
Kusbach at University of Strasbourg were founded by the FrenchGovernment Foundation (BGF). S.U. thanks David Mainprice,Andréa Tommasi and Alain Vauchez for stimulating discussions onLPO development of olivine.

eodyn

A

M

dpcetEiotob

D

miHtsfvcdaMh(tiDt(rof2w0

sasm(aRccsnRsnffpl

d


ppendix A. Supplementary data

-index and strength of lattice preferred orientation

Strength of a lattice preferred orientation can be expressed byifferent parameters. M-index was calculated for each olivine sam-le, also texture index (J-index) and fabric entropy parameters werealculated in Mtex (Bachmann et al., 2010). Texture index and fabricntropy parameters were calculated for both default kernel func-ion and optimal kernel function for ODF estimation from givenBSD dataset. Both parameters show linear correlation with M-ndex (Fig. A.1), which is in agreement with previous correlationsf M- and J-indexes in Soustelle et al. (2010). Therefore, we assumehe M-index as suitable representation of the fabric strength. Usingf the optimal kernel function causes systematic vertical shift foroth parameters.

eformation mechanism maps (DMMs)

Deformation mechanism maps were introduced in rock defor-ation studies, following their development in material science

n the seventies of the last century (Etheridge and Wilkie, 1979;andy, 1989). However, extrapolation of experimental deforma-

ion mechanism map to natural conditions of deformation is nottraightforward (e.g. Knipe, 1989; Means, 1990) and often wasar too large and locally even contrary to microstructural obser-ations. Nevertheless, several recent studies applied the DMMalculations to assess the rheological evolution of naturally peri-otites (Precigout et al., 2007; Warren and Hirth, 2006). Therefore,n attempt is made to assess the rheological behavior of theohelno peridotite body and adjacent granulites for the P-T-D

istory discussed above. Olivine deformation mechanism mapsDMM) were constructed for expected mantle–crust transitionemperature (900 ◦C) and for lower temperature limit for D3 fold-ng (700 ◦C) together with anorthite DMM for 700 ◦C (Fig. A.2).eformation mechanism maps show differential stress as a func-

ion of grain size for varied strain rates at constant temperatureFrost and Ashby, 1982). In our DMM the mantle rheology is rep-esented by experiments by Hirth and Kohlstedt (2003) for drylivine, whereas granulite is approximated by experimental dataor synthetic anorthite with 0.07% of water (Rybacki and Dresen,000). These plagioclase experiments were selected because ofater content of natural feldspars, which ranges from 0.02 to

.5 wt% (Johnson and Rossmann, 2003).Deformation mechanism maps (DMMs) are constructed to con-

train the deformation conditions of the peridotite and granulitend show differential stress as a function of grain size for differenttrain rates at constant temperature. Four different deformationechanisms are actually proposed for olivine: dislocation creep

GSI), diffusion creep (GSS), GBS creep (using terminology of Hirthnd Kohlstedt, 2003) and low-temperature plasticity (LTP) byaterron et al. (2004). Two different deformation mechanisms areonsidered for anortite (Rybacki and Dresen, 2000): dislocationreep (GSI) and diffusion creep (GSS). These experimental data onynthetic plagioclase were selected because of water content ofatural feldspars, which ranges from 0.02 to 0.5 wt% (Johnson andossmann, 2003). The overall strain rate ε is then defined as theum of the partial strain rates for each of the creeping mecha-isms: ε = εdisl + εdiff for feldspar and ε = εdisl + εdiff + εGBS + εLTP

or olivine, where εdisl is the strain rate for dislocation creep, εdiffor diffusion creep, εGBS for GBS creep and εLTP for low-temperaturelasticity. Partial strain rates can be calculated according power-

aw equations for

islocation creep (GSI) : εdisl = Adisl�ndisl exp

(−Qdisl

RT

), (A.1)

amics 56– 57 (2012) 108– 123 121

diffusion creep (GSS) : εdiff = Adiff�ndiff d−mdiff exp

(−Qdiff

RT

),

(A.2)

