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1, metamorphic Ceol., 1996, 14, 49-60 Probable anticlockwise P-T evolution in extending crust: Hlinsko region, Bohemian Massif P. P I T R A ' , ' A N D M. G U I R A U D 2 7 ' Ustav petrologie a strukturnigeologie, Universita Karlova, Albertov 6, 128 43 Praha 2, the Czech Republic 'Laboratoire de Min&alogie, URA CNRS 736, Muse'um National d'Histoire Naturelle, 61 rue Buffon, 75 005 Paris, France ABSTRACT In the Hlinsko region (Variscan Bohemian Massif, Czech Republic) a major extensional shear zone separates low-grade metasedimentary series (Hlinsko schists) and high-grade rocks of the Moldanubian terrane (Svratka Crystalline Unit). During late-Variscan extension, a tonalite intruded syntectonically into the normal ductile shear zone, and caused contact metamorphism of the overlying schists. Concurrent syntectonic sedimentation of a flysch series took place at the top of the hangingwall schists. In order to decipher the detailed petrological evolution of the Hlinsko unit situated in the hangingwall of this tectonic contact, a phase diagram approach and pctrogenetic grids, calculated with the thermocalc computer program, were used. The crystallization/deformation relationships and the paragenetic analysis of the Hlinsko schists define a P-T path with an initial minor increase in pressure followed by cooling. Calculated pseudosections constrain this anticlockwise P-T evolution to the upper part of the andalusite field between 0.36 and 0.40 G P a for temperatures ranging from 570 to 530 -C. A low uHZ0is required to explain the presence of andalusite-biotite-bearing assemblages, and could be related to the presence of abundant graphite. I n contrast, the footwall rocks of the Svratka Crystalline Unit record decompression from around 0.8 GPa at a relatively constant temperature, followed by cooling. Thus, the footwall and the hangingwall units display opposite, but convergent P-Thistories. Decompression in the footwall rocks is related to a rapid exhumation. We propose that the inverse, anticlockwise P-T path recorded in the hangingwall pelites is related to the rapid, extension-controlled sedimentation of the overlying flysch series. Key words: anticlockwise P-T path; Bohemian Massif; late-orogenic extension; low-P metamorphism; Variscan INTRODUCTION Numerous studies have shown that the final stages of the orogenic evolution of a collision belt are dominated by extensional tectonics (Coney & Harms, 1984; Dewey, 1988; Gaudemer et ul., 1988; Menard & Molnar, 1988; Burg et ul., 1994b). Although the petrology of rocks lying in the footwall of major normal faults is relatively well documented, knowledge concerning the evolution of the hangingwall units is lacking. In the Hlinsko region (Variscan Bohemian Massif, Czech Republic) a major extensional shear zone is observed between low-grade metasedimentary series (Hlinsko schists) a n d high-grade rocks of the Moldanubian terrane (Svratka Crystalline UnitSCU) that are situated in the hangingwall a n d in the footwall, respectively. The extensional character of this tectonic contact and of the related structures in both units has been demonstrated (Pitra et ul., 1994). The purpose of this paper is to present the detailed Corresponrlence: Pave1 Pitra, Laboratoire de Mineralogie, MNHN, 61 rue Buffon, 75 005 Paris, France (email: [email protected]). petrological evolution of the hangingwall Hlinsko unit and to discuss its possible geodynamic significance within the extensional context. GEOLOGICAL SETTING The Hlinsko Palaeozoic sediments form a N-Selongated, E-verging synform bounded by the Nasavrky plutonic complex to the west and by the high-grade metamorphics of the Svratka Crystalline Unit to the south-east (Fig. 1). From the bottom to the top, four units are identified (Vachtl, 1962): (1) a pelite and metavolcanic sequence (Vitanov series) supposedly of Upper Proterozoic age, (2) pelites and greywackes of the Hlinsko series for which a n Ordovician age has been proposed, (3) the Mrikotin series that contains pelites, graphitic pelites and lydites, and quartzite beds, a n d whose Silurian age has been confirmed by graptolites (Wurm, 1927; Horny, 1956), and (4) the Rychmburk series, which forms a thick sequence of pelites and greywackes with conglomerate intercalations. The age of these strata is not clear (Proterozoic, Siluro-Devonian o r Carboniferous). The latter series displays a flysch character with general grain coarsening observed towards the NE, i.e. upwards 49 50 P . P I T R A & M. C U I R A U D 1PE 16"E 51"N 49"N a, C 0 e, s CF 0) Ul 0 % .