Petrology and geodynamical interpretation of mantle xenoliths from Late Cretaceous lamprophyres, Villány Mts (S Hungary)

Tectonophysics 489 (2010) 43–54

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Petrology and geodynamical interpretation of mantle xenoliths from Late Cretaceous lamprophyres, Villány Mts (S Hungary) Zsuzsanna Nédli a,b, Tivadar M. Tóth b, Hilary Downes c, Géza Császár d, Andrew Beard c, Csaba Szabó a,⁎ a

Lithosphere Fluid Research Laboratory, Department of Petrology and Geochemistry, Eötvös University Budapest, Pázmány P. stny 1/c, 1117 Budapest, Hungary Department of Mineralogy, Geochemistry and Petrology, University of Szeged, Egyetem utca 2-4, 6722 Szeged, Hungary Department of Earth and Planetary Sciences, Birkbeck College, University of London, Malet Street, London WC1E 7HX, United Kingdom d Department of Regional Geology, Eötvös University Budapest, Stefánia út 14, 1143 Budapest, Hungary b c

a r t i c l e

i n f o

Article history: Received 27 May 2009 Received in revised form 26 February 2010 Accepted 21 March 2010 Available online 27 March 2010 Keywords: Villány Mts Tisza unit Mantle xenolith Thermobarometry Geodynamics

a b s t r a c t A Late Cretaceous lamprophyre dyke in the Villány Mts (S Hungary), situated in the Tisza unit, contains abundant spinel lherzolite xenoliths with porphyroclastic textures. Mineral chemistry suggests a relatively fertile mantle, which experienced only 5–7% melt extraction. Differences in porphyroclast and neoblast chemistry and thermobarometric calculations suggest that the mantle section represented by the xenoliths experienced recrystallization at lower PT as it was transported to shallow mantle depths close to the plagioclase stability field, followed by later relaxation. Based on volcanological and sedimentological constraints from the Villány Mts and the neighboring Mecsek Mts, we suggest that the uprise of the subcontinental mantle material was related to a Cretaceous rifting event and lithospheric deformation of the southwestern part of the Tisza unit. Mantle upwelling and formation of lamprophyre melts can be related to generation or reactivation of deep fractures of the lithosphere, during a period of lithospheric extension between the major nappe emplacements (Albian–Cenomanian and Paleocene) of the region. © 2010 Elsevier B.V. All rights reserved.

1. Introduction Mantle xenoliths are of great significance because they provide direct insight into the composition, physical state and evolution of the subcontinental lithospheric mantle. They are often hosted by alkaline mafic magmas but also occur rather rarely in lamprophyres. In the Carpathian–Pannonian Region xenolith-bearing lamprophyres of similar age (Late Cretaceous) are known on different microplates. A xenolith-bearing Late Cretaceous lamprophyre dyke swarm in the Alcapa unit (N Hungary) was discovered and described by Szabó et al. (1993, and references therein). Recently, other lamprophyre dykes were found in the Tisza unit, in the Villány Mts (S Hungary) (Nédli and M. Tóth, 2007) and one of them contains abundant, but altered mantle xenoliths. Upper mantle xenoliths from Neogene volcanics in the Carpathian–Pannonian region (Fig. 1) are common (predominantly from the Alcapa Unit) and have been studied in detail for decades, however such xenoliths are rare in Mesozoic volcanics in the region (Szabó, 1985; Downes et al., 1995). Therefore, the pre-Neogene subcontinental lithosphere beneath the region is poorly known. Thus, the recently discovered xenoliths in Late Cretaceous lamprophyre dykes of the Villány Mts (S Hungary) can provide important informa-

⁎ Corresponding author. E-mail address: [email protected] (C. Szabó). 0040-1951/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2010.03.013

tion to fill the gap of our knowledge about the pre-Neogene subcontinental lithosphere beneath the Tisza unit (Fig. 1). The petrology of the host lamprophyres was studied in detail by Nédli and M. Tóth (2007). The present study concerns the petrology, geochemistry and thermobarometry of spinel lherzolite xenoliths from the Beremend lamprophyre dyke (Fig. 1) and presents an attempt to integrate the information offered by these xenoliths into the lithosphere evolution of the Villány Mts region. 2. Geological background In the Alpine–Carpathian region, four major terrains have been distinguished: the Alcapa, Tisza, Dacia and Adria units (Fig. 1) (e.g. Kovács, 1984; Balla, 1988; Csontos, 1995; Fodor et al., 1999; Csontos and Vörös, 2004). The Tisza unit, now situated in the southern part of the Pannonian Basin, formed the northern (European) margin of the Mesozoic Tethys until the Middle Jurassic (Géczy, 1973). During the Jurassic and Cretaceous, several extensional basins opened along the southern margin of the European continental plate (Kázmér and Kovács, 1989), leading to the separation of the Tisza terrain from the European plate in Middle Jurassic times (Plašienka, 2000; Marroni et al., 2002) by anti-clockwise rotation (Balla, 1987). By the Late Cretaceous, the oceanic branches of the north-western Tethys were closed and reorganization of the microplates resulted from the collision between the African and European plates. The Alcapa and Tisza

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Z. Nédli et al. / Tectonophysics 489 (2010) 43–54

positions of pyroxenes we performed mapping of lamellar pyroxenes, using scanning with a defocused beam, on areas of ca. 50 μm2. Laser ICP-MS analysis of single clinopyroxenes for trace elements was performed at Birkbeck College, University of London. Clinopyroxene grains would be large enough for ICP-MS analysis, however it was impossible to find an appropiately large clinopyroxene in the samples because grains are often altered along cracks and fractures. Therefore, two hand-picked clinopyroxene single crystals, the same ones used for structural analysis, were used for ICP-MS analysis but only one crystal provided valuable analytical results. The laser-ICP-MS at Birkbeck/UCL consists of a New Wave Research UP213 laser aperture imaged frequency quintupled Nd:YAG solid state laser source operating at a wavelength of 213 nm, coupled to an Agilent 7500a quadrupole ICP-MS. Time-resolved analyses were performed using a laser diameter of 55 µm at pulse frequency of 20 Hz for a laser dwell time of 20 s. Instrumental background levels were established by a ‘gas blank’; i.e., analysis of the He/Ar mixed gases with the laser off for 8 s. Ablation was carried out in a He atmosphere and mixed with Ar carrier gas before the plasma touch. The synthetic glass reference material NIST 612 was used as the calibration standard, with the average composition of Pearce et al. (1997) being used in this study. Calcium (43Ca) was used as the internal calibration to correct for differences in ablation characteristics between samples and standards. The raw data was processed using the GEMOC glitter reduction program. On one crystal two measurements were possible to carry out. Fig. 1. Geological sketch map and major terranes of the Carpathian–Pannonian Region after Csontos and Vörös (2004) with localities of xenolith-bearing Mesozoic volcanics in the region (AD = Alcsútdoboz dyke swarm, Szabó, 1985; Poiana Rusca, Downes et al, 1995; Villány Mts, Nédli and M. Tóth, 1999 and this study). Inset: Late Cretaceous lamprophyre and xenolith localities in the Villány Mts (S Hungary).