GBS creep : εGBS = AGBS�nGBS d−mGBS exp(−QGBS

RT

)(A.3)

low-temperature plasticity : εLTP = Altp exp

(−F0

[1 − (�/�)p]q

RT

)

(A.4)

where Adisl, Adiff, AGBS, ndisl, ndiff, nGBS, mdiff, mGBS, Qdisl, Qdiff, QGBS,R, T, � and d, respectively are: according to each deformationmechanism, the pre-exponential constant, the stress exponent, thegrain size exponent, the activation energy and, the gas constant8.314 J mol−1 K−1, the temperature, the stress and the grain size. ForLTP equation the additional parameters are: Altp a pre-exponentialconstant, F0 the free energy for the dislocations to overcome frictionand obstacles (sometimes refered as activation energy), � the max-imum glide resistance which quantifies the Peierls’ stress, p and qtwo fitting parameters. Set of all values for olivine and anortite arelisted in Table A.1.

The deformation maps are divided into several fields: disloca-tion creep field (GSI), diffusion creep field (GSS) and specificallyfor olivine DMMs also GBS creep and low-temperature plasticityfields. Each field represents stress vs. grain size space for which onedeformational mechanism is dominant e.g. for the dislocation creepfield (GSI) εdisl > (εdiff + εGBS + εLTP) the bulk strain rate is domi-nated by dislocation creep strain rate. Similarly in other DMM fieldsthe other deformational mechanisms are dominant and control thebulk strain rate.

Equilibrium between recrystallised grain size and stress isdefined by different piezometric relations. For DMM of olivine thepiezometers of Karato et al. (1980); Twiss (1977); Van der Wal et al.(1993) are used, for feldspar are shown HT piezometer of Twiss(1977) and LT piezometer of Post and Tullis (1999). All piezome-ters are listed in Table A.2. Our calculations confirm that the DMMcalculation approach is not robust enough to predict the rheologi-cal evolution of mantle rocks with the current database of naturalmicrostructural observations and experimental data. However, theDMM approach should be cultivated in the future to avoid sim-plifying use of dislocation creep equations in lithosphere scalenumerical modeling (Burov and Watts, 2006; Thompson et al.,2001).

Supplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.jog.2011.06.004.

References

Artyushkov, E.V., 1973. Stresses in the lithosphere caused by crustal thickness inho-mogeneities. J. Geophys. Res. 78, 7675–7708.

Babuska, V., Plomerová, J., 2006. European mantle lithosphere assembled from rigidmicroplates with inherited seismic anisotropy. Phys. Earth Planet. Inter. 158,264–280.

Babuska, V., Plomerová, J., Vecsey, L., 2008. Mantle fabric of western BohemianMassif (central Europe) constrained by 3D seismic P and S anisotropy. Tectono-physics 462, 149–163.

Bachmann, F., Hielscher, R., Schaeben, H., 2010. Texture analysis with MTEX—freeand open source software toolbox. Solid State Phenom. 160, 63–68.

Bascou, J., Delpech, G., Vauchez, A., Moine, B.N., Cottin, J.Y., Barruol, G., 2008. Anintegrated study of microstructural, geochemical, and seismic properties of thelithospheric mantle above the Kerguelen plume (Indian Ocean). Geochem. Geo-phys. Geosyst. 9, Q04036.

Brace, W.F., Kohlstedt, D.L., 1980. Limits on lithospheric stress imposed by laboratoryexperiments. J. Geophys. Res. 85 (B11), 6248–6252.

Brueckner, H., Carswell, D., Griffin, W., 2002. Paleozoic diamonds within a Precam-
brian peridotite lens in UHP gneisses of the Norwegian Caledonides. Earth Planet.Sci. Lett. 203, 805–816.
Brueckner, H.K., 1998. Sinking intrusion model for the emplacement of garnet-bearing peridotites into continent collision orogens. Geology 26, 631–634.