-C z Greywackes (Rychmburk) Graphitic pelites a Greywackes -pelites (Hlinsko) Volcano-sedim. rocks(Vitanov) Basic rocks Deformed a Acid rocks Normal ductile shearzone t/ Fig. 1. Schematic geological map of the Hlinsko region and its location within the Bohemian Massif. Encircled numbers refer to samples in Table I . in the sequence and when approaching the faulted contact with the tonalite (Vachtl, 1950). Moldanubianderived clasts were found in this part of the series within the conglomerates (Chab, 1973), as well as numerous synsedimentary normal faults. Thus we interpret this series as a synextensional basin of Carboniferous age, which fits well with the age of extension in the Bohemian Massif (Pitra et al., 1994). The Svratka Crystalline Unit, probably a part of the Moldanubian terrane, comprises paragneisses, orthogneisses and migmatites intercalated with minor mica schist beds, lenticular bodies of amphibolites and skarns parallel to the main foliation. A tonalite laccolith has intruded syntectonically the contact with the Hlinsko sediments, triggering HT-LP metamorphic crystallization in the overlying schists (Pitra et a/., 1994). Three deformational events are regionally distinguished in the Hlinsko metasediments. The S 1 cleavage is marked by the alignment of muscovite, quartz, ilmenite and possibly chlorite and chloritoid in the lowermost pelites. Second-generation structures are well developed, with S2 being a crenulation cleavage of variable intensity. S2 average spacing decreases toward the contact with the tonalite and the base of the sedimentary pile. The SE-NW-striking subhorizontal and synmetamorphic lineation L2 is marked by the long axis of staurolite and chloritoid blasts, and/or by their boudinage. S2 is folded by the deformational event D3. The S3 crenulation cleavage is associated with the axial plane of the regional syncline and displays a strain gradient with an overall westward increase in intensity from the hinge of the Hlinsko syncline. It bears subvertical mineral stretching lineations near the contact with the Nasavrky granodiorite. Only S3 is observed within the uppermost part of the Rychmburk sedimentary series. The late Nasavrky granodiorite intruded the western part of the region producing limited contact metamorphism (Sachsel, 1933; Vachtl, 1962). Steep magmatic foliation and lineation that typically occur at the pluton margins are consistent with the D3 strain gradient in the country rocks. The granodiorite has been dated at 360-366 Ma (K/Ar, whole-rock; Smejkal, 1960, 1964). However, recent dating at 320+4Ma (Rb/Sr, whole-rock; Scharbert, 1987) suggest a younger age for the associated granites. It is assumed that D1 represents the local, and earliest stages of the D2 north-westward shear deformation. D3 is related to the emplacement of the Nasavrky granodiorite and is responsible for the regional Hlinsko syncline. PETROLOGY OF THE H L I N S K O SCHISTS Petrography and mineral chemistry In the Silurian (Mrakotin) pelites the S1 foliation is parallel to the sedimentary layering SO defined by muscovite-rich and quartz-rich layers. The relative A N T I C L O C K W I S E P-T PATH, B O H E M I A N M A S S I F ~~~ ~ ~ 51 ~ abundance of these layers varies between end-members with an equal proportion of both types, and rocks dominated by muscovite-rich layers (>90%). Pelites become generally more muscovite-rich when approaching the contact with the tonalite. The thickness of the layers is about 1 cm in rocks with preserved sedimentary layering and decreases to about 1 mm where S2 is intensely transposed near the contact with the tonalite. Muscovite-rich layers are composed of finegrained muscovite (70%), andalusite and/or Fe-Mg minerals (staurolite, biotite, garnet, cordierite, chlorite up to 25%), ilmenite and quartz ( < 5 % ) . The quartzrich layers are composed of quartz ( S O % ) , muscovite (40%), ilmenite, garnet and chlorite, with other Fe-Mg minerals being less abundant. All the Mrakotin pelites are relatively Fe- and Al-rich. A significant difference exists in bulk rock composition between the centre of the syncline where more aluminous and less ferriferous lithologies occur [AI/(Al + Fe + Mg) = 0.88, Fe/( Fe Mg) =0.60] and the SE limb, near the tonalite, which is characterized by more ferriferous and less aluminous lithologies [(Al/(Al+Fe+Mg)=0.79, Fe/(Fe+Mg)=0.64]. The representative mineral chemistry of individual phases is listed in Table 1. + Andalusite. Andalusite is zoned chiastolite and contains inclusions of biotite, garnet and euhedral staurolite (Fig. 2a). It encloses also oval-shaped aggregates of chlorite, biotite, muscovite and quartz. Straight inclusion trails of ilmenite and graphite suggest that andalusite grew partly before S2 as intense crenulation is present outside crystals. The andalusite is ferric with the Fe,O, content spanning 0.3-0.8 wt% and does not contain any other oxides in significant amount. Staurolite. Staurolite is euhedral and is often optically zoned. Crystals may be broken, with chlorite crystallizing between fragments. It includes crenulated S1 trails indicating that staurolite is late- to post-S2 (Fig. 2b). However, staurolite inclusions within peripheral parts of andalusite indicate that staurolite grew along with and later than andalusite during D2. No significant chemical zonation has been observed and the chemical variations are supposed to reflect only change in bulkrock composition. The staurolite is rich in ZnO (0.1-1.3 wt%), TiO, (0.15-0.65 wt%) and MnO (0.350.60 wt%) and X,, ranges from 0.85 to 0.93. Biotite. Biotite crystals have overgrown S1 and are deformed by S2. Although displaying microscopically the usual optical features, biotite has been systematically transformed to oxychlorite. In the AFM diagram projected from muscovite, compositions plot between the position of