terrains, rotating in opposite senses, came into juxtaposition during the Late Oligocene–Early Miocene (Csontos and Vörös, 2004). In the Villány Mts dykes of lamprophyric composition sporadically crosscut Mesozoic limestones (Nédli and M. Tóth, 2007) (Fig. 1, inset). They contain abundant pyroxene xenocrysts and the Beremend lamprophyre dyke is rich in spinel lherzolite xenoliths (Nédli and M. Tóth, 1999). Whole rock K/Ar data for the dykes indicate a Late Cretaceous age (76 Ma, Molnár and Szederkényi, 1996). Based on their geochemical characteristics, these dykes originated in a withinplate geotectonic setting from an enriched garnet lherzolite mantle source by low degree (approx. 1%) partial melting. Trace element geochemistry suggests that this source corresponds to asthenospheric mantle that was previously enriched by subduction-related melts/ fluids (Nédli and M. Tóth, 2007). Enriched mantle component and metasomatized mantle xenoliths are also known from the Late Cretaceous NE Transdanubian lamprophyre dykes (Szabó, 1985; Szabó et al., 1993) from the Alcapa microplate (Fig. 1). These two, geochemically similar, mantle xenolith-bearing lamprophyre localities of almost the same age and located far from each other, according to geodynamic reconstructions (Csontos and Vörös, 2004), offer the opportunity to study the nature and evolution of the late Mesozoic mantle at a large scale.

3. Analytical techniques Mineral chemical analyses on fresh phases were obtained using a JEOL 733 superprobe at Birkbeck College, University of London. Data were collected at 15 kV, 20 nA beam current, for 100 s per analysis, using an AN 10.000/55 s Link energy dispersive system and 1 μm spot diameter. Natural and synthetic minerals were used as standards. Exsolution lamellae in pyroxenes were avoided to determine host phase composition. In an effort to reconstruct pre-exsolution com-

4. Petrography 4.1. Host rock The Beremend lamprophyre is a 2–3 m thick subvertical dyke, which outcrops in a limestone quarry, crosscutting Aptian–Albian carbonate sediments. The dyke consists of olivine and clinopyroxene phenocrysts which form glomeroporphyritic aggregates, 3–4 mm in size. The olivine phenocrysts are 1.0–1.5 mm large, euhedral, and totally replaced by secondary minerals forming a typical mesh structure. Two clinopyroxene generations are distinguished: a) zoned phenocrysts, 1.5–2.0 mm in size with partially or totally resorbed cores, overgrown by fresh, unaltered clinopyroxene rims; b) euhedral prismatic augite, 0.1–0.5 mm in size, makes up 20–30 vol.% of the rock volume, in places surrounded by amphibole rims. Plagioclase is present only in the groundmass as small (0.1–0.5 mm) partially altered laths, which compose 10–30 vol.% of the rock. The groundmass also contains euhedral amphibole and biotite crystals, 0.1–0.2 mm large (1–2 vol.%), Fe–Ti-oxide grains, 0.3–0.5 mm large (8–10 vol.%), and apatite needles (1–2 vol.%) with a very high length to width ratio (1:50–1:100). Secondary minerals (chlorite, clay minerals and calcite) are also present in the matrix in variable amounts (up to 20–30 vol.%). Felsic, globular ocelli occur frequently. They are ellipsoidal or spherical and 2–15 mm in diameter. Most ocelli consist of an anhedral calcite core surrounded by a complex rim of silicate minerals (Nédli and M. Tóth, 2007). 4.2. Xenoliths The Beremend lamprophyre carries ultramafic xenoliths, which are rounded or ovoid, 1–5 cm in diameter. They are predominantly spinel lherzolites consisting of enstatite, Cr-diopside and Mg–Al chromite as primary phases. Although no fresh olivine could be detected, its previous presence as a dominant primary phase is indicated by the general mesh texture of the xenoliths. All xenoliths contain abundant secondary phases such as Fe-oxides, calcite, serpentine and smectite. The xenoliths contain large (3–4 mm) orthopyroxene porphyroclasts, disseminated among fine-grained, equigranular olivine, orthopyroxene and clinopyroxene grains (Fig. 2a). The majority of

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Fig. 2. Characteristic petrographic features in the Villány Mts xenoliths. 2a: large, deformed orthopyroxene porphyroclast surrounded by smaller, undeformed clinopyroxene neoblasts, cross-polarized photomicrograph; 2b; rounded spinel inclusions in clinopyroxene, plane polarized photomicrograph; 2c: spinel lamellae in orthopyroxene porphyroclast suggests re-equilibration of xenoliths, backscattered electron image; 2d: spongy rimmed spinel porphyroclast, backscattered electron image (sp = spinel, opx = orthopyroxene, cpx = clinopyroxene).

boundaries among the fine-grained, strain-free grains are straight. Well-developed 120° junctions in triple points are typical, showing textural equilibrium of constituent minerals. Spherical spinel inclusions are abundant in silicates (Fig. 2b). The xenoliths can be classified in the secondary recrystallized porphyroclastic group, according to Mercier and Nicolas (1975). Altered olivine occurs as large (1–4 mm) porphyroclasts, made up of a mesh structure, consisting of amorphous SiO2-domains and

secondary minerals of Fe-oxides, calcite and serpentine. Some relicts contain small spinel inclusions. Orthopyroxene forms porphyroclasts (3–4 mm) and smaller neoblasts (1–2 mm). Porphyroclasts contain abundant acicular or blebby lamellae of spinel or clinopyroxene (Fig. 2a, c), whereas neoblasts always lack lamellar texture. Clinopyroxene also occurs bimodally, but mainly as small neoblasts (1–2 mm) (Fig. 2a, b) situated interstitially among other silicate minerals. Some grains however are larger, amoeboidal and contain

Table 1 Clinopyroxene chemistry of Villány Mts xenoliths. mg#=Mg/(Mg + Fe).