Bunge, H., 1982. Texture analysis in materials sciences. Buttleworth, London.

http://dx.doi.org/10.1016/j.jog.2011.06.004

1 eodyn

B

B

C

C

C

C

C

C

C

D

D

D

D

E

E

E

F

F

F

F

F

G

G

H

H

H

H

I

JJ

22 V. Kusbach et al. / Journal of G

urov, E., Watts, A., 2006. The long-term strength of continental lithosphere: “jellysandwich” or “crème brûlée”? Geol. Soc. Am. Bull. 16 (1), 4–10.

ystricky, M., Kunze, K., Burlini, L., Burg, J.-P., 2000. High shear strain ofolivine aggregates: rheological and seismic consequences. Science 290, 1564–1567.

arswell, D.A., O’Brien, P.J., 1993. Thermobarometry and geotectonic significanceof high-pressure granulites: examples from the Moldanubian zone of thebohemian massif in lower Austria. J. Petrol. 34, 427–459.

arter, N.L., Ave’Lallemant, H.G., 1970. High temperature flow of dunite and peri-dotite. Geol. Soc. Am. Bull. 81, 2181–2202.

hemenda, A.I., Mattauer, M., Malavieille, J., Bokun, A.N., 1995. A mechanism forsyn-collisional rock exhumation and associated normal faulting: results fromphysical modelling. Earth Planet. Sci. Lett. 132, 225–232.

hopra, P.N., Paterson, M.S., 1984. The role of water in the deformation of dunite. J.Geophys. Res. 89 (B9), 7861–7876.

ooke, R.A., 2000. High-pressure/temperature metamorphism in the St. LeonhardGranulite Massif, Austria: evidence from intermediate pyroxene-bearing gran-ulites. Int. J. Earth Sci. 89, 631–651.

ooke, R.A., O’Brien, P.J., 2001. Resolving the relationship between high P-T rocksand gneisses in collisional terranes: an example from the Gföhl gneiss–granuliteassociation in the Moldanubian Zone, Austria. Lithos 58, 33–54.

ouvy, H., Frost, D.J., Heidelbach, F., Nyilas, K., Ungar, T., Mackwell, S., Cordier, P.,2004. Shear deformation experiments of forsterite at 11 GPa—1400 ◦C in themultianvil apparatus. Eur. J. Miner. 16, 877–889.

emouchy, S., Schneider, S.E., Mackwell, S.J., Zimmerman, M.E., Kohlstedt, D.L., 2009.Experimental deformation of olivine single crystals at lithospheric tempera-tures. Geophys. Res. Lett. 36, L04304.

ijkstra, A.H., Drury, M.R., Vissers, R.L.M., Newman, J., Van Roermund, H.L.M., 2004.Shear zones in the upper mantle: evidence from alpine- and ophiolite-typeperidotite massifs. Geol. Soc., Lond., 11–24 (Spec. Publ. 224).

obretsov, N., Mísar, Z., Popov, E., 1984. The P-T conditions of equilibrium for somepyrope peridotite and country rocks in the Moldanubian area at Mohelno (East-ern Moravia, Czechoslovakia). Miner. Slov. 16, 87–95.

rury, M.R., Vissers, R.L.M., Wal, D., Hoogerduijn Strating, E.H., 1991. Shear localisa-tion in upper mantle peridotites. Pure Appl. Geophys. 137, 439–460.

del, J., Weber, K., 1995. Cadomian terranes, wrench faulting and thrusting in thecentral Europe Variscides: geophysical and geological evidence. Geol. Rundsch.84, 412–432.

llis, S., 1996. Forces driving continental collision: reconciling indentation and man-tle subduction tectonics. Geology 24, 699–702.

theridge, M., Wilkie, J., 1979. Grainsize reduction, grain boundary sliding and theflow strength of mylonites. Tectonophysics 58, 159–178.

iala, J., Matejovská, O., Vanková, V., 1987. Moldanubian granulites and related rocks:petrology, geochemistry and radioactivity. Rozpr. Cesk. Akad. Ved.Rada Mat. Prír.Ved. 97, 1–102.