biotite and that of chlorite and X,, increases from 0.59 to 0.70 as it shifts towards chlorite. As the compositions also tend towards the staurolite position, it is suggested that the continuous destabilization of biotite to chlorite involves staurolite (see Fig. 4). In some thin sections, biotite zonation is marked by centres rich in ilmenite inclusions and clear rims. Textural equilibrium with andalusite, staurolite and garnet is common. Chloritoid. Chloritoid lies within, and locally across, S2. Near the centre of the regional syncline, it forms tiny elongated rods in chloritized biotite, possibly as a Table 1. Chemical compositions of principal metamorphic minerals of the Hlinsko pelites. Chl,,, represents the chlorite composition within the andalusite inclusion. Numbers in the first row refer to the location on Fig. 1. Locallon SIO, TIO, AI,O, CrD, M@ FeO MnO ZnO CaO Na,O K D Sum SI TI A1 Cr Mg + hl n Zn Ca Na K OH Sum X FC 2 I Ms H42 Ms H82 2 St H42 St 248 I Bt H82 27 06 0 28 55 42 0.11 1.28 13.53 0.34 1.31 0.00 0.00 0.00 99.33 2x31 0 29 55 40 0.02 0 84 13.44 0 33 0 74 0.00 0.00 0.00 99 37 34 06 0.86 20.01 0 00 5.99 23.62 0 12 0 00 0.00 0 16 7.03 91.85 35.65 I .07 20 95 0.02 7.93 20.61 0.12 0 00 0.00 0.34 8.14 94.83 35 97 1.06 21.80 0.00 5.93 18.46 0.07 0.00 0 09 0 08 5.14 88.60 35 81 0.71 23.81 0.00 5.04 20.21 0 12 0 00 0.00 0.00 3.34 89.04 24.50 0.00 40.95 0.00 1.83 24.85 0.57 000 000 no0 000 92.70 24 11 0.08 4045 0.00 2.41 2507 077 000 000 0.00 0.04 9293 37 07 0.00 20.81 0.00 1.33 29 62 10.46 0.00 136 0 00 0 00 100.65 37.44 0.03 21.18 0 00 135 31 99 8.21 0.00 1.17 0 05 0 04 101 46 36.74 0.00 21.49 0 07 170 33.68 4.61 0.05 1.60 0 02 0 03 100.05 23 47 0.08 23.11 0 00 10.16 31.32 0 50 0 00 0.00 0.00 0.00 88.64 23.14 0 09 23 99 0.00 10 75 29.84 0.19 0.00 0.00 0.00 0.27 88 21 24 14 0.03 23 58 0.00 13.23 27.56 031 0 00 0 00 n 00 0.00 88.85 0.55 53.16 0 26 0.00 0.00 43 43 1.83 0.00 0.04 0.00 0.13 100.00 45 64 0.14 37.37 0.00 0.29 1.16 0.02 0 00 0.00 195 8.48 95.05 46.97 o 04 37.1 I 0.00 0.32 0.88 0 01 0.00 0.00 I .40 9.20 95.93 7 77 0 06 1876 0 03 0 55 3 25 0.08 0 28 0 00 0.00 0 00 8.07 0 06 18.62 0 01 0 36 3 22 0 08 0.16 0.00 0 00 0 00 2.83 0.06 2.02 0.00 0.70 1.21 001 0 00 0.01 0.01 0.52 2.00 9.36 0.64 2.79 0.04 2 18 0.00 0.58 131 001 0 00 0.00 0.00 0.33 2 00 9 25 0.69 101 0.00 1.99 0.00 0.11 0.86 0.02 0.00 000 0.00 0.00 2.00 5 99 0.88 1.00 0.00 1 97 0.00 015 0.87 003 0.00 000 0.00 000 2.00 6.02 085 3.01 0 00 2.01 0.00 0 I6 2.15 0.56 0.00 0 10 0 01 0 00 2 98 0.00 2.05 0 00 021 2.29 0.32 0.00 0.14 0.00 0 00 8.00 0.93 8.00 0.93 7.99 0 92 2.53 001 2 94 0 00 1 64 2 83 0.05 0.00 0.00 0.00 0.00 8 00 17.99 0.63 2.49 001 3.04 0.00 1.72 2.68 0.02 0.00 0.00 0 00 0.04 8.00 18.00 0.61 2.54 0.00 2 92 0.00 2.08 2.43 0 03 0 00 0.00 0.00 0.00 8.00 I 8.00 0.54 001 I01 0 01 0.00 0.00 091 0 04 0.00 0.00 0.00 0.00 30.50 0 90 2.71 0 06 1.88 0.00 0.90 I31 0.01 0.00 0.00 0.05 0.79 2 00 9 71 0.59 3 01 0.00 1.99 0.00 0.16 2.01 0.72 0.00 0.12 0.00 0.00 30.68 0.86 2.71 0.05 1 88 0.00 0.71 0.57 001 0 00 0.00 0.03 0.71 2.00 9.67 0 69 3 02 0.01 2 91 0.00 0.03 0.06 0 00 0 00 0.00 0.25 0.72 2.00 9 00 0 69 3.07 0.00 2 86 0.00 0 03 0 05 0 00 0.00 0.00 0.18 0.77 2.00 8.97 061 3 3 3 3 I 3 3 3 4 2 4 Bt 248 G Bt 24811 Bt 248 P Cld H82 Cld 248 Grt 248c Grt 2481. Grt H2r. Chl H42 Chi H2 2 Chl,,, 3 Ilm 248 198 52 P . P I T R A t; M. C U I K A U D ANTICLOCKWISE P-T product of the destabilization of biotite (Fig. 2c). However, although rarely, it can also be observed in equilibrium with biotite near the boundary with the tonalite (Fig. 2d). Textural equilibrium exists between chloritoid, staurolite and chlorite. Locally, chloritoid was retrogressed into chlorite (Fig. 2d). It remains unclear whether small elongated plates of chloritoid parallel to S1 represent a first generation, or result from a preferential crystallization of chloritoid conformable with the early planar structure. No chemical difference is observed between the two possible generations of chloritoid. Chloritoid contains MnO (0.5-1.15 wt%), but no recalculated Fe,O, (on the basis of stoichiometry). The X,,(Cld) (0.85-0.89) is systematically smaller than that of coexisting staurolite. Gurnet. Small dioblastic, equidimensional Garnet is common. It is an Mn-rich almandine (Alm73-63, Prp7-5, Sps29-19, Grs4-1). The rimward decrease in spessartine (from 0.29 to 0.19) is compensated by the correlative increase in almandine and pyrope, with the X,, remaining constant at about 0.92. There is no significant difference in the range of garnet composition whatever its structural position. Garnet is in textural equilibrium with andalusite, staurolite, biotite and chlorite. Chlorite. Chlorite occurs mainly in the quartz-rich layers. Tiny crystals are exceptional within S 1, larger ones occur mainly in crystallization tails of andalusite, staurolite and garnet. Chlorite is also present in the oval-shaped inclusions within andalusite. Isolated grains form at the expense of biotite and/or chloritoid. Chlorite is ferrous with an X,., (0.60 up to 0.68) systematically slightly higher than that of biotite in the same thin section. However, chlorite occurring within andalusite inclusions has an X,, (0.53-0.59) lower than biotite and matrix