Seismic crustal structure between the Transylvanian Basin and the ...

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In order to study the lithospheric structure in Romania a 450 km long WNW–ESE trending ... remains constant with about 6.0 km/s along the basement top and 7.0 km/s above the Moho. ...... 100 km offsets, but weak and hardly visible at sub-.
Tectonophysics 430 (2007) 1 – 25 www.elsevier.com/locate/tecto

Seismic crustal structure between the Transylvanian Basin and the Black Sea, Romania F. Hauser a,1 , V. Raileanu b , W. Fielitz c,⁎, C. Dinu d , M. Landes a , A. Bala b , C. Prodehl a a

Geophysical Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany National Institute for Earth Physics, P.O.Box MG-2, RO-077125 Bucuresti-Magurele, Romania c Geological Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany Faculty of Geology and Geophysics, University of Bucharest, 6, Traian Vuia Street, sector 1, RO-70139 Bucuresti, Romania b

d

Received 11 May 2005; received in revised form 17 October 2006; accepted 19 October 2006 Available online 11 December 2006

Abstract In order to study the lithospheric structure in Romania a 450 km long WNW–ESE trending seismic refraction project was carried out in August/September 2001. It runs from the Transylvanian Basin across the East Carpathian Orogen and the Vrancea seismic region to the foreland areas with the very deep Neogene Focsani Basin and the North Dobrogea Orogen on the Black Sea. A total of ten shots with charge sizes 300–1500 kg were recorded by over 700 geophones. The data quality of the experiment was variable, depending primarily on charge size but also on local geological conditions. The data interpretation indicates a multi-layered structure with variable thicknesses and velocities. The sedimentary stack comprises up to 7 layers with seismic velocities of 2.0–5.9 km/s. It reaches a maximum thickness of about 22 km within the Focsani Basin area. The sedimentary succession is composed of (1) the Carpathian nappe pile, (2) the postcollisional Neogene Transylvanian Basin, which covers the local Late Cretaceous to Paleogene Tarnava Basin, (3) the Neogene Focsani Basin in the foredeep area, which covers autochthonous Mesozoic and Palaeozoic sedimentary rocks as well as a probably PermoTriassic graben structure of the Moesian Platform, and (4) the Palaeozoic and Mesozoic rocks of the North Dobrogea Orogen. The underlying crystalline crust shows considerable thickness variations in total as well as in its individual subdivisions, which correlate well with the Tisza-Dacia, Moesian and North Dobrogea crustal blocks. The lateral velocity structure of these blocks along the seismic line remains constant with about 6.0 km/s along the basement top and 7.0 km/s above the Moho. The Tisza-Dacia block is about 33 to 37 km thick and shows low velocity zones in its uppermost 15 km, which are presumably due to basement thrusts imbricated with sedimentary successions related to the Carpathian Orogen. The crystalline crust of Moesia does not exceed 25 km and is covered by up to 22 km of sedimentary rocks. The North Dobrogea crust reaches a thickness of about 44 km and is probably composed of thick Eastern European crust overthrusted by a thin 1–2 km thick wedge of the North Dobrogea Orogen. © 2006 Elsevier B.V. All rights reserved. Keywords: Seismic refraction; Crustal velocity structure; Vrancea zone; Eastern Carpathians; Moesian Platform; Transylvanian Basin

⁎ Corresponding author. Tel.: +49 721 6082139; fax: +49 721 6082138. E-mail address: [email protected] (W. Fielitz). 1 Now at Geophysics Section, Dublin Institute for Advanced Studies, Ireland. 0040-1951/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2006.10.005

1. Introduction This study focuses on a crustal transect in Romania with a complex geological history. It crosses from W to E the Transylvanian Basin, the Carpathian Orogen and

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the Carpathian Foredeep with the exceptionally deep Focsani Basin, as well as the stable Moesian Platform and the North Dobrogea Orogen in the foreland (Fig. 1). Several plate or microplate boundaries (Tisza-Dacia in the W, Moesia in the centre and North Dobrogea in the E) are intersected, which are largely covered by the nappes of the Carpathian Orogen or the Neogene sedimentary basins. Whereas these plates were consolidated during complex pre-Mesozoic deformations, the Carpathian Orogen and the younger sedimentary basins are the result of plate convergence during Mesozoic and Neogene times. The youngest part of the orogen and basins in Romania show still ongoing neotectonic activity. This is manifested by near-surface crustal deformation resulting in a very pronounced geomorphology (Fielitz and Seghedi, 2005; Necea et al., 2005), subsidence in the Focsani foreland basin (Tarapoanca et al., 2003; Matenco et al., 2003) and a strong seismicity at intermediate depths, which is concentrated in the Vrancea area of the southeastern Carpathian bend (Fig. 1; Fuchs et al., 1979; Oncescu et al., 1998). While the shallow seismic activity scatters widely and has moderate magnitudes (Mw b 5.6), the epicentral region of the intermediate depth seismicity is confined to an area of only about 40 × 80 km2 (Oncescu et al., 1998). The epicenters of these earthquakes lie between 60 and 180 km depth within an almost vertically elongated narrow zone and frequently have large magnitudes (Mw = 6.9–7.4). Several major earthquakes occurred there in the last century (1940, 1977, 1986 and 1990) causing many fatalities and an enormous economical damage. In the 1977 earthquake alone, 1570 people died and more than 11,300 were injured. In order to study this seismically high-risk area a joint German–Romanian research program was initiated by the Collaborative Research Center 461 (CRC 461) “Strong Earthquakes — a Challenge for Geosciences and Civil Engineering” at the University of Karlsruhe (Germany) and the Romanian Group for Vrancea Strong Earthquakes (RGVE) at the Romanian Academy in Bucharest (Wenzel, 1997; Wenzel et al., 1998a). The joint geoscientific and civil engineering research activities of this project seek to better understand the tectonic processes responsible for the strong intermediate-depth seismicity, and to reduce the risk by applying appropriate techniques from civil engineering (e.g. Wenzel et al., 1998a). Two major active-source seismic refraction experiments were carried out in 1999 and 2001 as a contribution to this research program (Hauser et al., 2001, 2002). They were designed to study the crustal and

uppermost mantle structure to a depth of about 70 km underneath the Vrancea epicentral region and were jointly performed by the Geophysical and Geological Institutes of the University of Karlsruhe (Germany), the GeoForschungsZentrum in Potsdam (Germany), the National Institute for Earth Physics in Bucharest (Romania) and the University of Bucharest (Romania). The obtained crustal velocity models provide additional important a priori information for passive tomographic studies and help to calibrate the relative velocity variations which are obtained by teleseismic tomography (Wenzel et al., 1998b; Martin et al., 2005, 2006). The 1999 seismic refraction experiment was published by Hauser et al. (2001). In this paper we present the results of the VRANCEA2001 seismic refraction project, which covers the 450 km long segment from the town of Aiud in the Transylvanian Basin to the town of Tulcea near the Black Sea (Fig. 1). 2. Geological and tectonic setting The Eastern Carpathians are part of the Alpine– Carpathian orogenic belt which resulted from the convergence and collision of several microplates with the Eurasian plate during the closure of the Tethys Ocean (Sandulescu, 1984; Stampfli et al., 1998; Neugebauer et al., 2001). In this context several tectonic units have been accreted in the Carpathian area (Fig. 1). The External (Flysch) Carpathians or Moldavides in the east originated along the European (Moesian and East European) margin (Sandulescu, 1988; see Fig. 1 for location). Their convergence and the outline of the present day geometry occurred mainly during the Miocene (20–11 Ma) and climaxed in the early Late Miocene (11–12 Ma, Sarmatian; Sandulescu, 1988; Matenco and Bertotti, 2000; Matenco et al., 2003). However, the Internal Carpathians further west (i.e. Dacides and Transylvanides of Sandulescu, 1988; see Fig. 1), which are part of the Tisza-Dacia microplate (Csontos, 1995), converged already in mid-Cretaceous time (Aptian/Albian), were then partially covered by smaller basins with mid-Cretaceous (Aptian/Albian) to Palaeogene sedimentary strata and were finally transported to their present position during the Sarmatian accretion of the Moldavides against Europe (Sandulescu, 1988; Royden, 1988; Matenco and Bertotti, 2000). Their sediments accumulated on a thinned continental (Sandulescu, 1988) or an oceanic domain (Csontos, 1995; Stampfli et al., 1998; Neugebauer et al., 2001) of Jurassic to Early Cretaceous age. The geodynamic setting of the Sarmatian deformation is inadequately understood, but it appears that

F. Hauser et al. / Tectonophysics 430 (2007) 1–25 Fig. 1. Geological overview of the Eastern Carpathian bend area and its foreland with the main crustal units, nappe structures, faults and basins. The location of the VRANCEA'99 (N–S, small E–W transverse) and VRANCEA2001 NE–SW seismic refraction profiles are shown with their shot points. Compiled from various sources given in the text.