ranke, W., 2000. The mid-European segment of the variscides: tectonostratigraphicunits, terrane boundaries and plate tectonic evolution. In: Franke, W., Haak,V., Oncken, O. (Eds.), Orogenic Processes: Quantification and Modelling in theVariscan Belt. Geological Society, London, pp. 35–61 (Special Publications 179).

ranek, J., Schulmann, K., Lexa, O., 2006. Kinematic and rheological model of exhuma-tion of high pressure granulites in the Variscan orogenic root: example ofthe Blansky les granulite, Bohemian Massif, Czech Republic. Miner. Petrol. 86,253–276.

ranek, J., Schulmann, K., Lexa, O., Tomek, C., Edel, J.-B., 2011. Model ofsyn-convergent extrusion of orogenic lower crust in the core of the Variscanbelt: implications for exhumation of high-pressure rocks in large hot orogens. J.Metamorph. Geol. 29, 53–78.

rost, H., Ashby, M., 1982. Deformation Mechanisms—Recognition from Natural Tec-tonites. Pergamon Press, Oxford.

uillope, M., Poirier, J.P., 1979. Dynamic recrystallization during creep ofsingle-crystalline halite: an experimental study. J. Geophys. Res. 84,5557–5567.

uy, A., Edel, J.-B., Schulmann, K., Tomek, C., Lexa, O., 2011. A geophysical model ofthe Variscan orogenic root (Bohemian Massif): implications for modern colli-sional orogens. Lithos 124, 144–157.

andy, M., 1989. Deformation regimes and the rheological evolution of fault zonesin the lithosphere: the effects of pressure, temperature, grainsize and time.Tectonophysics 163, 119–152.

asalová, P., Janousek, V., Schulmann, K., Stípská, P., Erban, V., 2008b. Fromorthogneiss to migmatite: geochemical assessment of the melt infiltrationmodel in the Gföhl Unit (Moldanubian Zone, Bohemian Massif). Lithos 102,508–537.

asalová, P., Schulmann, K., Lexa, O., Stípská, P., Hrouda, F., Ulrich, S., Haloda, J.,Tycová, P., 2008a. Origin of migmatites by deformation-enhanced melt infiltra-tion of orthogneiss: a new model based on quantitative microstructural analysis.J. Metamorph. Geol. 26, 29–53.

irth, G., Kohlstedt, D., 2003. Rheology of the upper mantle and the mantle wedge:a view from the experimentalists. Geophys. Monogr. 138, 83–105.

ldefonse, B., Billiau, S., Nicolas, A., 1995. A detailed study of mantle flow away fromdiapirs in the Oman ophiolite. In: Vissers, R., Nicolas, A. (Eds.), Mantle and LowerCrust Exposed in Oceanic Ridges and in Ophiolites. Kluwer, pp. 163–177.

ackson, J., 2002. Strength of the continental lithosphere. GSA Today 12 (Part 9), 4–9.anousek, V., Finger, F., Roberts, M., Fryda, J., Pin, C., Dolejs, D., 2004. Deciphering the

petrogenesis of deeply buried granites: whole-rock geochemical constraints onthe origin of largely undepleted felsic granulites from the Moldanubian Zone ofthe Bohemian Massif. Earth Env. Sci. T. R. So. 95, 141–159.

amics 56– 57 (2012) 108– 123

Janousek, V., Holub, F.V., 2007. The causal link between HP-HT metamorphism andultrapotassic magmatism in collisional orogens: case study from the Moldanu-bian Zone of the Bohemian Massif. Proc. Geol. Assoc. 118, 75–86.

Jin, D., Karato, S.-I., Obata, M., 1998. Mechanisms of shear localization in thecontinental lithosphere: inference from the deformation microstructures ofperidotites from the Ivrea zone, northwestern Italy. J. Struct. Geol. 20,195–209.

Johnson, E., Rossmann, G., 2003. The concentration and speciation of hydro-gen in feldspars using FTIR and 1H MAS NMR spectroscopy. Am. Miner. 88,901–911.

