Hauser Et Al Tectonophysics2007

25
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  b  National Institute for Earth Physics, P .O.Box MG-2, RO-0771 25 Bucur esti-Magur ele, Romania c Geological Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany d  Faculty of Geology and Geophysics, University of Buchar est, 6, Traia n Vu ia Street, sector 1, RO-7013 9 Bucur esti, Romania Received 11 May 2005; received in revised form 17 October 2006; accepted 19 October 2006 Available online 11 December 2006 Abstract In or de r to st udy the li thospher ic str uc ture in Roma ni a a 450 km long WNWESE tren din g seis mic refr act ion pro ject was car ried out in August/September 2001. It runs from the Transylvanian Basin across the East Carpathian Orogen and the Vrancea seismic region to the forel and ar ea s wit h the ve ry deep Ne og ene Focsani Ba sin and the North Dobr ogea Or ogen on the Bla ck Sea. A total of ten shot s wi th charge sizes 3001500 kg we re recorded by over 700 geop hones. The data quality of the experimen t was variable, depend ing primarily on cha rge sizebut als o on loc al geo log icalcondi tions. The dat a inte rpr eta tion indica tes a mul ti- layere d stru cture withvaria ble thi ckn esse s and veloc ities. The sedime ntary stac k comprise s up to 7 layers with seismi c velocitie s of 2.0 5.9 km/s. It reaches a ma ximum thickness of about 22 km within the Focsani Basin area. The sedimentary succession is composed of (1) the Carpathian nappe pile, (2) the post- collis ional Neog ene Tra nsylv anian Ba sin, which cover s the local Late Cretac eous to Paleog ene Ta rnava Ba sin, (3) the Ne ogen e Focsan i Basin in the foredeep area, which covers autochthonous Mesozoic and Palaeozoic sedimentary rocks as well as a probably Permo- Triassic graben structure of the Moesian Platform, and (4) the Palaeozoic and Mesozoic rocks of the North Dobrogea Orogen. The und erly ing crysta llin e crust sho ws con side rable thic kne ss var iati ons in tota l as well as in its indivi dua l sub divisi ons , whi ch cor rela te well with the Tisza-Dacia, Moesian and North Dobrogea crustal blocks. The lateral velocity structure of these blocks along the seismic line remains cons tant wi th about 6.0 km/s al ong the baseme nt top and 7. 0 km/s abov e the Moho. The Ti sza- Daci a bl ock is about 33 to 37 km thick and shows low vel oci ty zones in its upp ermost 15 km, whi ch are pres umabl y due to bas eme nt thru sts imbr icat ed with sed ime ntar y 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 12 km thick wedge of the North Dobrogea Orogen. © 2006 Elsevier B.V. All rights reserved.  Keywor ds:  Seismic refraction; Crustal velocity structure; Vrancea zone; Eastern Carpathians; Moesian Platform; Transylvanian Basin 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 Tectonophysics 430 (2007) 1 25 www.elsevier.com/locate/tecto  Corres ponding aut hor . Tel.: +49 721 6082139; fax: +49 721 6082138.  E-mail address: werner.fieli [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

Transcript of Hauser Et Al Tectonophysics2007

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 1/25

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 b  National Institute for Earth Physics, P.O.Box MG-2, RO-077125 Bucuresti-Magurele, Romania

cGeological Institute, University of Karlsruhe, Hertzstr. 16, D-76187 Karlsruhe, Germany

d  Faculty of Geology and Geophysics, University of Bucharest, 6, Traian Vuia Street, sector 1, RO-70139 Bucuresti, Romania

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 sizebut also on local geologicalconditions. The data interpretation indicates a multi-layered structure withvariable 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 post-

collisional 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 Permo-

Triassic 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

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

Tectonophysics 430 (2007) 1 – 25

www.elsevier.com/locate/tecto

⁎ 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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 2/25

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 thenappes of the Carpathian Orogen or the Neogene

sedimentary basins. Whereas these plates were conso-

lidated 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 theFocsani 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 (Mwb5.6), 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 largemagnitudes (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 StrongEarthquakes (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 intermedi-

ate-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 experi-

ments 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 addi-

tional 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 TransylvanianBasin 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 theeast 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-Cretaceoustime (Aptian/Albian), were then partially covered by

smaller basins with mid-Cretaceous (Aptian/Albian) to

Palaeogene sedimentary strata and were finally trans-

 ported to their present position during the Sarmatian

accretion of the Moldavides against Europe (Sandu-

lescu, 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 deforma-

tion is inadequately understood, but it appears that 

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 3/25

Fig. 1. Geological overview of the Eastern Carpathian bend area and its foreland with the main crustal units, nappe structures, faults and basins. The locati

transverse) and VRANCEA2001 NE–SW seismic refraction profiles are shown with their shot points. Compiled from various sources given in the text