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subduction with slab break-off and lithospheric delamination during closure of the eastern prolongation of the Penninic Ocean played a major role (Radulescu and Sandulescu, 1973; Constantinescu et al., 1973; Airinei, 1977; Fuchs et al., 1979; Constantinescu and Enescu, 1984; Oncescu, 1984; Csontos, 1995; Girbacea and Frisch, 1998; Mason et al., 1998; Seghedi et al., 1998). Recently, Sperner et al. (2001, 2002) suggested a model of Miocene subduction of oceanic lithosphere beneath the Carpathian arc and subsequent “soft” continental collision, which transported denser lithospheric material into the mantle. Cloetingh et al. (2004) proposed postcollisional processes after subduction to explain the actual geodynamic scenario, which are controlled by the thermo-mechanical heterogeneities within the underthrusted lithosphere and lateral variations in the interplay between the lithosphere and surface processes. This subduction and collision connected the Tisza-Dacia microplate to the Moesian and East-European plates and

a proposed steep Miocene suture zone, which is concealed by the overthrusted Eastern Carpathian flysch nappes, would separate both lithospheric domains from each other (Sandulescu, 1988; Girbacea and Frisch, 1998; Sperner et al., 2001). The External Eastern Carpathian fold- and thrust-belt consists of a complex pile of nappes (Fig. 1; Sandulescu, 1984, 1988; Ellouz et al., 1994; Badescu, 1998; Zweigel et al., 1998; Matenco and Bertotti, 2000). They are made-up of Cretaceous marine basinal sediments and of Paleogene to Neogene flysch and Neogene molasse deposits. The outer nappes contain also Neogene evaporitic formations with salt and/or gypsum. The Internal East Carpathians are dominated by crystalline rocks and a cover of Late Palaeozoic–Cretaceous mostly marine sediments and Early Cretaceous flysch deposits. The Moldavidian nappe pile has an estimated thickness of 8–10 km (Stefanescu and Working Group, 1985; Ellouz et al., 1994; Morley, 1996; Matenco and

Fig. 2. Topographic map showing the VRANCEA2001 seismic line (thick line), as well as older parallel or transecting refraction and reflection lines (thin and dashed lines). For detailed information and references see text.

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Bertotti, 2000). This thickness is constrained by surface geology, seismic reflection and borehole data, mainly from oil exploration, but for the inner nappes thicknesses are only estimated through balanced crosssections. Deeper structures are fairly unknown and even the International Geotraverse XI (Figs. 2 and 3), which crossed the Carpathians between Focsani and Targu Secuiesc parallel to the VRANCEA2001 line, gave only some vague indications (Radulescu et al., 1976). Magnetotelluric data from Stanica and Stanica (1998) indicate a depth of 8 km for the base of the Moldavidian nappes and a depth of about 14–16 km for the basement in the same area. Below the more internal parts of the East Carpathians the magnetotelluric data show the crystalline basement at 16–18 km and north of the town of Focsani at approximately 10 km depth. The crustal thickness reaches 50 km in both areas. The Carpathian foreland is underlain by Precambrian crust with a relatively undeformed Palaeozoic–Mesozoic autochthonous platform cover and by deformed rocks of the Triassic–Jurassic North Dobrogea Orogen (Fig. 1; Sandulescu, 1984; Visarion et al., 1988; Tari et al., 1997; Seghedi, 1998; Matenco et al., 2003). The crystalline basement of the Moesian Platform is made-up of metamorphic and intrusive magmatic rocks, which sometimes have a weak acoustic contrast to their older sedimentary cover (Raileanu et al., 1994). Near the orogenic front of the Carpathians the platform sediments are partly covered by the foredeep sediments, which dip slightly towards the central part of the foredeep (e.g. Tarapoanca et al., 2003).

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The Moesian Platform to the south and the Scythian Platform to the north are characterised by distinct magnetic anomalies, which result from petrological differences in the crystalline basement, and by lithological differences of the detritic and carbonaceous platform cover (Tari et al., 1997; Seghedi, 1998 and references therein; Raileanu et al., 2005). The major platform structures in the foreland are Late Permian/Early Triassic rifts (Tari et al., 1997; Seghedi, 1998; Landes et al., 2004; Raileanu et al., 2005; Panea et al., 2005; Bocin et al., 2005). The deformed rocks of the North Dobrogea Orogen include a complex polydeformed Variscan basement and a Permian–Cretaceous sedimentary and volcanic cover (Seghedi, 1998). The whole complex was overthrusted NNE-ward onto the Scythian Platform between the Late Triassic and the Late Jurassic. The Scythian, Moesian and North Dobrogea crustal blocks are thought to belong to the southeastern prolongation of the Trans-European suture zone (TESZ; e.g. Pharao, 1999; Debacker et al., 2005). They are separated by the Trotus (TF) and Peceneaga–Camena (PCF) faults (Fig. 1). On the Moesian Platform the Capidava–Ovidiu (COF) and Intramoesian (IMF) faults separate basements of different composition. The COF separates a greenschist basement to the north from a higher-grade metamorphic basement to the south (Seghedi, 1998 and references therein). The PCF and the COF are outcropping in the Dobrogea area near the Black Sea, while sediments and nappes cover the supposed NW-prolongation of the faults as well as the IMF itself (Fig. 1; Visarion et al., 1988;

Fig. 3. International Geotraverse GT XI crustal seismic refraction line (Radulescu et al., 1976; Cornea et al., 1981; slightly modified; 1 — sediments, 2 — grantic layer, 3 — basaltic layer, 4 — boundary between sediments and basement, 5 — Conrad discontinuity, 6 — Moho, 7 — Eastern Carpathian nappes, 8 — volcanic rocks, 9 — faults, 10 — crustal and subcrustal earthquake foci). For location see Fig. 2.

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Polonic, 1996; Ellouz et al., 1994; Seghedi, 1998). In the autochthonous and overthrusted areas of the Moesian Platform the seismic refraction line is expected to cross two of these major crustal faults, the Capidava–Ovidiu Fault (COF) and the Peceneaga–Camena Fault (PC). Recent field studies show, that the Trotus/Peceneaga– Fault system is a particularly active structure with a pronounced tectonic geomorphology and local offsets of 200 m, which were achieved during the Quaternary (Matenco et al., 2003; Tarapoanca et al., 2003). Based mainly on reflection seismic data, these crustal faults are thought to extend down to the Moho discontinuity and possibly even deeper (e.g. Visarion et al., 1988; Radulescu and Diaconescu, 1998). Post-collisional (i.e. post-Sarmatian) sediments overlie structures of the Carpathian Orogen and its foreland but are also deformed by younger structures. The most important of these structures is the Focsani Basin (Fig. 1). The sediments of the foredeep and the Focsani depression reach up to 8 km for the last 14 Ma (Matenco et al., 2003; Tarapoanca et al., 2003). Late Pliocene to Early Quaternary contractional deformation in the foreland reaches as far as the Pericarpathian Front (Hippolyte and Sandulescu, 1996; Hippolyte et al., 1999; Matenco and Bertotti, 2000; Fig. 1). Along the western flank of the Focsani Basin an eastward dipping monocline developed

from the Late Pliocene to the Holocene (Tarapoanca et al., 2003; Cloetingh et al., 2003; Necea et al., 2005). Apatite fission-track data show that strong uplift and exhumation occurred in the Eastern Carpathians during the Pliocene (2–7 Ma; Sanders et al., 1999). Ongoing vertical crustal movements are revealed by geodetic measurements (Popescu and Dragoescu, 1987; Radulescu et al., 1998; Zugravescu et al., 1998). The inner part of the Carpathians, which is partially covered by sediments of the Tertiary Transylvanian Basin (Fig. 1; Paleogene to Late Miocene in the central and northern part and mainly Middle to Late Miocene in the southern part) was also affected by some Neogene to Quaternary compressional deformation (Ciulavu et al., 2000). However, along its eastern and southeastern margin the Transylvanian Basin is made-up of the postcollisional Late Miocene to Quaternary calc-alkaline volcanic Calimani–Gurghiu–Harghita (CGH) mountains (Szakács and Seghedi, 1995, 1996; Seghedi et al., 1998; Mason et al., 1998) and the Latest Pliocene to Quaternary alkalic–basaltic volcanism in the adjacent Persani mountains (Seghedi and Szakács, 1994; Downes et al., 1995; Panaiotu et al., 2004), which are structurally connected to the sinistral–transtensional Gheorgheni–Ciuc–Brasov graben system (Fielitz and Seghedi, 2005 and references therein; Fig. 1).