chlorites. Oval-shaped inclusions within andalusite are composed of chlorite, muscovite, scarce biotite and quartz, and a very fine-grained mineral mass similar to pinite (Fig. 2a). This chlorite is more Mg-rich than chlorites in the matrix, supporting growth by destabilization of cordierite. Crystallization/deformation relationships We have seen that S1 is formed of muscovite, quartz, ilmenite, chlorite(?) and chloritoid(?). Cordieritec?), andalusite, biotite and garnet contain straight or only very slightly curved S1 inclusion trails and are thus post-S1. Staurolite contains inclusions of S1 crenulated by S2 and is therefore syn- to post-S2. As andalusite. biotite and garnet are in equilibrium with staurolite, and andalusite contains some staurolite as inclusions in the peripheral parts, we interpret these minerals as having grown as soon as the early stages of S2 development. They are also in textural equilibrium with S1 muscovite, ilmenite and quartz. Cordierite (contained only as inclusions in andalusite) is the oldest S2 mineral, followed by the crystallization of andalusite, biotite and garnet. Staurolite containing curved inclusion trails has grown longer than andalusite in which the S1 trails are straight. Accordingly, we assume a sequential, yet continuous, metamorphic crystallization (Fig. 3). Chloritoid and chlorite grow at the expense of biotite but are also contained within S2 shear-bands. Staurolite, chloritoid and chlorite have locally overgrown S2, but pre-date D3 structures. Therefore we postulate syn- to late-S2 peak metamorphism. Svratka Fe-rich Al-rich 0 . 0 -- 0 =- 111. 111. 53 White micu. White mica is fine-grained muscoviterich phengite (Si=3.02-3.08 for 11 oxygens) with an Na/(Na+K) ranging from 0.17 to 0.31. It defines S1 and is crenulated by S2. Muscovite is in equilibrium with all other mineral species. Hlinsko Grt Crd And St Bt Cld Chl PATH, BOHEMIAN MASSIF 0 0 Grt Crd And St Bt Cld Chl -El- I- -I --11. Grt KY St Rt Bt Sil Chl + Ms, Qtz, Ilm, Cr, fluid Fig. 3. Summary of the crystallization/deformation relationships within the hangingwall Hlinsko pelites and in the footwall Svratka mica schists. 54 1'. P I T R A & M . C U I R A U D MnKFMASHTC Paragenetic analysis A In a first approximation mineral parageneses can be described using the classical pelitic system K,O-FeOMg0-A1,0,-Si0,-H2O-TiO, (KFMASHT). Yet, Mn is abundant in garnet and in smaller but still significant amounts in staurolite, chloritoid and to a lesser extent also in other minerals and is therefore an independent component which must be taken into account in describing the phase relationships. Moreover, graphite is present as inclusions and in graphitic shales intercalated in the Silurian series. The interaction of graphite with water introduces C-bearing fluids into the hydrous fluid generated during metamorphism (e.g. Ohmoto & Kerrick, 1977). Hence CO, (kCH,) is an independent component, and the most appropriate system to describe the observed parageneses is the system MnKFMASHTC. In the eastern limb of the Hlinsko syncline, garnet occurs in all the parageneses and is considered to be in excess, along with quartz, muscovite, ilmenite, graphite and an H,O-CO, fluid. From the textural evidence, we deduce that the peak paragenesis is andalusite-cordierite-biotite near the centre of the syncline (followed by andalusite-staurolite-biotite) and staurolite-biotite at the base of the sedimentary sequence. Although staurolite began to grow in equilibrium with biotite, the latter is progressively chloritized, with the local appearance of chloritoid, and staurolite-chlorite is the typical association of this stage. Chloritoid is abundant in the lower part of the sequence and chloritoid-biotite-chlorite equilibrium may be exceptionally observed. This evolution is documented in the sequence of compatibility diagrams on Fig. 4. The peak parageneses are consistent with one another within the same compatibility diagram (Fig. 4) and therefore may reflect the same P- T conditions. Andalusite-cordierite-bearing assemblages occur in the most aluminous lithologies, in the centre of the syncline, whereas they are absent in the more ferriferous and less aluminous lithologies that predominate at the base: the observed difference in mineralogy would then reflect the difference in whole rock compositions. Petrogenetic grids and f - T conditions Standard geothermometry was used to put rough constraints on the temperature stability field of the observed parageneses. The garnet-biotite thermometer of Williams & Grambling (1990) was used because it allows the high Mn content in garnet to be taken into account. It yields reasonable temperatures ranging from 510 to 560 ^C. However, the poor analyses of biotite cast doubts on the reliability of these results. Garnet-chlorite thermometry (Dickenson & Hewitt, 1986; Ghent et al., 1987; Grambling, 1990) yields coherent temperatures between 490 and 545 "C. The geothermometer of Pownceby et al. (1987) based on the Fe-Mn exchange between garnet and ilmenite = 0 95 A1203 + 0 05 MnO = 0 95 FeO + 0 05 MnO ' = 0 95 MgO + 0 05 MnO +~rt +Ms a) +Qtz +Ilm +fluid "L A/F'+M' I Oxycilorite And c 1- F' F/F'+M' Bt Fig. 4. Compatibility diagrams representing