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 4/25

subduction with slab break-off and lithospheric dela-

mination 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 andFrisch, 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 post-

collisional processes after subduction to explain the

actual geodynamic scenario, which are controlled by the

thermo-mechanical heterogeneities within the under-

thrusted 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 estimatedthickness 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.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 5/25

Bertotti, 2000). This thickness is constrained by surface

geology, seismic reflection and borehole data, mainly

from oil exploration, but for the inner nappes thick-

nesses are only estimated through balanced cross-

sections. 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 meta-

morphic 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 thecentral part of the foredeep (e.g. Tarapoanca et al., 2003).

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; Raileanuet 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 poly-

deformed 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 pro-

longation 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 faultsas 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—

EasternCarpathian nappes, 8 — volcanic rocks, 9 — faults, 10 — crustal and subcrustal earthquake foci). For location see Fig. 2.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 6/25

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 post-

collisional Late Miocene to Quaternary calc-alkalinevolcanic Calimani–Gurghiu–Harghita (CGH) moun-

tains (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 exceptfor 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

andMoesia is from Sandulescu (1988), Girbacea and Frisch (1998) and Sperner et al.(2001). Neogene to Quaternary andesites are tentativelyprojected

down-strike the East Carpathian Calimani–

Gurghiu–

Harghita volcanic chain into the cross-section to indicate the possibility of occurrence of suchrocks in the crust. The Neogene filling of the Transylvanian Basin is based on Ciulavu et al. (2000).

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 7/25

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-surfacegeology 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). Thiscross-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 velocity-

depth model obtained from the seismic experiment as

 presented in this paper. It can later be integrated into a

complex larger scale geodynamic model using additionalgeological 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, magneto-

telluric and heat flow measurements, seismic reflection

and refraction data played an important role. Here we

summarize the results, which are relevant for theVRANCEA2001 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 

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). Thecrustal 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 ( K 0 boundary), and an intra-crustal boundary

( K 1) 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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 8/25

interface K 1 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 the1970'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 ( K 1)

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 K 0-boundary

lies at 7 km, K 1 at 20 km and the Moho at 40 km depth.

In the Transylvanian Basin K 0 is at 3 km, K 1 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 cross-

section parallel to the VRANCEA2001 line with up to

scale projections of the Geotraverse XI, RmS and Tr1lines 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 seismo-

genic “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 experi-

ment, 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 crust-mantle boundary (Moho); Pn = diving wave through the upper mantle; Diff. = supposedly diffracted waves.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 9/25

completion of the first north–south running VRAN-

CEA'99 line, briefly described above, it was the second

experiment to target the deep structure of the seismo-

genic Vrancea area with the aim of testing new and

existing geodynamic models.

The experiment recorded refraction/wide-angle re-

flection 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 geo-

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

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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 10/25

 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 SouthCarolina, USA. Field recording instruments were

 provided by University of Texas at El Paso and the

IRIS/PASSCAL instrument pool, USA (TEXAN 1-

component 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.

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

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

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

Table 1

Summary of data quality as a function of offset for each shot point 

 No. Shot 

 point 

Branch Good Acceptable Poor 

Distance

(km)

Distance

(km)

Distance

(km)

1 Z  West  – – –

East 0–140 140–220

2 Y  West 0–70 – –

East 0–95 – 95–175

3 X  West 0–40 – –

East 0–30 – 30–130

4 W  West 0–100 – 100–145East 0–90 – 90–130

5 U  West 0–50 – 50–150

East 0–40 – 40–100 very poor;

170–220 poor PmP

6 T  West 0–95 – 95–135

East 0–40 – 40–85; 120–200

7 S  West 0–105 105–215 215–290

East 0–50 50–130 130–155

8 R West 0–105 – 105–205

East 0–120 – –

9 P  West 0–70 – 70–235

East 0–70 – –

10 O West 0–95 95–220 220–285

East 0–30 –

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 11/25

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 asegment 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 sec-

tions, 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 b6 km/s could be identified.