Fig. 4. Pre-2001 geological section along the main WNW–ESE VRANCEA2001 seismic-refraction line. The section is mostly transverse to the trend of the main geological structures except for the basement structures of the foreland, which are highly oblique to the seismic line. For location see Fig. 1. Compiled from various sources. The near-surface geology is based on mapped exposures as published in geologic cross-sections by Stefanescu and Working Group (1985) and Matenco and Bertotti (2000). The extensional structures of the inner Eastern Carpathians (Brasov basin system) are as interpreted by the authors. The Moho depth is from Radulescu et al. (1976). The intracrustal thrust, alternative base of the Moho (dashed line) and the low velocity zone (L.V.Z.) are from Hauser et al. (2001). The proposed steep Miocene suture zone separating the lithospheric domains of Tisza-Dacia and Moesia is from Sandulescu (1988), Girbacea and Frisch (1998) and Sperner et al. (2001). Neogene to Quaternary andesites are tentatively projected down-strike the East Carpathian Calimani–Gurghiu–Harghita volcanic chain into the cross-section to indicate the possibility of occurrence of such rocks in the crust. The Neogene filling of the Transylvanian Basin is based on Ciulavu et al. (2000).

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Using published as well as unpublished geological data, a crustal cross-section for the VRANCEA2001 refraction line was compiled (Fig. 4). The section is mostly transverse to the trend of the main geological structures except for the basement structures of the foreland, which are highly oblique to the seismic line. The near-surface geology is based on mapped exposures as published in geological cross-sections by Stefanescu and Working Group (1985) and by Matenco and Bertotti (2000). The extensional structures of the inner Eastern Carpathians (Brasov basin system) are as interpreted by the authors. The Moho depth is from Radulescu et al. (1976). The intracrustal thrust, alternative base of the Moho (dashed line) and the low velocity zone (L.V.Z.) are from Hauser et al. (2001). The proposed steep Miocene suture zone separating the lithospheric domains of Tisza-Dacia and Moesia is from Sandulescu (1988), Girbacea and Frisch (1998) and Sperner et al. (2001). Neogene to Quaternary andesites are tentatively projected down-strike the East Carpathian Calimani–Gurghiu–Harghita volcanic chain into the cross-section to indicate the possibility of occurrence of such rocks in the crust. The Neogene filling of the Transylvanian Basin is based on Ciulavu et al. (2000). This cross-section, which is simplified and adapted to the scale and resolution of the seismic line, forms the base for a first geological interpretation of the velocitydepth model obtained from the seismic experiment as presented in this paper. It can later be integrated into a complex larger scale geodynamic model using additional geological and geophysical (e.g. mantle tomography) data. This is however not the focus of this paper. 3. Earlier seismic investigations Parts of the study area have been investigated by different geophysical methods during the last decades of the past century. Besides gravity, magnetic, magnetotelluric and heat flow measurements, seismic reflection and refraction data played an important role. Here we summarize the results, which are relevant for the VRANCEA2001 seismic refraction project. On the Moesian Platform and in the Focsani and Transylvanian basin areas the VRANCEA2001 seismic line is either intersected by or sub-parallel to nearby seismic reflection profiles recorded for oil and gas exploration. These seismic lines mainly mapped the structure of the Neogene cover and to a lesser extent the pre-Neogene sedimentary succession and the crystalline basement. A representative deep seismic reflection line across the Moesian Platform extends from Ramnicu Sarat some 40 km towards the east (RmS in Fig. 2). It shows a rapid thickening of the Neogene sediments from about 8 km at

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the eastern end of the line to 12 km near Ramnicu Sarat (Raileanu and Diaconescu, 1998). The interpretation of deeper-crustal features defines the crystalline basement at 16 km, an intra-crustal boundary at 32 km and the Moho at about 42 km depth (Raileanu and Diaconescu, 1998). Within the Transylvanian Basin two deep reflection lines cross the basin from NNW to SSE and from W to E, respectively and both lines end near the VRANCEA2001 profile (Tr1 and Tr2 in Fig. 2; Raileanu and Diaconescu, 1998; Raileanu, 1998). The sedimentary sequence is composed of a pre-Cretaceous succession, a post-tectonic Late Cretaceous to Palaeogene cover and a Neogene succession with a total thickness of 9–10 km. The crystalline crust is ‘transparent’ and no significant reflections are observed except for some short, dipping signals with weak coherency and most probably of diffractive nature. Sometimes, and mainly within the central part of the basin, the basement is marked by reflections. Structures are observed between 9 and 10 s two way time (TWT) or approximately 27–30 km depth and 12–14 s TWT or approximately 36–42 km depth. These reflections are interpreted as a transition zone from crust to mantle, between approximately 27 km and 42 km (Raileanu and Diaconescu, 1998; Raileanu, 1998). Based on seismic reflection and well data, Polonic (1998) produced a crystalline basement map of Romania. This map predicts a dramatic deepening of the basement along the VRANCEA2001 seismic line, from 1–3 km in the Dobrogea area to 15 km in the Focsani Basin, 8–10 km in the Carpathian Orogen area and 5–8 km in the Transylvanian Basin. Previous crustal refraction lines recorded within the North Dobrogea (Pompilian et al., 1993), the Focsani Basin (Enescu et al., 1972), the Moesian Platform and the Focsani and Transylvanian basins (Radulescu et al., 1976; Cornea et al., 1981) give further information on the deep crustal structures. Within the North Dobrogea a 6 km thick sedimentary cover is observed and the Moho reaches a depth of about 42–43 km (Pompilian et al., 1993). The crustal refraction line GT XI 1 (Fig. 2), recorded in the 1970's from Focsani southwards over a length of 60– 70 km, points out two sedimentary boundaries at 5 km and 10 km depth, respectively, a basement at approximately 17–18 km (K0 boundary), and an intra-crustal boundary (K1) at 26 km (Enescu et al., 1972). Another crustal seismic refraction line GT II (Fig. 2) runs from the territory of Ukraine and the Moldavian Republic through Galati with a SSW direction to Calarasi near the Bulgarian border. It crosses the Peceneaga–Camena Fault south of Braila (Fig. 1), where a large offset at all crustal levels is recorded. The

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interface K1 deepens from about 15 km south of the fault to about 21 km to the north of it and the Moho deepens from 40 km to 48 km (Radulescu et al., 1976; Cornea et al., 1981). The most important crustal seismic refraction line GT XI (Figs. 2 and 3) was recorded in the first half of the 1970's (Radulescu et al., 1976). It runs nearly parallel to the VRANCEA2001 profile at a distance of about 20 km north of Braila and of about 50 km north of Medias. It shows the following main features. In the Galati area the crust has a total thickness of 42–43 km with an upper crustal thickness of 20–22 km. In the Focsani Basin the sedimentary cover reaches a depth of 20 km. A crustal fault within the centre of the basin separates, at basement level, an uplifted block to the east from a subsided one to the west. An intra-crustal boundary (K1) that delimits the upper and lower crust is located at 32 km depth to the west and approximately 27 km to the east of the fault. The Moho is located at 47 km to the west and at approximately 42 km to the east of the fault. Further to the west the crustal boundaries are shallower. In the center of the Eastern Carpathians the K0-boundary lies at 7 km, K1 at 20 km and the Moho at 40 km depth. In the Transylvanian Basin K0 is at 3 km, K1 at 12– 18 km and the Moho at 30 km (Radulescu et al., 1976). Based on data collected before 1999 Radulescu (1988) and Enescu et al. (1992) produced contour-maps of the two major boundaries (Conrad and Moho discontinuities) in the crust. Along the VRANCEA2001

line depth values of 22 km and 45 km are predicted for the Dobrogea area, 26–28 km and 42 km for the Focsani Basin, 24 km and 52 km for the Carpathian Orogen, and 14 km and 34 km for the Transylvanian Basin. A crosssection parallel to the VRANCEA2001 line with up to scale projections of the Geotraverse XI, RmS and Tr1 lines is shown in Fig. 5 of Knapp et al. (2005). For the Geotraverse XI see also Fig. 3. The most recent refraction seismic line is the VRANCEA'99 line, the first seismic line of the present joint research project within the German–Romanian program mentioned above. This VRANCEA'99 line intersects the VRANCEA2001 profile in the East Carpathian mountain area at the centre of the seismogenic “Vrancea area” (Figs. 1 and 4). Its interpretation shows a multi-layered structure (Hauser et al., 2001) and at the intersecting point it depicts 3 sedimentary layers, an upper and a lower crustal layer. The base of the sedimentary succession is at 11 km, the intra-crustal boundary at 29–30 km and the Moho at about 41 km depth. The P-wave velocities for the sedimentary successions are 3.90–5.70 km/s, for the crystalline crust 5.90–7.00 km/s and for the top of the upper mantle 7.90 km/s. 4. The VRANCEA2001 seismic experiment In August/September 2001 a large seismic experiment, called VRANCEA2001, was performed. After the

Fig. 5. Trace normalised P-wave record section from shot point O with 6 km/s reduced time. The calculated travel times from the Vp model in Fig. 11 are superimposed on the data. Travel times are labelled as follows: P1–P5 = first arrival phases refracted within the sedimentary cover; Pg1 and Pg2 = diving waves through the upper and middle crust, respectively; Pi1P = reflected waves from the top of the middle crust; Pi2P = reflected waves from the top of the lower crust; P6P and P8P = reflected waves from the base of the low velocity layers L6 and L8; PmP = reflected wave from the crustmantle boundary (Moho); Pn = diving wave through the upper mantle; Diff. = supposedly diffracted waves.