two stages of the paragenetic evolution of the Hlinsko pelites. Projections have been calculated using the rim composition of garnet present in the rock. The composition of cordierite is that predicted by the therrnocalc program. The biotite bar depicts the theoretical biotite composition, empty points stand for biotite analyses in the Hlinsko schists. The shaded oval represents the cluster of 'biotite' (oxychlorite) compositions from one thin section. ChlCrdrepresents the composition of chlorite analysed within the andalusite inclusions. (a) Peak-S2; (b) late-S2. provides temperatures around 550-570 "C. Thus, the part of the P-T evolution including garnet can be bracketed between 500 and 600 "C. In order better to constrain the P-Tevolution of the Hlinsko pelites the petrogenetic grid approach using the thermocalc computer program (Powell & Holland, 1988; version 2.2b2, dataset April 1992) was chosen. This approach allows us (1) to confine the P-T conditions and to quantify the P-T evolution using textural criteria of equilibrium, the chemical equilibria being possibly broken during the P-T evolution, (2) to consider a system containing Mn in metamorphic phases and (3) to calculate the phase diagrams at aHZOimposed by the presence of graphite in the schists. In fact, the phase diagram for the metamorphic facies of interest strongly varies with respect to aHZOand ignoring this parameter could hamper the chosen approach. ANTICLOCKWISE P-T As shown by Connolly & Cesare (1993), the composition of the graphite-saturated C-0-H fluid produced by dehydration reactions is buffered by the condition of a constant H/O ratio, and is also uniquely determined at isothermal-isobaric conditions. Using the diagram of Ohmoto & Kerrick (1977), the X,,, ranges from 0.85 to 0.90 for the P-T area of interest (0.3-0.4 GPa, 500-600 C). Connolly & Cesare (1993) show that considering nonideality of mixing, these values should be higher. They find, however, their own values (0.90-0.95 for the same P-T field) are slightly overestimated. Thus, we consider that uH2, = 0.9 is a good approximation of fluid conditions reigning in the Hlinsko Silurian sequence. The petrogenetic grid for the system MnKFMASH has been calculated for the P-T-uH,, area of interest (Fig. 5 ) . Ideal mixing between mineral end-members was assumed, except for staurolite, where a mixing model based on Darken's Quadratic Formalism was used (Vance & Holland, 1993). Contents of Zn and Ti in staurolite are significant and should vary with the P-T conditions of equilibrium. In the absence of the Ti and Zn end-members it is not possible to calculate the effect of these components. Nevertheless, we use a constant value of Ti + Zn=0.2 (=50/0) for 4 atoms in PATH, BOHEMIAN MASSIF the M2 site and an activity corrected expression for staurolite a(fst)= 0.954 x4, where x = [Fe/( Fe + Mg)],,. Interpreting the observed mineral parageneses with the calculated P-T grid (Fig. 5) leads to constraints on the peak P-T conditions of 0.3-0.4GPa and 550-590 "C, based on the following reactions: 1 MnKFMASH reaction (Cld,Chl), since staurolitecordierite is not stable; 2 the lowermost stability of the biotite-andalusite is restricted to more than 0.3 GPa by the KFMASH reaction (Cld,Chl,St); 3 towards high pressures and temperatures, the limit is formed by And=Sil and the MnKFMASH reaction (Cld,St). In general, the differences in mineral parageneses depend as much on the bulk rock composition as on the P-T conditions of metamorphism. In order to explain the observed differences and constrain the P-T evolution of the Hlinsko pelites, P-T pseudosections were drawn for the two dominating bulk rock compositions: (a) more aluminous and less ferriferous lithologies from the centre of the syncline and (b) a more ferriferous and less aluminous lithology from the SE limb, near the tonalite (Fig. 6). The syntectonic crystallization of andalusite-stauro- 0.45 0.4 0.35 0.3 0.25 520 55 540 560 Fig. 5. Part of the P Tpetrogenetic grid for the KFMASH and MnKFMASH systems, calculated for a,,,=0.9. 56 P . P I T R A & M. C U I R A U D P (GPa) MnKFMASH pseudosection + Grt, Ms, Qtz, fluid ( a ~ 2 0= 0.9) Ik rock composition: I I I I 560 5s0 I 520 540 P (GPd) MnKFMASH pseudosection + Grt, Ms, Qtz, fluid (aH20 = 0 9) 0 sol R T(T) b) R 0 tst Cld Chl 0 15 ~~~ ~ ~~~ T y _- 520 ~1~ _., l---- ' ~~ 540 ! 560 - 580 T(T) Fig. 6. P-Tpseudoscctions calculated for the two main lithologies in the Hlinsko pelites yield complementary information about the P~-Tevolution.