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Total mg#

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Total mg#

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be41

Be43

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

53.80 0.23 1.71 0.27 2.18 0.08 17.32 24.36 0.13 100.08 0.93

53.64 0.14 2.04 0.26 2.24 0.07 17.35 24.15 0.16 100.04 0.93

53.69 0.19 1.81 0.50 2.23 0.09 17.33 24.57 0.09 100.50 0.93

54.12 0.24 2.01 0.51 2.20 0.13 17.65 24.53 0.15 101.55 0.93

53.84 0.29 1.89 0.30 2.23 0.05 17.25 24.54 0.11 100.50 0.93

54.25 0.28 1.89 0.33 2.16 0.11 17.37 24.36 0.11 100.87 0.93

54.10 0.24 1.76 0.35 2.13 0.14 17.36 24.46 0.14 100.67 0.94

53.76 0.23 1.82 0.45 2.28 0.06 17.11 24.50 0.11 100.30 0.93

54.11 0.20 1.96 0.35 2.21 0.08 17.50 24.46 0.16 101.04 0.93

53.73 0.24 1.88 0.34 2.21 0.10 17.26 24.48 0.13 100.37 0.93

53.81 0.16 1.85 0.27 2.29 0.21 17.11 24.44 0.12 100.25 0.93

54.08 0.20 1.85 0.31 2.15 0.06 17.39 24.46 0.13 100.64 0.94

53.49 0.21 2.40 0.34 2.33 0.07 17.08 24.04 0.18 100.14 0.93

Be43

Be43

Be43

Be43

Be43

Be46

Be46

Be46

Be46

Be42

Be42

Be46

Be46

Neoblast

Neoblast

Neoblast

Neoblast

Neoblast

Porph

Porph

Porph

Porph

Porph

Porph

Area1

Area2

53.33 0.26 2.21 0.35 2.34 0.13 17.10 24.01 0.17 99.91 0.93

53.50 0.28 2.42 0.42 2.30 0.09 17.20 24.05 0.22 100.47 0.93

53.21 0.27 2.42 0.33 2.21 0.12 17.02 23.97 0.21 99.75 0.93

52.98 0.22 2.88 0.42 2.20 0.11 16.96 23.73 0.23 99.73 0.93

53.14 0.29 2.40 0.43 2.32 0.15 16.89 23.87 0.20 99.68 0.93

52.87 0.28 3.16 0.53 2.72 0.00 16.60 23.38 0.43 99.98 0.92

52.90 0.34 3.13 0.44 2.62 0.00 16.64 23.39 0.44 99.89 0.92

52.77 0.21 3.15 0.55 2.63 0.00 16.53 23.32 0.36 99.54 0.92

52.81 0.27 3.13 0.43 2.71 0.00 16.53 23.27 0.31 99.46 0.92

52.28 0.26 3.65 0.66 2.76 0.08 15.94 22.45 0.79 98.86 0.91

52.32 0.32 3.61 0.58 2.79 0.06 15.96 22.21 0.71 98.55 0.91

50.29 0.32 4.65 1.02 3.01 0.17 16.59 22.53 0.51 99.10 0.91

50.14 0.36 5.32 1.00 3.01 0.18 16.40 22.30 0.44 99.16 0.91

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abundant, minute (10–50 μm) acicular or blebby lamellae of orthopyroxene. Exsolution features occur both in ortho- and clinopyroxene; they are lamellar or blebby, about 20–50 μm in length, representing approx. 5-10 vol.% (Fig. 2c). Orientation of exsolution features generally varies between two planes at 20–30°. Spinel is present in 3 different textural positions: a) in the form of large (2– 3 mm), amoeboidal porphyroclasts situated interstitially among the silicate minerals, sometimes showing vermicular textured rims (Fig. 2d), b) as small, rounded inclusions in silicates (Fig. 2b) and c) as lamellae in relict porphyroclasts (Fig. 2c). 5. Mineral chemistry 5.1. Clinopyroxene We carried out detailed microprobe measurements, analysing 6– 10 points on virtually fresh, unzoned minerals. Clinopyroxenes in different textural positions were analysed: neoblasts wich lack exsolution lamellae, and porphyroclasts which contain large amounts of minute lamellae. Pre-exsolved composition of porphyroclasts was also analysed by defocused beam measurements. Clinopyroxene chemistry data are shown in Table 1. Neoblast and porphyroclast chemistry is slightly different (Fig. 3). Neoblasts show higher mg# (Mg/(Mg+ Fe)) with values in a narrow range between 0.92 and 0.94. Na2O (0.1–0.9 wt.%) and TiO2 (0.1–