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Fig. 6. Trace normalised P-wave record section from shot point S. For further explanations see Fig. 5.

completion of the first north–south running VRANCEA'99 line, briefly described above, it was the second experiment to target the deep structure of the seismogenic Vrancea area with the aim of testing new and existing geodynamic models. The experiment recorded refraction/wide-angle reflection seismic data along a 700 km line from east of Tulcea through the Vrancea zone to Karcag in Hungary (see Figs. 1 and 2). Ten large drill hole shots (300– 1500 kg charge) were fired in Romania between Aiud at the western margin of the Transylvanian Basin and the Black Sea which resulted in an average shot point spacing of 40 km. To the west an additional shot

(500 kg) was fired in Hungary. These shots generated 11 seismogram sections recorded by almost 800 geophones. The spacing of the geophones was variable. It was around 1 km from the eastern end to Aiud (ca. 450 km length), 6 km from Aiud to Oradea (Romanian– Hungarian border) and about 2 km on the Hungarian territory. Between shot points T and U the geophones were deployed at a spacing of 100 m. In total, 790 recording instruments were available. The 640 onecomponent geophones (TEXAN type) were mostly deployed in the open field outside of localities, while for security reasons the 150 three-components geophones (REFTEK and PDAS type) were deployed in guarded

Fig. 7. Trace normalised P-wave record section from shot point T. For further explanations see Fig. 5.

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Fig. 8. Trace normalised P-wave record section from shot point W. For further explanations see Fig. 5.

properties within towns and villages. The experiment was jointly performed by research institutes and universities from Germany, Romania, the Netherlands, Hungary and the United States. They included the University of Karlsruhe and the GeoForschungszentrum Potsdam, Germany, the Free University of Amsterdam, Netherlands, the National Institute for Earth Physics and the University of Bucharest, Romania, the Eötvös Loránd Geophysical Institute in Budapest, Hungary and the Universities of Texas at El Paso and South Carolina, USA. Field recording instruments were provided by University of Texas at El Paso and the IRIS/PASSCAL instrument pool, USA (TEXAN 1component stations), and the GeoForschungsZentrum, Germany (3-component REFTEK and PDAS stations). In this paper we only deal with the 450 km long main line extending from the Transylvanian Basin to the Black Sea.

a point along the refraction model with respect to shot point Z. Due the complex geological structures of the Carpathian Orogen, the data quality of the experiment is quite variable (Tables 1 and 2) but seems to depend primarily on the charge size. It was also influenced by local geological conditions around the shot points, the receivers, as well as by propagation conditions between the source and the receivers. Nevertheless, the shots Table 1 Summary of data quality as a function of offset for each shot point No.

Shot point

Branch

1

Z

2

Y

3

X

4

W

5

U

6

T

7

S

8

R

9

P

10

O

5. Seismic sections The seismic record sections were compiled and plotted using the SeismicHandler program package of Stammler (1994). All seismic sections presented in this paper are displayed with a reduction velocity of 6 km/s (Figs. 5–8). Seismograms are normalized with respect to the maximum amplitude per trace. A general bandpass filter between 3 and 12 Hz was applied to the data to improve the signal-to-noise ratio. For this paper the term “travel time” refers to the reduced travel time of the seismic section. The term “offset” is used for the distance between the source and the receivers, while the term “distance” is used to locate

Good

Acceptable

Poor

Distance (km)

Distance (km)

Distance (km)

West East West East West East West East West East

– 0–140 0–70 0–95 0–40 0–30 0–100 0–90 0–50 0–40



West East West East West East West East West East

0–95 0–40 0–105 0–50 0–105 0–120 0–70 0–70 0–95 0–30

– – 105–215 50–130 – – – – 95–220 –

– 140–220 – 95–175 – 30–130 100–145 90–130 50–150 40–100 very poor; 170–220 poor PmP 95–135 40–85; 120–200 215–290 130–155 105–205 – 70–235 – 220–285

– – – – – – – –

F. Hauser et al. / Tectonophysics 430 (2007) 1–25

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Table 2 Summary of the P-wave phase correlation and their relative quality

O–e O–w P–e P–w R–e R–w S–e S–w T–e T–w U–e U–w W–e W–w X–e X–w Y–e Y–w Z–e

P1

P2

P3

P4

P5

P6P

Pg1

Pi1P

P8P

Pg2

Pi2P

PmP

Pn

– – – – 3 3 3 3 – – – – – – – – – – –

– – – – – – 3 3 3 3 3 3 3 3 3 3 3 3 3

– – – – – – 3 3 – – – – – – – – – – –

– – – – – – 3 3 3 3 3–2 3 3 3–2 3 3 3 3 3

3 3 3 3 3 3 2–1 2 – 2 2 2 2–3 3–2 2 2 3 2–3 3

– – – – – – – 3–2 2 2–3 3 2–3 2–3 3 2 3 2 – –

3 3 3 3–2 3 – 2 – – 2–3 – 3–2 3 3 – 2 3–2 2 3

2–3 2–3 2–3 2 2 2–1 2 2–3 2 2–1 1 2 2–3 3–2 1 2 2 2 2–3

– – – – – – – 3–2 1 2–1 1 1–2 1–2 2 1 2 2 2 2

– 2 – – – – – 2 – – – – – 2 – – 1–2 – 3

1 2 2 2 2 1–2 2 2 1 2–1 – 2–1 2 2–1 – 2–1 2 1 2

1 2–3 1 1 1 – 2 3 1 2–3 1 2–1 2 1 1 1 1 1 2–1

– 1 – – – – – 3 1 – – – – – – – – – 2

Correlation ranks: 3—easy correlation, 2—difficult correlation, 1—very difficult correlation. Combination of two numbers means intermediate assessments between two ranks. (–): phase is not present or the high noise did not allow any correlation.

from the eastern half of the line were more effective then the shots from the western half. Especially shot point S (300 kg charge) generated coherent signals up to the western end of the line, although some secondary arrivals are suspected to be local diffractions. Some seismic sections show segments of good or acceptable data (i.e. clean and easy to correlate traces), followed by a segment of noisy traces (Figs. 5–9). This is probably due to the local noise levels during the recordings. Higher quality data seem to be located below the Carpathians

and lower quality data below the Transylvania Basin and partially the Focsani Basin and North Dobrogea Orogen. Table 1 summarizes the data quality as a function of the offset. Not all phases could be identified on all record sections, sometimes because the phase is not present, but sometimes also due to low signal-to-noise ratio (Table 2). We picked only those phases, which are coherent over several traces. Up to five first-arrival refracted phases with apparent velocities b 6 km/s could be identified.

Fig. 9. Trace normalised P-wave record section from shot point Z. For further explanations see Fig. 5.

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F. Hauser et al. / Tectonophysics 430 (2007) 1–25

find the model that fits the seismic data best we used several major steps in the modelling procedure:

They have been labelled P1 to P5 and are attributed to waves travelling through different sedimentary layers or intercalations of sedimentary, metamorphic, and volcanic layers (Figs. 5–9). Most of the data show a clear first arrival Pg1 phase (a diving wave through the upper crust) from 20 km to over 100 km offset (Figs. 5 and 7–9). This phase is characterised by strong undulations and sometimes very small amplitudes, making picking generally difficult beyond 80 km offset. Especially in Transylvania (shot points Z–W) this phase shows very high and variable apparent velocities (6.0–6.2 km/s) and appears to be split into two phases Pg1 and Pg2 (e.g. Fig. 9), separating a low velocity layer which generated the reflected phase P6P on its base (Figs. 6–8). A Pg2 phase is also observed in the eastern part, but only for shot points O and S (Figs. 5 and 6). A reflected Pi1P phase from the top of a low velocity layer in the western half and from an intracrustal boundary in the eastern half of the seismic line is correlated almost along the whole line (Figs. 5–9). Another reflected phase P8P was generated at the base of a second low velocity layer within the western half of the seismic line (Figs. 6–9). Within the middle crust we could identify secondary Pi2P arrivals (the reflection from the top of the lower crust (Figs. 5–9), between 30 and 140 km offset. The phase is prominent and laterally coherent beyond 100 km offsets, but weak and hardly visible at subcritical offsets. Critical offsets for Pi2P ranges from 120–130 km for shots that recorded data at both ends of the line to 90–110 km within the central part. The reflection from the crust-mantle boundary (PmP) is well observed on several record sections, especially from the larger shots at both ends of the profile and from shot point S (Figs. 5–9). The critical offset for this phase varies from ca. 90 km in the west to 120 km in the east and an apparent velocity of 6.8–6.9 km/s is observed at larger offsets. A Pn diving wave through the upper mantle can only be observed for a few shot points (Figs. 5, 6 and 9; Table 2). If present, Pn can be seen between 140 and 230 km offset. The apparent velocities lie between 7.9 and 8.1 km/s. Table 2 shows the correlated phases and their relative quality on a scale from 1 (low) to 3 (high). In some cases certain strong and coherent signals are displayed on seismograms which might be interpreted as diffractions (Fig. 5, 6, 8, 9). Sometimes those diffracted waves overlap with useful signals.