(a) The Al-rich rocks havc preserved the initial stages of the P-Tevoluiion, involving a slight increase in pressure; ( b ) the Fe-rich lithology has well recorded thc late stages. dominated by cooling, whereas the incrcase in pressure is inferred from (a). lite-biotite (preceded by andalusite-cordierite-biotite) and staurolite-biotite are contemporaneous in rocks with different compositions. As deduced from the compatibility diagrams, these parageneses are consistent with each other under the same P-T conditions. Comparing the two pseudosections, the peak P-Tmay further be limited to the high-pressure part of the andalusite field (say higher than 0.35 GPa) as cordierite does not occur in the Fe-rich Al-poor rocks. Moreover, a P-T evolution can be observed within the narrow pressure interval. The sequence of mineral assemblages observed in the aluminous rocks in the centre of the syncline is And-Crd-Bt > And-~ St-Bt > And-St-Chl. The crystallization of andalusite-staurolite-chlorite from andalusite-staurolitebiotite implies a path involving cooling from about 570 C to 550-530 C (stability of andalusiteestaurolite-chlorite). The observed crystallization of andalusite-staurolite-biotite replacing andalusite-cordieritebiotite precludes isobaric cooling only and requires a slight initial increase of pressure from about 0.36 GPa to 0.4 G P a according to the pseudosection on Fig. 6. The paragenetic sequence in the less aluminous and more ferriferous rocks, St-Bt > St-Chl(-Cld) > CldChl, records cooling at pressures higher than 0.35 GPa, as no trace of cordierite has been found in this rock type. The cooling is recorded to temperatures under 530 'C, in the chloritoid-chlorite assemblage stability field. The two different bulk rock compositions yield complementary information about the P-T evolution. The Al-rich rocks have better preserved the initial stages involving the slight increase in pressure. The late stages comprise cooling to low temperatures, and are better recorded in the more ferriferous rocks, abundant near the rim of the syncline. This may be related to the more intense fluid circulation near the major shear zone allowing diffusion processes to take place at lower temperatures. The X(Sps) values analysed in the garnets fit well with the P-T stability fields of our parageneses. However, calculated evolution of X(Sps) does not match the observed one. X ( Sps) should increase along the P-Tpath shown in Fig. 6, whereas it decreases from the garnet core to the rims. On the other hand, the small garnets do not display any synkinematic features, which suggests rather fast nucleation and growth. Consequently, the Mn zoning in garnet may not reflect a change in P-T conditions but rather diffusion processes in a reservoir with limited Mn content. This is consistent with the overall same composition range of garnet whatever its structural position. Our calculations confirm the observation of Pattison (1989, 1991) that the presence of graphite is often needed for rocks to display andalusite-bearing assemblages. In fact, lowering the uH20 shifts the reactions involving aluminosilicate to lower temperatures and into the andalusite stability field. The topology of the calculated MnKFMASH pseudosections is in excellent agreement with that of Hudson (1980) drawn for the KFMASH system. It shows that the major effect of adding Mn to the chemical system is only the stabilization of a spessartine-rich garnet, the original KFMASH topology being conserved. DISCUSSION An increase in pressure in the hangingwall Hlinsko schists during the main metamorphic event is the most interesting geological result. This increase is very slight and probably lies within the uncertainties of the calculated reactions. Moreover, the limits of stability are also sensitive to the bulk rock composition. Nevertheless, whatever the precise location of the equilibria in the grid, the relative position of the stability fields of the different parageneses remains reliable. The cordierite-out reaction has a very flat slope and therefore the disappearance of cordierite ANTICLOCKWISE corresponds to an increase in pressure within the andalusite field. It would be theoretically possible for cordierite to have crystallized within the sillimanite field and to have reacted out during isobaric cooling. However, no sillimanite is found in the rocks. Given the abundance of aluminosilicate in the Al-rich rocks, an early sillimanite would have been preserved. Therefore, we consider that cordierite grew only within the andalusite field. Although no cordierite could be analysed with the microprobe, there is strong evidence to indicate that the chlorite-muscovite fbiotite inclusions within andalusite represent former cordierite: 1 inclusions have a characteristic oval shape and optical attributes typical for pseudomorphs after cordierite; 2 chlorite is more Mg-rich in the inclusions than in the matrix (see Chl,,, in Table 1) and the calculated equilibrium compositions of the phases involved in the reaction Crd + Bt + And + H,O = Grt + Chl + Ms + Qtz at T= 560 ‘ C and P=O.39 GPa correspond to the compositions we measured in the inclusions. Therefore, it is assumed that the increase in pressure is realistic. In order to explain this P-T evolution, we have to consider the Hlinsko schists within their geodynamical context. The Hlinsko syncline lies between the underlying, NW-dipping high-grade rocks of the Svratka Crystalline Unit on the south-east and the granodiorite of the Nasavrky complex that forms its western, vertical to reverse limb (Fig. 1). Its uppermost part is formed by a nearly unmetamorphosed flysch-type (Rychmburk) sequence. Svratka Crystalline Unit In the Svratka Crystalline Unit, mica schists are intercalated within migmatites and migmatitic paragneisses that form nearly 80% of the unit. Previous studies have suggested that the Svratka metamorphics exhibit early planar and linear fabrics