0.5 wt.%) contents of both pyroxene groups are low and vary in a narrow range, whereas Cr2O3 varies in a wider range (0.2–0.7 wt.%). SiO2 and CaO contents are higher in neoblasts than in porphyroclasts, Al2O3, Cr2O3 and Na2O are lower. Porphyroclast have mg# between 0.90 and 0.92, which are similar to those of their pre-exsolution composition (0.90–0.91), obtained by mapping of lamellar grains. Porphyroclasts pre-exsolved compositions show lower SiO2 (50– 51 wt.%) and CaO (22 wt.%) but higher Al2O3 (4.3–5.3 wt.%), FeO (3 wt.%) and Cr2O3 (0.7–1.0 wt.%) than porphyroclasts. We compared the Villány Mts xenoliths' chemistry to clinopyroxenes from petrographically similar porphyroclastic or recrystallized xenoliths from the region (Bakony-Balaton Highland Volcanic Field (BBHVF) (Downes et al., 1992; Embey-Isztin et al., 2001), Poiana Rusca (Romania) (Downes et al., 1995), E-Serbia (Cvetković et al., 2004) and from W-Eifel (Witt and Seck, 1987). From this latter locality petrographically identical xenoliths were described. Our samples are rather low in Cr2O3, Al2O3 and Na2O relative to other suites, however they have high mg#. Compositionally, they are similar (Fig. 3) to mosaic and poikilitic textured xenoliths from the BBHVF and to W-Eifel sheared porphyroclastic xenoliths. Clinopyroxene in xenoliths from Poiana Rusca are higher in Na2O and Al2O3, whereas those from E-Serbia are lower in Al2O3. Compositional trends from the porphyroclasts to neoblasts in the Villány Mts xenoliths (decrease in Na, Cr, Al and increase of mg#) coincide with trends in the sheared porphyroclastic xenoliths from W-Eifel (Witt and Seck, 1987). Porphyroclasts from Villány Mts xenoliths plot close to the poikilitic xenoliths from the BBHVF, whereas neoblast compositions are similar to clinopyroxene from the mosaic textured xenoliths from the same localities. REE abundanes in clinopyroxene of the studied xenolits are generally low (Table 2), the REE pattern of the analysed single crystal (Fig. 4) is flat for heavy REEs, but generally MREE- and LREE-depleted relative to HREEs (LaN/YbN = 0.2 and SmN/YbN = 0.4), and show a slight La-enrichment (LaN/CeN = 3.2). This REE pattern is rather more depleted than those of well-described xenolith series in Neogene alkali basalts from the Carpathian–Panonnian region (Downes et al., 1992), but is very similar to the “tick-shaped” REE-patterns of some clinopyroxenes from Massif Central (Downes et al., 2003) and the Rhön (Germany) (Witt-Eickschen and Kramm, 1997). Extended trace element pattern show notable Ta and Nb negative anomalies and slight Nd positive anomaly. Pattern is similar to that of clinopyroxenes from Rhön (Germany) and the Massif Central (France) apart from their more significant heavy incompatible element (Th, U, Sr and LREEs) enrichment and strong Zr-depletion. 5.2. Orthopyroxene As for clinopyroxene, we analysed orthopyroxene neoblasts and porphyroclasts separately. Pre-exsolved analyses with a defocused

Table 2 LA ICP-MS analysis (in ppm) of clinopyroxene single crystal from Villány Mts xenolith. agk 19.3

Fig. 3. mg# vs. Al2O3 and Na2O vs. Cr2O3 (wt.%) diagrams for clinopyroxenes from Villány Mts xenoliths. Comparative data after: Serbia — Cvetković et al. (2004); W-Eifel — Witt and Seck (1987); Poiana Rusca — Downes et al. (1995); CPR (Carpathian Pannonian Region) — Embey-Isztin et al. (2001).

Th U Nb Ta La Ce Pr Sr Nd Zr Hf

0.04 0.07 0.04 0.00 0.37 0.30 0.05 4.38 0.59 1.94 0.19

agk 19.3 Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu

0.47 0.22 0.98 0.24 1.86 11.31 0.45 1.25 0.21 1.24 0.19

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47

Fig. 4. Chondrite-normalised extended trace element and REE diagrams for Villány Mts xenolith clinopyroxene. Comparative data of xenoliths from French Massif Central (Downes et al., 2003) and Rhön volcanic field, Germany (Witt-Eickschen and Kramm, 1997). Normalising values after Taylor and McLennan (1985).

beam were also carried out on porphyroclasts (Table 3). All orthopyroxenes show homogeneous enstatitic composition (Wo1En89Fs10,) although small differences are observable: porphyroclasts are slightly richer in Cr2O3 but lower in MgO, SiO2 and TiO2 than neoblasts (Fig. 5). Mg# is high with values of 0.89–0.91 in all groups. CaO is low (0.41– 0.50 wt.%), Al2O3 ranges between 2.2 and 3.3 wt.% showing no significant difference between the two textural types. Pre-exsolved compositions show slightly higher Al2O3 (4.8–4.9 wt.%), Cr2O3 (0.6 wt.%) and CaO (0.8 wt.%) contents (Fig. 5). Orthopyroxenes from Villány Mts xenoliths are similar in composition to porphyroclasts and neoblasts of the sheared porphyroclastic xenoliths from the W-Eifel (Witt and Seck, 1987) and plot close to the field of mosaic-poikilitic xenoliths from the BBHVF (Downes et al., 1992; Embey-Isztin et al., 2001) (Fig. 5). Similarly to the W-Eifel xenoliths, neoblasts are lower in Al and Cr, but have higher mg#. Poiana Rusca (Downes et al., 1995) xenoliths show a wide range in composition, some samples resemble the chemistry of the Villány Mts xenoliths; whereas E-Serbia xenoliths (Cvetković et al., 2004) are highly Al-poor and rather Cr-rich relative to the other series. On the mg# vs. Cr2O3 diagram (Fig. 5) there is a strong similarity between the pre-exsolution composition of Villány Mts orthopyroxene porphyroclasts and poikilitic samples from BBHVF. 5.3. Spinel We analysed spinel porphyroclasts and inclusions in silicates, as well as exsolution lamellae in pyroxenes (Table 4). The composition of texturally different grains varies in narrow ranges (Fig. 6): they have low content of Cr2O3 (13–17 wt.%) and TiO2 (0–0.2 wt.%) but are high

in Al2O3 (49–52 wt.%). Low cr# (Cr/(Cr + Al)) (0.17–0.25) is homogeneous in samples; mg# (0.69–0.72) is also homogeneous and lower than in silicate minerals. Slightly higher mg# (0.72–0.74) is found only in spinel inclusions. Spinel porphyroclasts are homogeneous, only minor differences are detectable: slight rimward decrease in Al and increase in Cr content accompanied by constant Mg and Fe. In composition (low cr#, low TiO2, high Al2O3) Villány Mts spinels are similar (Fig. 6) to slightly depleted mantle spinels in xenoliths from Poiana Rusca (Downes et al., 1995), to those of mosaic samples from BBHVF (Downes et al., 1992; Embey-Isztin et al., 2001) and to spinel in W-Eifel sheared xenoliths (Witt and Seck, 1987). Xenoliths from ESerbia (Cvetković et al., 2004) show very scattered compositions, some are more depleted than those of Villány Mts and others have been more enriched by Ti. 6. Discussion 6.1. Mafic melt extraction from the subcontinental mantle and enrichment process Different degrees of melting of the subcontinental mantle are traceable in the chemical compositions of minerals in xenoliths: e.g. cr# of spinel and mg# of clinopyroxene are considered sensitive indicators of the extent to which mantle peridotites have lost their basaltic components (Frey and Prinz, 1978; Dick and Bullen, 1984; Press et al., 1986). Mg# around 90 in the pyroxenes suggests a slightly depleted mantle. Also low cr# and mg# of spinels, 0.17–0.25 and 0.69–0.72, respectively, (Fig. 6) resemble values in fertile spinel lherzolites, indicating that the mantle beneath the region has been

Table 3 Orthopyroxene chemistry of Villány Mts xenoliths. mg# = Mg/(Mg + Fe).