An initial 2D velocity model was obtained by Hauser et al. (2002) by picking first arrival times only and inverting them using the non-linear tomographic technique of Hole (1992). The most prominent result of this inversion is seen in the center of the model extending to a depth of some 20 km. It is associated with the Carpathian Orogen and the Focsani Basin (see Fig. 3 in Hauser et al., 2002). Since reflections from the crust-mantle boundary and from an intra-crustal layer are visible on most sections, this information was used to construct initial 1D models for individual shot points. Both informations were then used to construct a 2D starting model for ray-tracing, making due allowance for the offsets in the different phases. Next, the data were interpreted by 2D forward ray-tracing. For this we used the RAYINVR program, which is based on the method of Zelt and Smith (1992) and which allows the inclusion of reflected phases, therefore using more of the information present in the seismic wave field. In the ray-tracing procedure, travel times are successively re-calculated for a sequence of models which are constructed by the interpreter. In the initial modelling, the uppermost velocity structure is determined using the appropriate seismic phases. This is followed by stepwise modelling of phases from deeper layers, usually in a sequence in which the velocity is determined from the refracted phase, after which the thickness is determined by the relevant reflected phase. Usually only super-critical reflections are strong enough to be identified in the seismic record sections. This procedure is continued until a reasonable fit is achieved, as judged from the assessed uncertainty of each modelled seismic phase.

6. Interpretation techniques and velocity model

6.1. Uncertainties

For the interpretation of the P-wave seismic data only vertical component seismograms were used. In order to

A total of 13 seismic phases were correlated and picked (see Table 2), which already indicates the

(1) Travel times and associated errors were picked for each of the seismic phases described above. The integrity of the picks and the consistency of the phase identification were checked by comparing reciprocal travel times where possible. (2) One-dimensional (1D) and two-dimensional (2D) velocity-depth functions were calculated using tomographic inversion methods as well as raytracing techniques.

F. Hauser et al. / Tectonophysics 430 (2007) 1–25

complexity of the area. The pick uncertainties of the seismic phase arrivals were estimated to ± 0.1 s for all phases, except PmP and Pn which were estimated as

13

± 0.15 s. At the end of the forward modelling, when a reasonable agreement between observed and calculated travel times was reached, an inversion of all 5142 picked

Fig. 10. (a) Ray coverage through the final model connecting all shot and receiver pairs for all phases (upper figure) and only for Pi2P, PmP and Pn (lower figure). Small black dots indicate the shot points named as in Figs. 1, 11 and 12. (b) Comparison of observed and calculated travel times for all shots and all phases. Vertical bars indicate observed data with height representing pick uncertainties: ±0.10 s for all phases except PmP and Pn with ±0.15 s. Solid lines indicate calculated travel times.

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data points was carried out. The final model gave a rootmean-square (RMS) travel time residual of ± 0.117 s and a normalized χ2 = 1.244. The largest deviations from χ2 = 1 were observed for P4, P5, Pg1 and Pn with values between 2 and 5. A ray coverage through the final model for all shot and receiver pairs is presented in Fig. 10a, and a comparison of the observed and calculated travel times for all shots and all correlated phases in Fig. 10b. From this we estimate the uncertainties for the P-wave velocity from ± 0.1 km/s for the upper and middle crust to ± 0.15–0.20 km/s for the lower crust and upper mantle. For the interfaces we estimate depth uncertainties from ±0.5–1.0 km for shallow layers of the model to ±1.0–1.5 km for deeper interfaces. Based on all above data the model interfaces were drawn with thick solid lines for well resolved regions, and with dashed lines for less well resolved regions (Fig. 11).

6.2. The 2D velocity model The final 2D velocity model derived using the methods described above is shown in Fig. 11. It has a multi-layered character and reflects the different tectonic units (Transylvanian Basin, Eastern Carpathian Orogen, Moesian Platform including the Focsani Basin and North Dobrogea Orogen), which are crossed by the seismic line. The main horizontal structures can be separated into two different groups : (1) the sedimentary cover with imbricated crystalline and volcanic rocks showing velocities b 6 km/s; (2) the crystalline crust down to the crust-mantle boundary (Moho). The sedimentary succession with imbricated crystalline and volcanic rocks along the profile consists of up to seven layers (L1–L6 and L8 in Fig. 11) with velocities ranging from 2.0 to 5.9 km/s. Within the

Fig. 11. VRANCEA2001–2D velocity-depth model. Labelled dots at the top of the model indicate the shot points from O to Z and the decimal numbers show the P-wave velocities in km/s. The VR 99 arrow marks the intersection with the VRANCEA'99 seismic line and PCF the Peceneaga– Camena Fault. L1 to L11 indicate the seismic layers used in the interpretation. Thick solid lines indicate areas which are well constrained by reflections and/or refractions. Dashed lines indicate less well constrained areas, while thin lines are extrapolations (see also Table 2).

F. Hauser et al. / Tectonophysics 430 (2007) 1–25

Eastern Carpathian Orogen we find four layers (L2, L4, L5 and L6) above the crystalline basement (L7 with 6.0 km/s) with variable thicknesses and increasing velocities from 3.60 km/s at the surface to 5.70 km/s at the base of the succession. The lowest layer (L6) is a low-velocity zone with 5.30 km/s sandwiched between higher-velocity layers. Further to the west and into the Transylvanian Basin we find three layers (L2, L4 and L5) with variable thicknesses and seismic velocities from 3.00 km/s near the top to 5.80 km/s at the sediment-basement interface. To the east and into the Focsani Basin the sediment thickness increases drastically to about 20–22 km. The seismic velocities cover a wide range and increase from 2 km/s at the surface to 5.6 km/s in layer L5. The deepest part of the Focsani Basin shows three layers (L6–L8) with velocities between 5.1 and 5.9 km/s, which are attributed to a succession of different sedimentary rocks characterised by a high-velocity layer (L7) wedged in between layers with lower velocities (L6, L8). The North Dobrogea Orogen is covered by a thin wedge of sediments, volcanics and imbricated basement (L5; 1– 3 km thick) with rather high velocities (5.00–5.90 km/s). The seismic basement (L7–L10) coincides with a depth where velocities exceed 5.9 km/s. The upper crustal velocities are very heterogeneous and seem to reflect different tectonic units. In addition, we find a low-velocity layer (L8 with 5.50–6.00 km/s) within the upper crystalline crust, which extends from the western end of the seismic line to the Focsani Basin. A distinct intra-crustal boundary separates the middle crust (L9; 6.1–6.5 km/s) from the lower crust (L10; 6.7–7.1 km/s) with varying depth from 27 km at the western end to 29 km below the Focsani Basin and 27 km at the eastern end of the seismic line. Wide-angle Moho reflections (PmP) indicate the existence of a first-order crust-mantle boundary (between L10 and L11). The Moho topography shows a thickening of the crust from 37 km in the west to 42– 46 km below the Focsani Basin and 44 km in the North Dobrogea Orogen area in the east. No pronounced crustal root below the Carpathian Orogen is recognizable. Some constraints on upper mantle seismic velocities are provided by Pn arrival times picked from several shot points. 7. Discussion and interpretation The VRANCEA2001 seismic refraction model in Figs. 11 and 12 demonstrates considerable lateral thickness and velocity variations within the sedimentary succession as well as in the deeper crust. Three different

15

crustal blocks characterised by clearly distinct geometries and velocity structures were identified (Fig. 12, from west to east): (1) the Tisza-Dacia crustal block, which underlies the Transylvanian Basin and most of the Eastern Carpathian Orogen, (2) the Moesian Platform crustal block, which underlies the Focsani Basin, and (3) the North Dobrogea Orogen crustal block. The last two tectonic units are separated by the Peceneaga–Camena Fault while the contact between the first two units is less well defined and concealed by the East Carpathian nappes. A proposed steep Miocene suture zone is, however, thought to separate both lithospheric domains (Sandulescu, 1988; Girbacea and Frisch, 1998; Sperner et al., 2001) and we therefore attribute the complete crust with its crystalline and sediment-derived nappes to the individual plates (Tisza-Dacia and Moesian crustal blocks; details in Section 7.3.). Additional constraints for the interpretation of the seismic model were provided by seismic reflection data, borehole data for which published velocity logs are available, and other geophysical data. 7.1. The sedimentary sequence with imbricated crystalline rocks The sedimentary sequence is composed of the East Carpathian flysch nappes, the Neogene infill of the Transylvanian Basin in the hinterland and the Focsani Basin in the foredeep, the autochthonous Mesozoic and Palaeozoic sedimentary rocks of the Moesian Platform and the North Dobrogea Orogen, and the deeper sedimentary basins below the Transylvanian and the Focsani basins (Fig. 12). The different outcropping geological units are reflected by the laterally variable velocity structure in the seismic model (Fig. 11). The sedimentary sequence along the line, which also comprises some imbricated crystalline rocks as specified below, was subdivided into several geological units (shown in Fig. 12 with different colours) made up of a single or several seismic layers as identified in the velocity model. The first geological unit represents the western part of seismic layer L2 and the upper part of seismic layer L4 and extends from west of shot point Z to east of shot point X (yellow unit in Fig. 12). The thickness of the upper seismic layer L2 is nearly constant, at about 1 km, and the velocities within this layer range from 3.0 to 3.3 km/s. In seismic layer L4 the upper approximately up to 2 km of this geological unit with velocities of 3.9 km/s are not well separated from the next second geological unit probably because of similar composition and/or consolidation of the sedimentary rocks. Because