associated with high- to medium-pressure metamorphic conditions (Nemec, 1968; Pertoldova, 1986). They are reworked, showing a muscovite-sillimanite-biotite-bearing foliation associated with NW-dipping shear bands. Semibrittle NW-dipping shear zones post-date both structures and indicate the continuation of the deformation at low temperatures. To the East, the SCU is directly thrust over the Moravian nappe complex, where Devonian rocks are affected by Variscan thrusting and metamorphism. The Barrovian metamorphic zonation in both the Moravian nappes and the Svratka Crystalline Unit show apparent continuity of the thermal structure, i.e. identical dP/dT profile and synmetamorphic prograde mineral growth and kinematics (Schulmann et al., 1991). Thus, the early metamorphism of the overlying SCU is directly linked with the evolution in its footwall and should be Variscan in age. P-r P A T H , B O H E M I A NM A S S I F 57 Nemec (1968) deduced approximate peak metamorphic conditions of about 450-590 “C for around 1 GPa from mica schists of the central part. Pertoldova (1986) has studied the petrology of a skarn body in the SE part of the unit and has calculated peak P-Tconditions of 600 “C for 0.6-0.9 GPa. We have investigated several mica schist samples from the NW part of the SCU in order to verify these results. Mica schists range in composition from rare light quartz- and muscovite-rich mica schists to more common darker biotite- and sillimanite-rich ones. In the ‘dark’ mica schists the main foliation comprises an intercalation of layers rich in biotite (XFe = 0.57-0.7) intergrown with sillimanite and muscovite, and layers rich in quartz and plagioclase (XAn 10-13). This foliation wraps around scarce crystals of garnet (Alm83-78, Prp8-12, Sps3-10, Grs3-5), tourmaline or large plagioclase. Garnet is small (1-2 mm) and often forms atoll-like porphyroblasts that include quartz, plagioclase, muscovite and biotite crystals. In the light mica schists, mica-rich and quartz-rich layers define the main foliation. Mica-rich layers are dominated by muscovite intergrown with biotite and sillimanite and contain minor quartz, staurolite, kyanite, chlorite and ilmenite. Porphyroblasts of staurolite and kyanite are generally parallel to, but also deformed by, the foliation that wraps around large poikiloblastic garnets. Rutile and ilmenite are common inclusions but only ilmenite is observed in the matrix. Garnet forms large subeuhedral porphyroblasts (up to 3 mm) and contains inclusions of quartz, rutile (up to 0.1 mm), less abundant ilmenite, small muscovite and aluminosilicate. Optical zonation, characterized by a core rich in inclusions of large quartz grains, suggests two stages of garnet growth. Elongate sigmoidal trails of quartz inclusions within garnet cores demonstrate its synkinematic crystallization. The same structural position of staurolite (up to 3 mm, Fe/( Fe + Mg) = 0.85-0.90), kyanite and garnet (Alm80-87, Prp8-14, Sps4-1, Grs8-1) with respect to the main foliation suggest their contemporaneous growth during a first metamorphic event. The main foliation (muscovite-biotitesillimanite) wraps around large garnets, meaning it developed later; yet, biotite (XFe= 0.57-0.64) is in textural equilibrium with the minerals of the first paragenesis. The last, retrogressive stage is documented by the alteration of garnet, staurolite. kyanite and biotite to chlorite (XFe= 0.60-0.64) and muscovite. We deduce from the textural evidence that the first recorded paragenesis is garnet-kyanite-staurolite-biotite, followed by the crystallization of garnet-sillimanite-biotite (Fig. 3). The presence of different Al,SiO, polymorphs in the two successive parageneses shows that they have equilibrated at significantly different P-T conditions. Applying standard geothermometers yields unreasonably scattered data, resulting probably from insufficient equilibration of phases coexisting in textural equilibrium. However, garnet-sillimanitebiotite is stable within the domain limited by the 58 P. P l T R A & M. C U I R A U D KFMASH reactions Sil+ Grt = Bt + St and Sil+ Bt = Crd+Grt (+Ms, Qtz. H,O). Taking into account the widespread occurrence of migmatites in the region, and the fact that muscovite is stable in the mica schists, we can confine the temperature at around 600-650 T. According to the KFMASH petrogenetic grid of Xu et al. (1994), calculated using the same version of thermocalc,we can thereafter bracket the pressure between 0.4 and 0.6 GPa. Garnet-kyanite-staurolite-biotite is a univariant paragenesis in the KFMASH system. However, significant amounts of Ca in garnet suggest that garnet could be stabilized by this additional component. The GRAIL barometer (Bohlen et ul., 1983) was used because garnet contains rutile, ilmenite and aluminosilicate a s inclusions and because it is largely independent of temperature estimates. It yields pressures between 0.8 and 0.9 GPa. which implies temperatures from 610 to 660 C (KFMASH petrogenetic grid of Xu et al., 1994), consistent with the lack of evidence of melting at this stage. The two metamorphic events recorded within the SCU occurred at nearly the same temperature but at clearly different pressures. The late crystallization of