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Na2O Total mg#

Be41

Be41

Be42

Be42

Be42

Be42

Be42

Be42

Be42

Be42

Neoblast

Neoblast

Neoblast

Neoblast

Porph

Porph

Porph

Porph

Porph

Area

57.20 0.09 2.28 0.22 6.48 0.23 34.44 0.42 0.27 101.64 0.90

56.82 0.07 2.26 0.24 6.41 0.23 34.07 0.41 0.16 100.69 0.90

56.35 0.12 3.36 0.17 6.69 0.12 33.78 0.58 0.24 101.41 0.90

55.66 0.07 3.04 0.21 6.43 0.19 33.12 0.48 0.22 99.43 0.90

54.90 0.00 3.29 0.36 6.68 0.05 32.68 0.47 0.27 98.71 0.90

54.69 0.10 3.11 0.44 6.70 0.16 32.86 0.48 0.24 98.78 0.90

54.83 0.04 3.24 0.42 6.67 0.16 32.39 1.16 0.25 99.18 0.90

55.48 0.06 3.21 0.27 6.73 0.18 33.27 0.50 0.26 99.96 0.90

54.77 0.11 3.16 0.35 6.39 0.17 31.52 1.96 0.25 98.71 0.90

54.18 0.00 4.78 0.64 6.75 0.21 32.54 0.83 0.00 99.93 0.90

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estimations and show that our xenoliths went through low degree (approx. 4.5–7%) of basaltic melt extraction. Xenolith clinopyroxene is rather poor in REE-s, showing 0.2– 5 × chondritic values. The steeply convex chondrite–normalized pattern (except La) (Fig. 4) and low LREE/HREE ratio (LaN/YbN = 0.2; CeN/YbN = 0.06) indicates the LREE-depleted character of xenoliths, whereas the heavy REEs are rather flat. This indicates a low degree of melting, which removed the light (and partly the middle) rare earth elements from the mantle and did not affect the HREEs. Notable is the slight La enrichment relative to the Ce and other LREEs of the sample (LaN/CeN = 3.2) which points to a La enrichment after the LREE depletion. Similar La addition to LREE-depleted xenoliths was described from other suites in Europe (e.g. Downes and Dupuy, 1987; Vannucci et al., 1994; Vaselli et al., 1996; Zangana et al., 1997) and can be considered as extremely weak cryptic metasomatism. However, because of the small dimension and altered character of the xenoliths in study, we cannot exclude the possibility of some La-contamination by the host magma during ascent. 6.2. Thermobarometric estimations of upper mantle evolution

Fig. 5. mg# vs. Al2O3 and Cr2O3 diagrams for orthopyroxenes from Villány Mts xenoliths. Comparative data after: Serbia — Cvetković et al. (2004); W-Eifel — Witt and Seck (1987); Poiana Rusca — Downes et al. (1995); CPR (Carpathian Pannonian Region) — Embey-Isztin et al. (2001).

subjected to only low degrees of mafic melt extraction. On the TiO2 vs cr# diagram (Fig. 6), for comparison, partial melting line of a depleted MORB mantle (DMM) source after Pearce et al. (2000) is also shown. However, xenolith data from Villány Mts does not fit with the melting trend, although the low cr# of the samples suggests a low degree of melting (5–10%) of the mantle. Estimation of basaltic melt extraction for the studied xenoliths is also possible by use of a simple calculation, based on cr# in spinels (Hellebrand et al., 2001) (see F values in Table 4). These F values are also in good agreement with the above

Thermometry of mantle xenoliths constrains the thermal state and evolution of the mantle. Pyroxene geothermometry is a wellestablished method of estimating equilibrium temperatures in mantle xenoliths. We calculated equilibration temperatures (Table 5) using widely accepted different thermometers (Wood and Banno, 1973; Wells, 1977; Mercier, 1980; Brey and Köhler, 1990) on chemically homogeneous pyroxenes, which are thought to have equilibrated under mantle PT conditions and escaped re-equilibration after entrainment in the melt. Similar to other studies on porphyroclastic xenoliths (e.g. Bohrson and Clague, 1988; Falus et al., 2007), we consider that the earlier mantle conditions were preserved in porphyroclastic relicts. Their abundant exsolution features suggest that they suffered re-equilibration at lower PT conditions. Therefore, we also calculated temperature on porphyroclast compositions before their exsolution (referred to as pre-exsolution porphyroclast composition), which are thought to preserve an earlier mantle condition. After re-equilibration of this porphyroclastic assemblage, the mantle section experienced deformation(s) and the final re-equilibrated state is represented by finer-grained neoblasts. Fe–Mg and Cr–Al distributions among the pyroxene generations show slight but clear disequilibrium (Figs. 3, 5); such small differences indicate slightly different chemistry and hence physical conditions. The Cr2O3 and Al2O3 in ortho- and clinopyroxene generations show similar behavior (Fig. 7). Regarding the pyroxene generations, there is a notable compositional similarity between the studied xenolith and the W-Eifel sheared porphyroclastic xenoliths (Witt and Seck, 1987), which could suggest a similar thermal history.