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F. Hauser et al. / Tectonophysics 430 (2007) 1–25

Fig. 12. Interpreted geological cross-section (top: 4.5 × vertical exaggeration, bottom: without vertical exaggeration) from the 2D seismic model of Fig. 11 along the main VRANCEA2001 seismic refraction line between the Transylvanian Basin and the Black Sea. The upper crustal geological structures of the Tisza-Dacia and the Moesian crustal blocks are transverse to the section. The proposed out-of-sequence thrusting in the crystalline basement (labeled with number 1) and the geologic structures of the North Dobrogea crustal block in the foreland (labeled with number 2) are oblique to the seismic line. For location of the section and for location of the major geological structures compare with Fig. 1.

of their proximity to the surface and from reflection seismic lines and boreholes from gas exploration (Ciulavu, 1999; Ciulavu et al., 2000) they are however known to exist. This geological unit with in total up to 3 km of thickness represents the Neogene cover of the Transylvanian Basin (Fig. 12). The second geological unit (seismic layer L4 between west of shot point Z and shot point Y in Fig. 11) underlies geological unit 1 within the Transylvanian Basin (green unit in Fig. 12). It has an asymmetric shape

with a steeper eastern flank and its thickness increases from 0 km in the west to about 2 km in the east. The velocity within this layer increases from 3.9 km/s at the top to 4.2 km/s at the bottom. It represents the Tarnava Basin, which has a half-graben geometry and is filled with Late Cretaceous to Paleogene sediments and is underlain by Early Cretaceous sediments and Jurassic volcanics. It is known from reflection seismic lines and boreholes from gas exploration (Ciulavu, 1999; Ciulavu et al., 2000).

F. Hauser et al. / Tectonophysics 430 (2007) 1–25

The third geological unit comprises seismic layers L2 and L4 from west of shot point W to shot point T (Fig. 11; brown unit in Fig. 12) and has a thickness of 4– 6 km. Its velocities range between 3.6 and 5.1 km/s with strong lateral as well as vertical variations. This unit represents the presumed unrooted sedimentary nappe pile of the Carpathian Orogen made up mainly of Triassic to Neogene rocks (mostly Cretaceous and Tertiary flysch of the Moldavide and Outer Dacide nappes; see Figs. 1 and 4). The fourth geological unit (seismic layer L4 from shot point Y to west of shot point W and layer L5 from west of shot point Z to halfway between shot points U and T in Fig. 11) underlies geological units 1, 2 and 3 (Transylvanian and Tarnava Basins and sedimentary nappe pile of the Carpathian Orogen; uppermost part of violet unit in Fig. 12). It reaches the surface between shot points X and W, where it can be correlated to surface outcrops. The thickness of this unit ranges between 2 and 5 km with highly variable velocities between 4.0 and 5.8 km/s reaching the maximum velocity (5.4 to 5.8 km/s) at the bottom of layer L5 at a depth of about 7 km. This unit represents imbricated basement nappes of the Median Dacides (mainly Bucovinian nappes; Sandulescu, 1984) of the Internal and External Carpathian Orogen. It is composed of metamorphic and other crystalline rocks. A heterogeneous composition with participation of lower-grade metamorphic, possibly Palaeozoic rocks, and a localized thin cover of autochthonous Palaeo and Mesozoic rocks could explain the velocity variations, especially the low velocities (b 4.9 km/s) within seismic layer L4. This geological unit may, west of shot point Y, contain thin thrust sheets of ophiolithic rocks from the Transylvanide nappes as proposed in the section of Fig. 4. But velocities point to a more generally crystalline composition of the crust. There is also no indication of the proposed Miocene suture zone below the third geological unit (Eastern Carpathian flysch nappes) since velocities seem to be laterally and horizontally relatively continuous and lower than expected for mafic or ultramafic ophiolitic rocks. The fifth geological unit (seismic layer L6 in Fig. 11) is an about 2 to 4 km thick low velocity zone with velocities of 5.3 km/s (blue unit in Fig. 12). It underlies geological unit 4 (basement nappes of the Carpathian Orogen) and its western boundary near shot point Y is in western down-dip prolongation of the overthrusted basement nappes at the surface (see preview crosssection in Fig. 4, east of shot point X). Because of this relationship it can be interpreted as the autochthonous Palaeozoic and/or Mesozoic cover of the basement

17

overthrusted by the deeper, higher velocity metamorphic crystalline basement of geological unit 4 (Fig. 12; Sandulescu, 1984). The sixth geological unit (seismic layers L1–L3 east of shot point T and L4 below and east of shot point T in Fig. 11; eastern yellow unit in Fig. 12) partially underlies geological unit 3 (sedimentary nappe pile of the External Carpathians). It has its greatest thickness of about 10 km east of shot point T at the front of the Carpathian nappes and thins continuously from shot point S towards shot point R to only 1 km and disappears further east. In its lower part (seismic layer L4) it displays an asymmetric shape with a steep western flank, which clearly separates this unit from the western area, and a relatively gentle dipping eastern flank. In the upper parts (seismic layers L1 to L3) its more symmetrical shape correlates with data known from surface geology and reflection seismic lines (Tarapoanca et al., 2003). Velocities increase from 2.0 to 4.8 km/s at the base of the unit and are laterally continuous. This geological unit corresponds to the Middle Miocene to Quaternary sedimentary fill of the Focsani Basin. The seventh geological unit (seismic layer 5 in Fig. 11) begins about halfway between shot points U and T and can be followed along the base of geological unit 6 (Focsani Basin) until it reaches the surface between shot points R and P (dark-blue unit in Fig. 12), from where it stays at the surface to the eastern end of the section (easternmost light blue unit in Fig. 12). Its thickness is about 2–3 km, while slightly decreasing to about 1– 1.5 km east of shot point R. Velocities are in the range of 5.4–5.9 km/s with only minor lateral variations. The deeper part of this unit (dark blue segment) probably represents the autochthonous Mesozoic and maybe the very thin Cenozoic sedimentary cover rocks of the Moesian Platform below the Neogene Focsani Basin. The high velocities would correlate with widely occurring carbonate rocks in these layers (Tari et al., 1997). Observations of Poisson's ratio along the VRANCEA'99 seismic refraction line seem to confirm this interpretation (Raileanu et al., 2005). In the Dobrogea area, this unit is also composed of Triassic volcanic rocks and imbricated Palaeozoic sedimentary, magmatic or metamorphic rocks. The whole layer in this crustal block probably represents the NNE-ward overthrusted North Dobrogea Orogen (light blue segment in Fig. 12). The eighth geological unit (seismic layer 6 between east of shot point T and shot point S in Fig. 11; orange unit in Fig. 12) is a homogeneous low-velocity layer with a thickness of about 5 km and velocities of 5.1 km/s. Its isolated appearance with sharp lateral boundaries point to a graben-like structure, while the lower velocities