Jung, H., Karato, S.-I., 2001. Water-induced fabric transitions in olivine. Science 293,1460–1463.

Kamei, A., Obata, M., Michibayashi, K., Hirajima, T., Svojtka, M., 2010. Two contrast-ing fabric patterns of olivine observed in garnet and spinel peridotite from amantle-derived ultramafic mass enclosed in felsic granulite, the MoldanubianZone, Czech Republic. J. Petrol. 51, 101–123.

Karato, S.-I., Toriumi, M., Fujii, T., 1980. Dynamic recrystallization of olivine singlecrystals during high-temperature creep. Geophys. Res. Lett. 7, 649–652.

Kennedy, L.A., Russell, J.K., Kopylova, M.G., 2002. Mantle shear zones revisited: theconnection between the cratons and mantle dynamics. Geology 30, 419–422.

Klemme, S., O’Neill, H.S., 2000. The near-solidus transition from garnet lherzolite tospinel lherzolite. Contrib. Miner. Petrol. 138, 237–248.

Knipe, R., 1989. Deformation mechanisms—recognition from natural tectonites. J.Struct. Geol. 11, 127–146.

Kratinová, Z., Závada, P., Hrouda, F., Schulmann, K., 2006. Non-scaled analogue mod-elling of AMS development during viscous flow: a simulation on diapir-likestructures. Tectonophysics 418, 51–61.

Kretz, R., 1983. Symbols for rock-forming minerals. Am. Mineral. 68, 277–279.Lexa, O., Schulmann, K., Janousek, V., Stípská, P., Guy, A., Racek, M., 2011. Heat sources

and trigger mechanisms of exhumation of HP granulites in Variscan orogenicroot. J. Metamorph. Geol. 29, 79–102.

Liang, X., Zhou, S., Chen, Y.J., Jin, G., Xiao, L., Liu, P., Fu, Y., Tang, Y., Lou, X., Ning, J.,2008. Earthquake distribution in southern Tibet and its tectonic implications. J.Geophys. Res. Sol. Earth 113, B12409.

Machek, M., Ulrich, S., Janousek, V., 2009. Strain coupling between upper mantleand lower crust: natural example from the Bestvina granulite body, BohemianMassif. J. Metamorph. Geol. 27, 721–737.

Mackwell, S.J., Kohlstedt, D.L., Paterson, M.S., 1985. The role of water in the defor-mation of olivine single crystals. J. Geophys. Res. 90 (B13), 319–333.

Mainprice, D., 2007. Seismic anisotropy of the deep Earth from a mineral and rockphysics perspective. In: Schubert, G. (Ed.), Treatise on Geophysics, vol. 2. ElsevierLtd., pp. 437–492.

Mainprice, D., 1990. A FORTRAN program to calculate seismic anisotropy from thelattice preferred orientation of minerals. Comput. Geosci. 16, 385–393.

Matejovská, O., 1967. Petrogenesis of the Moldanubian granulites near Námest’ nadOslavou. Krystalinikum 5, 85–104.

Matejovská, O., 1975. The Moldanubian gneiss series of south-western Moravia andits relation to granulites. Vestn. Ústred. Ústav. Geol. 50, 345–351.

Means, W., 1990. Kinematics, stress, deformation and material behavior. J. Struct.Geol. 12, 953–971.

Medaris, L., Ackerman, L., Jelínek, E., Toy, V., Siebel, W., Tikoff, B., 2009. The Sklenégarnet peridotite: petrology, geochemistry, and structure of a mantle-derivedboudin in Moldanubian granulite. J. Geosci. 54 (4), 301–323.

Medaris, L.J., Wang, H., Jelínek, E., Mihaljevic, M., Jakes, P., 2005. Characteristics andorigins of diverse Variscan peridotites in the Gföhl Nappe, Bohemian Massif,Czech Republic. Lithos 82, 1–23.

Medaris, L.J., Wang, H., Mísar, Z., Jelínek, E., 1990. Thermobarometry, diffusion mod-elling and cooling rates of crustal garnet peridotites: two examples from theMoldanubian zone of the Bohemian Massif. Lithos 25, 189–202.