chlorite suggests temperatures lower than 570 -C at the end of metamorphism. Thus, the rocks of the SCU record a decompression of at least 0.2 G P a at a relatively stable peak temperature, followed by cooling. Such an evolution is known in the footwalls of extensional terranes (e.g. Chauvet et al., 1992; Rey et ul., 1992). It is generally explained as the consequence of the rapid exhumation of rocks underlying major extensional faults, and thus the minor increase in pressure seen in the Hlinsko schists is not related to the history of thc SCU. Extensional tectonics within the Hlinsko-Svratka region took placc during the first period of late Variscan extension (Pitra et ul., 1994) that began in the inner part of the belt whilst convergence was still active (Burg et ul., 1994a). Therefore, we may expect that there was no major time gap between the first, crustal-thickening-related metamorphic event and the extension-related decompression of the Svratka Crystalline Unit. Numerical modelling shows that only fast exhumation rates ( > 2 km Myr-') provide an explanation for a stage of isothermal decompression (e.g. Rey et d..1992). Similar exhumation rates, reaching 3-6 km Myr-', were measured in recent extending orogens (Dallmeyer et ul., 1986; Davis, 1988; Hacker el ul., 1990; Lonergan & Mange-Rajetzky, 1994). Thus, a decompression of 0.2-0.4 GPa. corresponding to an uplift of 6-12 km, may be accomplished within 1L4 Myr. Nasavrky granodiorite The Nasavrky granodiorite intrusion produced a limited contact metamorphism of all series of the Hlinsko syncline (including the otherwise unmetamor- phosed Rychmburk sequence) and a deformation that clearly post-dates structures related to the peak metamorphism. Hence, its intrusion cannot explain the increase in pressure recorded in the Hlinsko schists. Rychmburk flysch basin The geological meaning of this pressure increase has therefore to be sought in the syntectonic sedimentation of the over 2-km-thick (Vachtl, 1962) flysch-type Rychmburk sequence. Sedimentary filling of flysch character with conglomerate beds that concentrate near the boundary fault is common in continental halfgrabens (e.g. Allen & Allen, 1990). It is recognized in the Rychmburk sequence (Vachtl, 1950), which also appears to have been deposited in a half-graben during normal faulting. High subsidence rates are commonly observed in half-grabens (e.g. Allen & Allen, 1990; Gordon & Heller, 1993), as well as related rapid unroofing rates of their footwalls. We have to consider a fast sedimentation associated with a decreasing heat supply from the cooling tonalite and Svratka migmatites in order to account for the increase in pressure at constant or even slightly decreasing temperature. Schlische & Olsen (1990) argue, on the basis of a onedimensional numerical model, that a half-graben may P l.o[ (GPa) Svratka (Mold an u bian) I I I 500 I I I I I 600 I I I I I 700 Fig. 7. Schematic P Tpaths for the Hlinsko pelites and the mica schists of the Svratka Crystalline Unit accompanied by a cartoon representing the tectonic evolution of both units. A crustal-scale normal ductile shear zone enhanced the emplacement of the syntectonic pluton that heated the overlying schists. The footwall Svratka rocks were exhumed, whereas the Hlinsko schists in the hangingwall were buried under a thick syntectonic basin, which would explain the recorded increase in pressure during the peak metamorphism. A N T I C L O C K W I S E P-T P A T H , BOHEMIAN MASSIF reach a cumulative thickness of more than 2 km in its deepest part within 3 Myr. Their calculations are consistent with field observations in the Mesozoic rift basins of eastern North America (Schlische & Olsen, 1990) and the late Variscan Saar-Nahe basin (Henk, 1993). These values are compatible with the exhumation rates recorded in extensional footwalls. This provides a plausible link between the metamorphic and sedimentary histories of the Hlinsko region and a constraint on the estimation of the metamorphism duration. CONCLUSIONS The Hlinsko area is characterized by a ductile normal fault that separates the footwall Svratka gneisses from the pelites of the Hlinsko syncline in the hangingwall. The main metamorphic crystallization in both units (staurolite-biotite i andalusite in the Hlinsko schists, garnet-sillimanite-biotite in the SCU) is related to concordant extensional movements and is consequently simultaneous. However, the footwall a n d the hangingwall units display opposite, but convergent, P-T histories (Fig. 7 ) . Decompression in the footwall rocks is related to rapid exhumation. We propose that the inverse, anticlockwise, P-T path recorded in the hangingwall pelites is related to the rapid, extensioncontrolled sedimentation of the overlying (Rychmburk) flysch series. Although, to our knowledge, such a P-T-t path has not yet been demonstrated in similar extensional terranes, it may be common, as syntectonic basins and fast subsidence are usually observed in the hangingwall of large extensional faults. ACKNOWLEDGEMENTS L. Latouche and R. Black helped to improve the manuscript. We are especially grateful to J.-P. Burg for a complete critical review. K. 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