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F. Hauser et al. / Tectonophysics 430 (2007) 1–25

indicate a more clastic sedimentary succession. Since it is located below the proposed autochthonous Mesozoic cover rocks of the Moesian Platform, we propose a Permo-Triassic graben structure (Fig. 12). Similar geological structures of the Moesian Platform in this or nearby areas have already been described or proposed by Tari et al. (1997), Landes et al. (2004), Panea et al. (2005), Bocin et al. (2005) and Raileanu et al. (2005). The ninth geological unit (seismic layers L7 from halfway between shot points U and T until shot point R and seismic layer L8 between shot points W and S in Fig. 11) has a relatively constant thickness of about 6 km while its velocities range from 5.5 to 5.9 km/s (light blue unit in Fig. 12). Its western part is located below the Carpathian nappes at 14 km depth, its central part at greater depth below the proposed Permo-Triassic graben, and the continuation to the east would be covered by Mesozoic platform rocks. Because this unit underlies the above graben structure, we interpret it for the central and eastern part as the autochthonous Palaeozoic cover rocks of the Moesian Platform (Fig. 12). The western part (west of halfway between shot points U and T) could be made up of similar rocks covering the eastern margin of the Tisza-Dacia block with a slightly thinned middle and lower crust, which were later overthrusted by crystalline Carpathian nappes with higher velocities. The high velocities can, again, be correlated in the Moesian domain with widely occurring carbonate rocks in these layers (Tari et al., 1997) and Poisson's ratio observations along the VRANCEA'99 seismic refraction line seem to confirm this as well (Raileanu et al., 2005). 7.2. The structure of the crystalline crust The tenth geological unit (seismic layer L7 west of halfway between shot points U and T and east of shot point R; seismic layer L8 west of halfway between shot points W and U and seismic layers L9 and L10 in Fig. 11) makes up the crystalline crust of the TiszaDacia, Moesian and North Dobrogea crustal blocks (violet, pink and gray–blue units in Fig. 12). It shows thickness variations related to the different crustal blocks, while there are only small lateral velocity variations (especially in the middle crust) along the entire seismic line. The velocities increase from 6.0 km/s at the top of basement to approximately 7.0 km/s at the Moho. The total thickness of the crystalline crust lies between 30 and 34 km for the western part of the model, which corresponds to the Tisza-Dacia crustal block. This block is characterised by basement thrusts in its upper

crustal layers down to 12–15 km depth as described in the previous section. A low-velocity layer (L8 with 6.0 km/s in Fig. 11) between 11 and 15 km depth is interpreted as being related to another intra-crustal basement thrust connected to the Carpathian Orogen, where a higher-velocity deeper crustal unit (L7 with 6.0– 6.2 km/s in Fig. 11) was thrusted over a lower-velocity shallower crustal unit (Fig. 12). Between shot point W and east of shot point U, the crystalline crust of the seismic layer L7 and the western part of layer L8 cover lower velocity rocks of the ninth geological unit (Fig. 12). We interpret this again as crystalline basement overthrusted on top of Palaeozoic sedimentary rocks inside the Tisza-Dacia block. This structure is also seen in the 3D crustal tomography model of Landes et al. (2004) and has been correlated with a SSE-ward directed Late Pliocene/Early Pleistocene out-of-sequence basement thrust (Landes et al., 2003; Fielitz and Seghedi, 2005). The middle crust comprises velocities of 6.3– 6.5 km/s at depth between 15 and 29 km, whereas the lower crust with velocities of 6.8–7.0 km/s is thinner and has a thickness of only between 8 and 10 km (Fig. 11). As discussed later on (Section 7.3.), the middle and lower crust of the Tisza-Dacia block have their eastern boundary between shot points U and T (Fig. 12). And like for the upper crust, there is also no indication of the proposed Miocene suture zone in the middle and lower crust, since velocities seem to be laterally and horizontally relatively continuous and lower than to be expected for mafic or ultramafic ophiolitic rocks. The velocity model gives also no indication of voluminous Neogene to Quaternary volcanic rocks (basalts and andesites) as tentatively proposed in the geological section of Fig. 4 for this part of the seismic profile. Small-volume dikes or sub-volcanic intrusions are however still possible, but cannot be resolved. To the east, until shot point R, a clearly distinct structure in the central part of the model is associated with the Moesian crustal block. The total thickness of the Moesian crystalline crust is 19 to 25 km. The only 7–9 km thick middle crustal layer with the top between approximately 20 and 22 km depth shows velocities of 6.1–6.3 km/s, while the 9–16 km thick lower crustal layer has velocities of 6.7–7.1 km/s. The eastern boundary of this Moesian crustal block correlates well with the down-dip prolongation of the Peceneaga– Camena crustal fault (Figs. 2 and 12). As described in Section 7.1, above the thinned middle to lower crust, the Moesian block is characterised by alternating high- and low-velocity layers (layers L1–L8 in Fig. 11), which we interpret from bottom to top as a thick Palaeozoic autochthonous sedimentary cover, a Permo-Triassic

F. Hauser et al. / Tectonophysics 430 (2007) 1–25

graben structure, a Mesozoic autochthonous sedimentary cover and a deep Middle Miocene to Quaternary sedimentary basin. The superposition of these different structures suggests a repeated reactivation of possibly Palaeozoic or older major crustal discontiniuties within the Moesian block. The verification of this geological model, however, has to take into account the 3Dorientation of the individual structures, which cannot be distinguished on this two-dimensional section. In addition, especially the deeper structures, which do not outcrop at the surface, are not yet clearly established. East of the Peceneaga–Camena Fault, the North Dobrogea crustal block shows a very distinct threelayered crystalline crust with a total thickness of 44 km. The upper crust has velocities of 6.0–6.2 km/s (seismic layer L7), while middle and lower crustal velocities (seismic layer L9 and L10) range between 6.3–6.4 km/s and 6.7–7.1 km/s, respectively. The thick North Dobrogea crystalline crust is connected with the Scythian Platform, which is a continuation of the East European Platform further to the east and northeast, as shown by large-scale tomographic data from this area (Wortel and Spakman, 2000). We therefore suggest, that the North Dobrogea crustal block is mainly composed of crust from the Scythian Platform (see Fig. 1) and that the uppermost layer of 1 to 2 km thickness represents the central or frontal parts of the overthrusted wedge of the North Dobrogea Orogen. Layers L7, L9 and L10 show no lateral velocity variations between shot points R and O. Therefore, an important deep reaching ultramafic nappe, as proposed in the geological section of Fig. 4, seems improbable. Additionally the velocity model in this area is sub-parallel to the geological structures. For this reason such a nappe would not be a steeply dipping structure but a near-surface horizontal to shallow dipping body. Velocities of 5.8 and 5.9 km/s in the thrust wedge of seismic layer L5 could relate to such an ultramafic nappe, whose down-dip continuation must however be searched south and parallel to the actual seismic section. 7.3. Plate boundaries As discussed in the introduction, the geodynamic setting of the region covered by the VRANCEA2001 seismic line is thought to relate to the final stages of a subduction process (e.g. Sandulescu, 1988; Csontos, 1995; Girbacea and Frisch, 1998; Sperner et al., 2001; Cloetingh et al., 2004). This subduction involved the upper Tisza-Dacia lithospheric plate, which was already affected by important contractional deformation (thrusts, nappes and a palaeosuture) related to an earlier Early Cretaceous subduction and collisional event. The

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composite lower plate originated by the accretion of different lithospheric domains, which involved the relatively undeformed but compositionally distinct Moesian, Scythian and East European platform areas as well as the Late Triassic to Late Jurassic North Dobrogea Orogen with its Variscan basement, all separated by important crustal faults. The collision and climax of deformation between both plates took place in the Middle to Late Miocene (Sarmatian) and resulted in a steep suture zone thought to be concealed by the overthrusted Eastern Carpathian flysch nappes, which would represent the unrooted accretionary prism. This model is represented in the geological section of Fig. 4. The VRANCEA2001 profile can be subdivided into three crustal domains with distinct characteristics concerning thickness, composition, structuring and geometry of the different crustal layers (Fig. 13). The western domain, which we relate to the TiszaDacia plate, has the thinnest crust with a Moho depth of 37–33 km. The middle crust is significantly thicker than the lower crust. The upper crust shows an alternation of high and low velocity zones, which we interpret as largely related to imbricated thrust sheets with alternating sedimentary and crystalline rocks. Much of this deformation could already have been emplaced during the earlier Early Cretaceous subduction and collisional event with some reactivation and new deformation during the later Sarmatian event. In this domain Bouguer anomalies show negative values decreasing across the Transylvanian Basin with the lowest values due to thick sediments around shot point U (− 50 mgal) and shot point Y (− 60 mgal; Visarion, 1998). Positive values of the magnetic anomaly component ΔZ of up to 300 gamma dominate the centre of the Transylvanian Basin, while negative values of up to −100 gamma were observed between shot points X and R (Airinei et al., 1985). These anomalies are related to different basement compositions (greenschist basement with negative anomalies and other basements with positive anomalies; Airinei et al., 1985). The central domain, which we relate to the Moesian plate, has a thick crust with a Moho depth down to 45 km. Here the middle crust (with slightly lower velocities than the adjacent plates) is thinner than the lower crust. The upper crust is well layered, wholly sedimentary and shows in its lower part also an alternation of high and low velocity zones. Deformation seems to be limited to Permo-Triassic extension and Miocene to Quaternary subsidence in the Carpathian foreland. In this domain the Bouguer anomalies show negative values with a small minimum of − 85 mgal around shot point T and an absolute minimum of