Molnar, P., Chen, W.-P., 1983. Focal depths and fault plane solutions of earthquakesunder the Tibetan plateau. J. Geophys. Res. 88, 1180–1196.

Molnar, P., England, P., Martinod, J., 1993. Mantle dynamics, uplift of the TibetanPlateau, and the Indian Monsoon. Rev. Geophys. 31, 357–396.

Molnar, P., Lyon-Caen, H., 1988. Some simple physical aspects of the support, struc-ture, and evolution of mountain belts. Geol. Soc. Am., 179–207 (Spec. Pap. 218).

Monsalve, G., McGovern, P., Sheehan, A., 2009. Mantle fault zones beneath theHimalayan collision: flexure of the continental lithosphere. Tectonophysics 477,66–76.

O’Neill, H.S.C., 1981. The transition between spinel lherzolite and garnet lherzolite,and its use as a Geobarometer. Contrib. Miner. Petrol. 77, 185–194.

Post, A., Tullis, J., 1999. A recrystallized grain size piezometer for experimentallydeformed feldspar aggregates. Tectonophysics 303, 159–173.

Precigout, J., Gueydan, F., Gapais, D., Garrido, C., Essaifi, A., 2007. Strain localisa-tion in the subcontinental mantle—a ductile alternative to the brittle mantle.Tectonophysics 445, 318–336.

Racek, M., Stípská, P., Pitra, P., Schulmann, K., Lexa, O., 2006. Metamorphic record ofburial and exhumation of orogenic lower and middle crust: a new tectonother-mal model for the Drosendorf window (Bohemian Massif, Austria). Miner. Petrol.86, 221–251.

Racek, M., Stípská, P., Powell, R., 2008. Garnet-clinopyroxene intermediate gran-ulites in the St. Leonhard massif of the Bohemian Massif: ultrahigh-temperature
metamorphism at high pressure or not? J. Metamorph. Geol. 26, 253–271.
Ramberg, H., 1981. The role of gravity in orogenic belts. Geol. Soc. Lond., 125–140(Spec. Publ. 9).

Ranalli, G., Murphy, D.C., 1987. Rheological stratification of the lithosphere. Tectono-physics 132, 281–295.

eodyn

R

R

R

R

R

SS

S

S

S

S

S

S

S

S

S

S

S

S

S

S

Further reading


aterron, P., Amiguet, E., Chen, J., Li, L., Cordier, P., 2009. Experimental deformation ofolivine single crystals at mantle pressures and temperatures. Phys. Earth Planet.Inter. 172, 74–83.

aterron, P., Wu, Y., Weidner, D.J., Chen, J., 2004. Low-temperature olivine rheologyat high pressure. Phys. Earth Planet. Inter. 145, 149–159.

oux, V.L., Tommasi, A., Vauchez, A., 2008. Feedback between melt percolation anddeformation in an exhumed lithosphere-asthenosphere boundary. Earth Planet.Sci. Lett. 274, 401–413.

utter, E., Brodie, K., 1992. Rheology of the lower crust. In: Fountain, D., Kay, A.R.(Eds.), Continental Lower Crust. Elsevier, pp. 201–267.

ybacki, E., Dresen, G., 2000. Dislocation and diffusion creep of synthetic anorthiteaggregates. J. Geophys. Res. 105, 26017–26036.

ander, B., 1930. Gefügekunde der Gesteine. Springer, Vienna.chmalholz, S.M., Podladchikov, Y., 1999. Buckling versus folding: importance of

viscoelasticity. Geophys. Res. Lett. 26, 2641–2644.chmidt, N.-H., Olesen, N.O., 1989. Computer-aided determination of crystal-lattice

orientation from electron channeling patterns in the SEM. Can. Miner. 27, 15–22.chmädicke, E., Evans, B.W., 1997. Garnet-bearing ultramafic rocks from the Erzge-

birge, and their relation to other settings in the Bohemian Massif. Contrib. Miner.Petrol. 127, 57–74.