Table 4 Spinel chemistry of Villány Mts xenoliths. mg# = Mg/(Mg + Fe); cr# = Cr/(Cr + Al); F = 10 ln(cr#) + 24, melt extraction in %, calculation after Hellebrand et al. (2001). Be41

SiO2 TiO2 Al2O3 Cr2O3 FeO MnO MgO CaO Total mg# cr# F (%)

0.20 0.08 50.98 16.29 14.28 0.00 18.46 0.05 100.33 0.70 0.18 6.65

Be41

0.11 0.06 52.38 14.70 13.09 1.40 18.97 0.12 100.83 0.72 0.16 5.57

Be41

0.17 0.08 51.95 14.85 14.02 0.00 18.59 0.03 99.69 0.70 0.16 5.73

Be41

0.22 0.05 52.00 15.16 13.54 0.00 18.73 0.10 99.82 0.71 0.16 5.89

Be41

0.30 0.00 52.31 14.92 13.80 0.00 18.88 0.03 100.24 0.71 0.16 5.71

Be42

0.22 0.08 51.94 13.37 13.73 1.17 18.81 0.08 99.40 0.71 0.15 4.84

Be43

0.25 0.05 52.08 14.48 13.69 0.39 18.81 0.07 99.82 0.70 0.17 6.54

Be43

0.23 0.06 52.01 13.93 13.71 0.78 18.81 0.08 99.61 0.70 0.17 6.12

Be43

0.24 0.06 52.05 14.20 13.70 0.58 18.81 0.08 99.72 0.70 0.17 6.44

Be43

0.24 0.06 52.05 14.20 13.70 0.58 18.81 0.08 99.72 0.72 0.15 5.24

Be46

0.18 0.12 52.07 12.80 14.80 0.00 19.12 0.07 99.14 0.70 0.14 4.44

Be46

0.21 0.08 51.86 13.86 15.32 0.00 19.16 0.05 100.53 0.69 0.15 5.16

Be41

Be41

Be41

Inclusion

Inclusion

Inclusion

0.23 0.02 52.14 14.45 12.47 1.16 19.41 0.32 100.20 0.74 0.16 5.47

0.27 0.05 53.92 14.15 13.18 0.00 19.76 0.23 101.55 0.73 0.15 5.00

0.18 0.08 51.68 14.88 13.78 0.00 18.56 0.24 99.41 0.71 0.16 5.79

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49

Fig. 6. mg# vs. cr# and TiO2 vs. cr# diagrams for spinels from Villány Mts xenoliths. Partial melting line after Pearce et al. (2000). Comparative data after: E-Serbia — Cvetković et al. (2004); W-Eifel — Witt and Seck (1987); Poiana Rusca — Downes et al. (1995); CPR (Carpathian Pannonian Region) — Embey-Isztin et al. (2001).

We applied Brey and Köhler's (1990) Ca-in-opx thermometer, which is widely accepted in mantle studies and this gave the closest values to the average of all thermometric formulations. Thermometry of porphyroclasts suggests that the mantle could have had an earlier temperature about 990–1010 °C (pre-exsolution porphyroclast composition). After cooling, accompanied by lamellae formation, the porphyroclasts reached equilibrium at 910–960 °C. Neoblasts represent the physical conditions after the final re-equilibration at 870– 920 °C. Absence of zoning in neoblast compositions indicates that equilibrium was nearly reached. Pressure estimation in the spinel stability field is a serious unsolved problem in mantle xenolith studies. Crystal structural studies on mantle xenolith clinopyroxenes (Dal Negro et al., 1984; 1989; Cundari et al., 1986) have shown that relations between unit cell and site (especially M1 and M2) volumes are indicative of different equilibrium pressures. These observations stimulated several mineralogists (e.g. Secco, 1988; Princivalle et al., 1994; 2000a, b) to use clinopyroxene structural parameters for estimation of equilibrium pressure for peridotitic assemblages. Therefore, we carried out clinopyroxene single crystal structural studies by combination of single crystal X-ray diffractometer and electron microprobe measurements on the studied Villány Mts xenoliths. For a detailed description of the method and results, see Nédli et al. (2008, 2009). We observed that the studied clinopyroxenes have rather large unit-cell volumes (437.1–437.4 Å3), which are significantly higher than those found in undeformed mantle xenoliths worldwide (cc. 432–435 Å3) (e.g. Dal Negro et al., 1984; Princivalle et al., 1994) and are similar to unit-cell volumes of plagioclase-spinel-bearing xenoliths (436–439 Å3) (Nimis, 1995). Therefore, these large unit-cell parameters are indicative of a relatively low equilibrium pressure in the upper mantle, close to the spinel–plagioclase transition. 6.3. Geodynamic implications Recrystallized texture and thermobarometric data on the xenoliths suggest that the mantle beneath Villány Mts cooled from an earlier

relatively high temperature (990–1010 °C) to a slightly lower temperature (910–960 °C), and went through deformation and subsequent re-equilibration. The abundant exsolution lamellae in pyroxenes (Fig.2c) shed light on the subsolidus re-equilibration event. Textural and chemical equilibrium suggest that deformation happened well before the entrainment of xenoliths, therefore reequilibration may have occurred long before the Late Cretaceous igneous activity, at slightly lower PT conditions. This scenario can be explained by emplacement and cooling of the mantle at shallower depths (sampled later by the lamprophyric melt). Diapiric uprise of mantle material into the subcontinental lithosphere may be a reasonable geodynamic model to account for such a cooling history and explain deformed mantle formation. Similar cases were described worldwide (e.g. Basu, 1975; Coisy and Nicolas, 1978; Nicolas et al., 1987; Witt and Seck, 1987; Downes et al., 1992; Embey-Isztin et al., 2001; Falus et al., 2007), involving uplift of a hotter deeper part of the mantle into a relatively colder uppermost mantle and subsequent thermal re-equilibration between these two thermally different units. Strong textural and chemical similarities of the Villány Mts samples to examples of xenolith series related to diapiric uplift (e.g. W-Eifel and Neogene basalt-hosted xenoliths from the Pannonian Basin), support a similar evolution. However the dimensions, mechanism and driving force of this uplift require further explanation. We can exclude a diapiric uprise associated with a large-scale upwelling of deep hot mantle material (i.e. plume activity) which should result in much more voluminous volcanic activity and highly depleted mantle (White and McKenzie, 1989; White, 1993) not observed in the Tisza Unit. Hence, either the presence of a small diapiric uprise or tectonic emplacement of the previously deformed lithospheric mantle section beneath the region appears more likely. The Villány mantle xenoliths strongly suggest that the mantle moved upward into a low pressure regime near to the spinel-plagioclase transition. These types of lithospheric movements are generally related to extensional tectonics. Therefore, in the following discussion, we examine the major Mesozoic geodynamic features of the Tisza unit, to which the mantle evolution, as deduced from the xenoliths, can be linked.