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− 100 mgal between shot points S and T in the Focsani Basin, where the sedimentary basin is deepest (Visarion, 1998). Variations of the magnetic anomaly component ΔZ are similar to the western domain with negative values of up to − 100 gamma (Airinei et al., 1985). The eastern domain, which we relate to the North Dobrogea Orogen, has a thick crust with a Moho depth of 44 km. Here the middle and lower crust are relatively thick, the latter being somewhat thicker. The upper crust is relatively thick, very homogeneous and crystalline, except for the very thin uppermost thrust wedge of mixed composition related to the Triassic to Jurassic deformation of the orogen. The thick crystalline crust would be the continuation of the Scythian/East European platform. In this domain the Bouguer anomalies show positive values (Visarion, 1998) and the magnetic anomaly component ΔZ rises to almost 200 gamma between shot points R and P and to approximately 150 gamma at the eastern end of the line (Airinei et al., 1985). The boundaries between the three plates are steep and relatively sharp (Figs. 12 and 13). The Peceneaga– Camena Fault between the North Dobrogea and the Moesian plate is a well known crustal discontinuity and well defined from surface geology and geophysical data (Visarion et al., 1988; Radulescu and Diaconescu, 1998; Seghedi, 1998; Matenco et al., 2003; Tarapoanca et al., 2003). The boundary between the Moesian and TiszaDacia plates is generally poorly constrained, because it is concealed by the Sarmatian overthrusted flysch nappes of the Eastern Carpathians (Figs. 4, 12 and 13). The proposed Miocene suture (Sandulescu, 1988; Girbacea and Frisch, 1998; Sperner et al., 2001) cannot be identified in the VRANCEA2001 profile, neither in the location shown in Fig. 4 nor further to the east. The only sharp boundary separating two distinct crustal domains is found between shot points U and T (Fig. 12). Therefore, we tentatively interpret this crustal discontinuity to be the boundary between the Tisza-Dacia and the Moesian plates. There is, however, no indication of a suture zone, since crustal velocities do not point to mafic or ultramafic rocks. The nature and detailed geometry of this contact is not known, but an alternative could be a crustal fault with lateral displacement, eventually a transfer zone reactivating older (Permo-Triassic ?) crustal structures. From the conventionally proposed subduction models important horizontal displacements between the Tisza-Dacia and Moesian plate would generally be expected, also for deeper parts of the

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crust. This cannot be confirmed from the presented plate characteristics (Fig. 13). It has however to be taken into account for the correlation, geometry and interpretation of the presented data that the orientation and age of the crustal structures change considerably between and inside the involved plates (Tisza-Dacia with Early Cretaceous and Sarmation deformation, Moesia with Triassic–Permian and Miocene–Quaternary deformation, North Dobrogea with Variscan and Triassic– Jurassic deformation, Scythian/East European platform with Precambrian deformation) and that the 3Dorientation of the individual structures can be difficult to distinguish in a two-dimensional section. Additionally, structures in the probably mostly Precambrian crystalline crust might be concealed because possible compositional and therefore structural differences might not be shown by differences in the seismic velocities. Also no obvious relation to the steep Vrancea seismic body can be seen. This could be because of decoupling of crustal and mantle processes. The three plates are generally thought to belong to the southeastern prolongation of the Trans-European suture zone (TESZ; e.g. Pharao, 1999; Debacker et al., 2005) and therefore the VRANCEA2001 profile also crosses this major plate boundary. The overall geometry of the presented velocity model shows a high degree of similarity to the velocity models and seismic profiles across the TESZ in Poland further to the northwest (e.g. Jensen et al., 2002; Janik et al., 2002; Grad et al., 2002). These similarities consist mainly in a thick three-layered crust of the Precambrian Craton (42–45 km in Poland, 44 km in North Dobrogea), a thinner crust with a thin (∼8 km) lower crust in the areas to the southwest (29– 32 km in the Palaeozoic terranes of Poland, 37–33 km in the Tisza-Dacia terran) and the TESZ itself is covered by a deep sedimentary basin with Permian origins or precursors (20 km thick Polish Basin, 22 km thick Focsani Basin area). This suggests strongly a southeastward prolongation of the TESZ structure into Romania along the southwestern margin of the East European Precambrian craton. However, there are also several differences, which have to be considered: The POLONAISE profiles cross the TESZ perpendicular to their overall structures, whereas the VRANCEA2001 profile is highly oblique to it. In Poland the southwestern terranes belong to Avalonia, which experienced Caledonian and Variscan deformation. In Romania they belong to the Moesian terran with its still poorly understood Palaeozoic evolution and to the North

Fig. 13. Geological characteristics and crustal thicknesses of the main crustal domains (plates) along the VRANCEA2001 seismic refraction line. Steep boundaries between the deduced Tisza-Dacia, Moesia and North Dobrogea plates seem to be recognizable although important horizontal displacements between the Tisza-Dacia and Moesian plate would generally be expected.

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Dobrogea orogen, which experienced Variscan and Triassic–Jurassic deformation. The Tisza-Dacia terran, which makes up the whole western half of the VRANCEA2001 profile and already experienced an Early Cretaceous collisional event, is generally thought to have collided with the TESZ only during the Miocene and, separated by an oceanic domain, was formerly located much farther to the west. Therefore its crustal geometry cannot be compared easily with the Avalonia terran. Also the younger basins show marked differences. The Polish Basin is a Carboniferous–Permian and Mesozoic structure that was inverted during the Late Cretaceous and Early Tertiary. The Focsani Basin is mainly a Late Cenozoic structure marginally affected by Late Alpine deformation. It possibly had a Permo-Triassic precursor basin, but its geometry and relation to the overlying Cenozoic basin is only very poorly constrained. In summary, the POLONAISE and VRANCEA2001 profils show globally many similarities, expecially due to the contrast between the exceptionally thick East European Precambrian crust and the thinner southwestern accreted terranes. Timing of accretion and deformation of these terranes might, however, be very different and crustal and basinal similarities partly result only from the mechanical differences between both crustal domains. A more in-depth interpretation of the VRANCEA2001 is only possible in the context of a complex larger scale geodynamic model using additional geological and geophysical (e.g. mantle tomography; Wenzel et al., 1998b; Martin et al., 2005, 2006) data. This is however not the focus of this paper. 8. Conclusions A 700 km long WNW–ESE trending seismic refraction line was carried out in Romania in order to study the lithospheric structure. Here we present results from a subsection between the Transylvanian Basin across the SECarpathians to the Carpathian foreland areas. The geophysical and geologic interpretation of the data by forward and inverse modeling gave the following results: The sedimentary succession can be subdivided into 7 layers with a total thickness of up to 22 km. It is composed of (1) the Carpathian nappe pile, (2) the postcollisional (post-Early Cretaceous) Paleo to Neogene Transylvanian Basin, which covers the local Late Cretaceous to Paleogene Tarnava Basin, (3) the Neogene Focsani Basin in the foredeep area, which covers autochthonous Mesozoic and Palaeozoic sedimentary rocks as well as a proposed Permo-Triassic graben structure of the Moesian Platform, and (4) the Palaeo and Mesozoic rocks of the North Dobrogea Orogen.

The underlying crystalline crust shows considerable thickness variations, in total as well as in its individual subdivisions, which correlate well with the Tisza-Dacia, Moesian and North Dobrogea crustal blocks, respectively. Only minor lateral changes in velocity structure of these blocks were observed. The Tisza-Dacia block is about 35 km thick and low velocity zones in its uppermost 15 km are presumably basement thrusts imbricated with sedimentary successions related to the Carpathian Orogen. The crystalline crust of Moesia does not exceed 23 km and is covered by up to 22 km of sedimentary rocks. The North Dobrogea crust reaches a thickness of about 44 km including an up to 2 km thick mixed sedimentary-volcanic-crystalline cover, which is mainly composed of a thin overthrusted wedge of the North Dobrogea Orogen. The presented velocity model intersects the TransEuropean suture zone (TESZ) and shows a high degree of similarity in its overall geometry and velocities to the velocity models and seismic profiles across the TESZ in Poland further to the northwest, although the specific crustal evolution of both areas appears to have clear differences. Acknowledgements This investigation was only possible by the continuous effort of many volunteers, in particular students from the Universities of Amsterdam, Bucharest and Karlsruhe. The National Institute for Earth Physics (NIEP) and the University of Bucharest (Geology and Geophysics Department) provided the logistics for the fieldwork in Romania. The Romanian Exploration Company PROSPECTIUNI S.A., Bucharest, was responsible for the environmental study as well as the drilling and shooting operations. Data were collected using the seismic equipment of the geophysical instrument pool of the GeoForschungsZentrum Potsdam (150 units) as well as the joint pool of IRIS /PASSCAL at Socorro, New Mexico and the University of Texas at El Paso (640 units). The Deutsche Forschungsgemeinschaft (German Science Foundation) funded the project through the Collaborative Research Centre 461 (CRC 461) at the University of Karlsruhe, Germany: “Strong Earthquakes — a Challenge for Geosciences and Civil Engineering”. The Romanian Ministry for Education and Research funded the Romanian researchers in this project via the CERES program (CERES 1 no. 34/ 2001 and CERES 4 no. 38/2004). The NATO Science Collaborative Research Linkage Grant no. EST.CLG 974792 assisted the project by additional travel funding. Laszlo Csontos and an anonymous reviewer are

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