chmädicke, E., Gose, J., Will, T.M., 2010. The P-T evolution of ultra high temper-ature garnet-bearing ultramafic rocks from the Saxonian Granulitgebirge CoreComplex, Bohemian Massif. J. Metamorph. Geol. 28, 489–508.

chulmann, K., Konopásek, J., Janousek, V., Lexa, O., Lardeaux, J.-M., Edel, J.-B., Stípská,P., Ulrich, S., 2009. An Andean type Palaeozoic convergence in the BohemianMassif. C. R. Geosci. 341, 266–286.

chulmann, K., Kroner, A., Hegner, E., Wendt, I., Konopásek, J., Lexa, O., Stípská, P.,2005. Chronological constraints on the pre-orogenic history, burial and exhuma-tion of deep-seated rocks along the eastern margin of the Variscan Orogen,Bohemian Massif, Czech Republic. Am. J. Sci. 305, 407–448.

chulmann, K., Martelat, J.-E., Ulrich, S., Lexa, O., Stípská, P., Becker, J.K., 2008. Evolu-tion of microstructure and melt topology in partially molten granitic mylonite:implications for rheology of felsic middle crust. J. Geophys. Res. 113, B10406.

chulmann, K., Schaltegger, U., Jezek, J., Thompson, A.B., Edel, J.-B., 2002. Rapid burialand exhumation during orogeny: thickening and synconvergent exhumation ofthermally weakened and thinned crust (Variscan orogen in Western Europe).Am. J. Sci. 302, 856–879.

kemer, P., Katayama, I., Jiang, Z., Karato, S., 2005. The misorientation index:development of a new method for calculating the strength of lattice-preferredorientation. Tectonophysics 411, 157–167.

oustelle, V., Tommasi, A., 2010. Seismic properties of the supra-subduction man-tle: constraints from peridotite xenoliths from the Avacha volcano, southernKamchatka. Geophys. Res. Lett. 37, L13307.

oustelle, V., Tommasi, A., Bodinier, J.L., Garrido, C.J., Vauchez, A., 2009. Deformationand reactive melt transport in the mantle lithosphere above a large-scale par-tial melting domain: the ronda peridotite massif, Southern Spain. J. Petrol. 50,1235–1266.

oustelle, V., Tommasi, A., Demouchy, S., Ionov, D.A., 2010. Deformation and fluid-rock interaction in the supra-subduction mantle: microstructures and watercontents in peridotite xenoliths from the Avacha volcano, Kamchatka. J. Petrol.51, 363–394.

ˇichtarová, I., 1981. Moldanubian amphibolites in the area SE of Námest’ nadOslavou. Vestn. Ústred. Ústav. Geol. 56, 203–214.

ˇtípská, P., Schulmann, K., Kroner, A., 2004. Vertical extrusion and middle crustalspreading of omphacite granulite: a model of syn-convergent exhumation
(Bohemian Massif, Czech Republic). J. Metamorph. Geol. 22, 179–198.
ˇtípská, P., Schulmann, K., Powell, R., 2008. Contrasting metamorphic histories oflenses of high-pressure rocks and host migmatites with a flat orogenic fabric(Bohemian Massif, Czech Republic): a result of tectonic mixing within horizontalcrustal flow? J. Metamorph. Geol. 26, 623–646.

amics 56– 57 (2012) 108– 123 123

Tajcmanová, L., Konopásek, J., Schulmann, K., 2006. Thermal evolution of the oro-genic lower crust during exhumation within a thickened Moldanubian root ofthe Variscan belt of Central Europe. J. Metamorph. Geol. 24, 119.

Thompson, A., Schulmann, K., Jezek, J., Tolar, V., 2001. Thermally softened continen-tal extensional zones (arcs and rifts) as precursors to thickened orogenic belts.Tectonophysics 332, 115–141.