Table 5 Estimated equilibrium temperatures in °C (WB — Wood and Banno, 1973; W — Wells, 1977; M — Mercier, 1980; BK — Brey and Köhler, 1990).

Neoblast Porphyroclast (host) Porphyroclast (pre-exsolution)

TM-opx

TM-cpx

TBK-Ca-opx

TBK-2pyx

TW

TWB

932–995 966–972 1059–1070

855–930 898–950 982–998

854–923 885–895 994–1010

672–808 763–823 818–873

800–839 818–871 864–904

816–887 877–921 913–952

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Fig. 7. Al2O3 vs. Cr2O3 (wt.%) diagrams for clino- and orthopyroxene porphyroclasts and neoblasts. Thermometrical estimation of the recrystallization and data for comparison after Witt and Seck (1987).

The southern and southwestern part of the Tisza unit was situated at the junction of differently moving and rotating units in the Mesozoic (e.g. Plašienka, 2000; Csontos and Vörös, 2004). By the middle-late Mesozoic, due to changes in movement and rotation of the African and European plates, microcontinent movements also notably rearranged. In such cases, compression can sometimes change to extension (e.g. Wortmann et al., 2001). From the southern margin of the Tisza unit, Mesozoic volcanics are only known from restricted areas (Harangi, 1994; Huemer, 1997; Nédli and M. Tóth, 2007) and in some boreholes from the central parts of the Tisza unit (Szepesházy, 1977; Molnár, 1985). Except for the Villány Mts lamprophyres, none carry mantle xenoliths. On the other hand, Cretaceous intraplate alkaline volcanics are known from the neighbouring units: from the Alcapa unit a mantle xenolith-bearing lamprophyre dyke swarm is known (Szabó et al., 1993), and Early Cretaceous alkaline lamprophyre dykes were also described from the Western Carpathians (Hovorka and Spišiak, 1988; Dostal and Owen, 1998). As little information is available about the Mesozoic mantle evolution and geodynamics of the Tisza unit (e.g. Csontos and Bergerat, 1992; Csontos and Vörös, 2004), we take a try to reconstruct the lamprophyre and xenolith evolution using also the main Mesozoic volcanic and sedimentological events, as constraints. Considering the volcanological and sedimentological features of the middle-late Mesozoic, we suggest that the Cretaceous volcanic activity of the southwestern part of the Tisza unit can be included in a scenario of large-scale geodynamic evolution (Fig. 8), relative to the Africa and Europe plate motions and the evolution of the neighbouring Penninic–Vahic–Magura and Vardar oceans (Fig. 9). (1) The Tisza microplate separated from the European plate in the Middle Jurassic, due to the spreading of the Ligurian–Penninic– Vahic ocean (Plašienka, 2000; Marroni et al., 2002). This is also suggested by the drastically changed sedimentation pattern in the Mecsek zone (fluvial sediments were replaced by lacustrine then paralic coarser-grained siliciclastic ones, deepening to marly then siliceous and pelagic limestone sedimentation) (e.g. Haas and Péró, 2004; Császár, 2005). This separation may have been accompanied by notable lithosphere movements and fracture of the southwestern part of the Tisza unit, possibly down to mantle depths. These movements can be traced in the presence of highly different sedimentological evolution in neighbouring areas of the microplate: formation of a large half-

graben structure in the Mecsek Mts and deposition of up to 4300 m thick sediments, meanwhile coeval strata in the adjacent Villány Mts are missing or reach only a few hundreds of metres. Intense basin deepening in the Mecsek zone suggests significant crustal thinning, whereas in the Villány zone there are no traces of crustal thinning, it was more likely the area from which the sediments were transported into the subsiding Mecsek zone. (2) Rifting in the Mecsek zone may have begun in the latest Jurassic–earliest Cretaceous and evolved slowly, accompanied by volcanic activity in the Mecsek Mts (e.g. Harangi, 1994; Huemer, 1997) and in the central part of the Tisza unit (Szepesházy, 1977; Molnár, 1985). It reached its peak in the Valanginian–Hauterivian (Fig. 8, Stage 1), in connection with an incomplete outward extension of the previously opened South Penninic–Vahic and North Penninic–Magura oceanic branches (Fig. 9a) (Haas and Péró, 2004). Between the volcanic build-ups, pelagic marl was formed, while in the southern marginal zones crinoidal limestone was deposited. Contemporaneously in the Villány Mts, local subaerial emergence took place and bauxite was formed (Császár, 2002). Rifting may have been accompanied by slow mantle upwelling along the earlier generated fractures, which caused the uprise and deformation of mantle later brought to the surface as xenoliths (Fig. 8, Stage 1). Absence of voluminous volcanic activity and highly depleted mantle beneath the region – characteristic features of a mantle plume (White and McKenzie, 1989; White, 1993) – can indicate that rifting aborted before reaching significant mantle upwelling and evolution of plume-like structures. At the southern margins of the Tisza unit, the closure and collision of the Vardar ocean determined the lithospheric evolution in this period (Fig. 9a) (Csontos and Vörös, 2004). (3) After the paroxysm, rapid decrease of volcanic activity in the Mecsek Mts was followed by abortion of rifting at the beginning of the Late Cretaceous, whereas in the northern part of the Mecsek Mts, volcanic activity was active until the Albian (Császár and Turnšek, 1996). In the Villány Mts the Early Cretaceous was dominated by the accumulation of platform carbonates (Császár, 1998, 2002), indicating a tectonically relatively calm period. Platform carbonates gave way to pelagic marl by the Albian, and flysch-type sediments by the Cenomanian, suggesting continuous deepening of the area

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51

Fig. 8. Geodynamical model of the evolution of volcanism at the south-western part of the Tisza Unit. Dark grey rectangle with X: location of mantle material sampled later as xenoliths; garnet – spinel – plagioclase transition after O'Neill (1981). For detailed location of surrounding units see Fig. 9, for further explanation of lithosphere evolution see the text.