Tikoff, B., Russo, R., Teyssier, C., Tommasi, A., 2004. Mantle-driven deformation oforogenic zones and clutch tectonics. In: Grocott, J., McCaffrey, K.J.W., Yaylor, G.,Tikoff, B. (Eds.), Vertical Coupling and Decoupling in the Lithosphere. GeologicalSociety, London, pp. 41–64 (Special Publications).

Tikoff, B., Teyssier, C., Waters, C., 2002. Clutch tectonics and the partial attachmentof lithospheric layers. EGU S. Mueller Spec. Publ. Ser. 1, 57–73.

Tollmann, A., 1982. Grossräumiger variszischer Deckenbau im Moldanubikum undneue Gedanken zum Variszikum Europas. Geotekton. Forsch. 64, 1–91.

Tommasi, A., Mainprice, D., Canova, G., Chastel, Y., 2000. Viscoplastic self-consistentand equilibrium-based modeling of olivine lattice preferred orientations:implications for the upper mantle seismic anisotropy. J. Geophys. Res. 105,7893–7908.

Tommasi, A., Tikoff, B., Vauchez, A., 1999. Upper mantle tectonics: three-dimensional deformation, olivine crystallographic fabrics and seismic proper-ties. Earth Planet. Sci. Lett. 168, 173–186.

Tommasi, A., Vauchez, A., Ionov, D.A., 2008. Deformation, static recrystalliza-tion, and reactive melt transport in shallow subcontinental mantle xenoliths(Tok Cenozoic volcanic field, SE Siberia). Earth Planet. Sci. Lett. 272,65–77.

Toy, V.G., Newman, J., Lamb, W., Tikoff, B., 2010. The role of pyroxenites in formationof shear instabilities in the mantle: evidence from an ultramafic ultramylonite,twin sisters massif, Washington. J. Petrol. 51, 55–80.

Twiss, R.J., 1977. Theory and applicability of a recrystallized grain size paleopiezome-ter. Pure Appl. Geophys. 115, 227–244.

Urban, M., 1992. Kinematics of the Variscan thrusting in the Eastern Moldanubicum(Bohemian Massif, Czechoslovakia): evidence from the Námest’ granulite massif.Tectonophysics 201, 371–391.

Van der Wal, D., Chopra, P., Drury, M., Gerald, J.F., 1993. Relationships betweendynamically recrystallized grain size and deformation conditions in experimen-tally deformed olivine rocks. Geophys. Res. Lett. 20, 1479–1482.

Vauchez, A., Tommasi, A., Barruol, G., 1998. Rheological heterogeneity, mechanicalanisotropy and deformation of the continental lithosphere. Tectonophysics 296,61–86.

Vissers, R.L.M., Drury, M.R., Hoogerduijn-Strating, E.H., Spiers, C.J., van der Wal,D., 1995. Mantle shear zones and their effect on lithosphere strength duringcontinental breakup. Tectonophysics 249, 155–171.

Vollmer, F.W., 1990. An application of eigenvalue methods to structural domainanalysis. Geol. Soc. Am. Bull. 102, 786–791.

Walter, M., Katsura, T., Kubo, A., Shinmei, T., Nishikawa, O., Ito, E., Lesher,C., Funakoshi, K., 2002. Spinel-garnet lherzolite transition in the systemCaO–MgO–Al2O3–SiO2 revisited: an in situ X-ray study. Geochim. Cosmochim.Acta 66, 2109–2121.

Warren, J., Hirth, G., 2006. Grain size sensitive deformation mechanisms in naturallydeformed peridotites. Earth Planet. Sci. Lett. 248, 438–450.

Weijermars, R., 1993. Progressive deformation of single layers under constantlyoriented boundary stresses. J. Struct. Geol. 15, 911–922.

Stípská, P., Powell, R., 2005. Does ternary feldspar constrain the metamorphicconditions of high-grade meta-igneous rocks? Evidence from orthopyroxenegranulites, Bohemian Massif. J. Metamorph. Geol. 23, 627–647.

Ductile deformation and rheology of sub-continental mantle in a hot collisional orogeny: Example from the Bohemian Massif

Documents