including Villany Mts and the beginning of the compressional phase (Fig. 9b, c). This is the first phase of nappe emplacement (Ianovici et al., 1976) and likely the period of mantle relaxation (lamellae formation) and perhaps metasomatism, related to dehydration/melting of the subducting slab of the Vardar ocean (Fig. 8, Stage 2). (4) Between the main nappe emplacement phases (Albian– Cenomanian and the second phase in the Late Cretaceous– Paleogene) (Ianovici et al., 1976), a relatively calm period occurred, characterized by extensional tectonics and high voluminous sedimentation. By the Late Cretaceous, convergent movements of the Adriatic promontory and the European plate caused rearrangement of intercalated plate fragments (including the Tisza unit) with horizontal displacement (e.g. Wortmann et al., 2001), therefore the extensional regime was replaced by a compressional one; the oceanic branches that previously surrounded the unit were consumed (Csontos and Vörös, 2004) (Fig. 9c, d). Also, formation of flysch-type sequences to the north of the Tisza Unit can be related to the

second (end of Late Cretaceous–Paleogene) phase of nappe emplacement. Extensional movements between these compressional phases could govern the generation and/or reactivation of deep lithospheric fractures at the southern part of Tisza, and generation of the lamprophyric melts in the Villány Mts, carrying the xenoliths to the surface (Fig. 8, Stage 3). 7. Concluding remarks Small but significant differences in chemistry and equilibrium PT conditions in the pyroxene porphyroclasts and neoblasts of the Villány Mts xenoliths indicate that the sampled mantle section went through cooling and re-equilibration. According to the spinel and pyroxene composition, the mantle experienced a low degree of melt extraction. Thermobarometry suggests re-equilibration at lower PT conditions, in the shallower zones of the upper mantle, near to the MOHO. Porphyroclasts record temperatures of approx. 1000 °C, whereas neoblasts indicate a re-equilibration at slightly lower PT conditions. Extended exsolution

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Fig. 9. Geodynamical reconstruction maps of the major units composing the Carpathian–Pannonian Region in Late Jurassic–Cretaceous after Csontos and Vörös (2004).

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features in porphyroclasts are due to long residence and cooling at ambient conditions after emplacement. This suggests that the mantle, represented by the studied xenoliths, reached low PT conditions in the subcontinental mantle beneath the Tisza unit before the Late Mesozoic. Strong textural and chemical similarities of the studied xenoliths to xenolith series deformed by diapiric uplift point to their similar evolution, related to a mantle upwelling or tectonic emplacement of hot mantle material into colder upper mantle. Based on volcanological and sedimentological constraints, we suggest that the xenolith and lamprophyre evolution is related to the Cretaceous rifting event that affected the southern part of the Tisza unit. Mantle upwelling and formation of lamprophyre melts can be related to generation or reactivation of deep fractures of the lithosphere, during lithospheric extension period between the major nappe emplacements of the region (Albian–Cenomanian and Paleocene). Acknowledgments The DCM Limestones Mine Co. is gratefully acknowledged for making the fieldwork and sampling possible and László Pongrácz for providing several xenolith hand specimens. We give many thanks to Jens Hopp and an anonymous reviewer for their constructive comments and helpful suggestions and M. Liu for editorial care of manuscript. We are grateful to F. Princivalle and the members of the Department of Earth Sciences at the University of Trieste for making possible the clinopyroxene single crystal structural analysis. Many thanks to all members of the Department of Mineralogy, Geochemistry and Petrology (University of Szeged) and Lithosphere Fluid Research Lab (Eötvös L. University, Budapest) for technical support and fruitful discussions. EMP analyses were carried out at the University of London with the financial support of the NATO Science Fellowship Foundation. The first author's studies in Villány Mts were supported by the Hungarian National Scientific Research Foundation (N. 64020 and 62468 projects). This is the 42nd publication of the Lithosphere Fluid Research Lab, in collaboration with the Department of Mineralogy, Geochemistry and Petrology (University of Szeged), Department of Regional Geology (Eötvös University) and with the Birkbeck Department of Earth and Planetary Sciences (University of London). References Balla, Z., 1987. Anti-clockwise rotation of the Mecsek (south west Hungary) in the Cretaceous: interpretation of the paleomagnetic data in the light of the geology. (in Hungarian with English abstract) Ált. Földt. Szemle 22, 55–98. Balla, Z., 1988. On the origin of the structural patterns of Hungary. Acta Geol. Hung. 31, 53–63. Basu, A.R., 1975. Hot spots, mantle plumes and a model for the origin of ultramafic xenoliths in alkalic basalts. Earth Planet. Sci. Lett. 28, 261–274. Bohrson, W.A., Clague, D.A., 1988. Origin of ultramafic xenoliths containing exsolved pyroxenes from Hualalai volcano, Hawaii. Contrib. Mineral. Petrol. 100, 139–155. Brey, G.P., Köhler, T., 1990. Geothermobarometry in four-phases lherzolites II. New thermobarometers, and practical assessment of existing thermobarometers. J. Petrol. 31, 1353–1378. Coisy, P., Nicolas, A., 1978. Regional structure and geodynamics of the upper mantle beneath the Massif Central. Nature 274, 429–432. Császár, G., 1998. The stratigraphy of the Early and Mid-Cretaceous formations of Mecsek and Villány Unit. In: Bérczi, I., Jámbor, Á. (Eds.), The stratigraphy of the geological formations of Hungary (in Hungarian). MOL Rt and MÁFI, Budapest, pp. 353–369. Császár, G., 2002. Urgon formations in Hungary with special reference to the Eastern Alps, the Western Carpathians and the Apuseni Mountains. : Acta Geol. Hun. Ser. Geol., 25. 206 pp. Császár, G. (2005) Magyarország és környezetének regionális földtana. I. Paleozoikumpaleogén – Egyetemi tankönyv, (Regional geology of Hungary and the surrounding areas. I. Palaeozoic-Palaeogene) (in Hungarian) Budapest, Eötvös Kiadó, 328 p. Császár, G., Turnšek, D., 1996. Atoll-like vestiges in the Lower Cretaceous of the Mecsek Mountains, Hungary. Cretaceous Res. 17, 419–442. Csontos, L., 1995. Tertiary tectonic evolution of the Intra-Carpathian area: a review. Acta Vulcanol. 7, 1–13. Csontos, L., Bergerat, F., 1992. Reevaluation of the Neogene brittle tectonics of the Mecsek-Villány area (SW Hungary) Annales Univ. R. Eötvös nominate. Ser. Geol. 29, 3–12.

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