Table 2

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

P1 P2 P3 P4 P5 P6P Pg1 Pi1P P8P Pg2 Pi2P PmP Pn

O–e – – – – 3 – 3 2–3 – – 1 1 –

O–w – – – – 3 – 3 2–3 – 2 2 2–3 1

 P –

e– – – –

3–

3 2–

3 – – 2 1–

 P –w – – – – 3 – 3–2 2 – – 2 1 –

 R–e 3 – – – 3 – 3 2 – – 2 1 –

 R–w 3 – – – 3 – – 2–1 – – 1–2 – –

S –e 3 3 3 3 2–1 – 2 2 – – 2 2 –

S –w 3 3 3 3 2 3–2 – 2–3 3–2 2 2 3 3

T –e – 3 – 3 – 2 – 2 1 – 1 1 1

T –w – 3 – 3 2 2–3 2–3 2–1 2–1 – 2–1 2–3 –

U –e – 3 – 3–2 2 3 – 1 1 – – 1 –

U –w – 3 – 3 2 2–3 3–2 2 1–2 – 2–1 2–1 –

W –e – 3 – 3 2–3 2–3 3 2–3 1–2 – 2 2 –

W –w – 3 – 3–2 3–2 3 3 3–2 2 2 2–1 1 –

 X –e – 3 – 3 2 2 – 1 1 – – 1 –

 X –w – 3 – 3 2 3 2 2 2 – 2–1 1 –

Y –e – 3 – 3 3 2 3–2 2 2 1–2 2 1 –

Y –w – 3 – 3 2–3 – 2 2 2 – 1 1 –

 Z –e – 3 – 3 3 – 3 2–3 2 3 2 2–1 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.

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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 12/25

They have been labelled P1 to P5 and are attributed to

waves travelling through different sedimentary layers or 

intercalations of sedimentary, metamorphic, and volca-

nic layers (Figs. 5–9). Most of the data show a clear first 

arrival Pg1 phase (a diving wave through the upper crust)

from 20 kmto over 100 km offset (Figs. 5 and 7–9). This phase is characterised by strong undulations and some-

times 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 sub-

critical 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.

6. Interpretation techniques and velocity model

For the interpretation of the P-wave seismic data only

vertical component seismograms were used. In order to

find the model that fits the seismic data best we used

several major steps in the modelling procedure:

(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 ray-

tracing techniques.

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 usedthe 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 infor-

mation 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 deter-

mined using the appropriate seismic phases. This is

followed by stepwise modelling of phases from deeper 

layers, usually in a sequence in which the velocity isdetermined 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.

6.1. Uncertainties

A total of 13 seismic phases were correlated and

 picked (see Table 2), which already indicates the

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 13/25

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

±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.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 14/25

data points was carried out. The final model gave a root-

mean-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. 10 b.

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 b6 km/s; (2) the crystalline crust 

down to the crust-mantle boundary (Moho).

The sedimentary succession with imbricated crystal-

line 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 byreflections and/or refractions. Dashed lines indicate less well constrained areas, while thin lines are extrapolations (see also Table 2).

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 15/25

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 betweenhigher-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 areattributed 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 westernend 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 (be-

tween 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 NorthDobrogea Orogen area in the east. No pronounced

crustal root below the Carpathian Orogen is recogni-

zable. Some constraints on upper mantle seismic velo-

cities 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 

crustal blocks characterised by clearly distinct geome-

tries 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 constraintsfor 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 andPalaeozoic 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 thevelocity 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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 16/25

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 be-

tween 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).

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.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 17/25

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 betweenshot 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 heteroge-

neous composition with participation of lower-grade

metamorphic, possibly Palaeozoic rocks, and a localizedthin cover of autochthonous Palaeo and Mesozoic rocks

could explain the velocity variations, especially the low

velocities (b4.9 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 compo-

sition of the crust. There is also no indication of the

 proposed Miocene suture zone below the third geolo-

gical unit (Eastern Carpathian flysch nappes) since

velocities seem to be laterally and horizontally relativelycontinuous 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 cross-

section 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 

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) partiallyunderlies 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 fromsurface 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 isabout 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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 18/25

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-Triassicgraben, 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 occurringcarbonate 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 inFig. 11) makes up the crystalline crust of the Tisza-

Dacia, 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 varia-

tions (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 bet-

ween 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 base-ment 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 horizon-

tally relatively continuous and lower than to be expectedfor 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 only7–9 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 au-

tochthonous sedimentary cover, a Permo-Triassic

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 19/25

graben structure, a Mesozoic autochthonous sedimen-

tary 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 geologicalmodel, however, has to take into account the 3D-

orientation 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 three-

layered 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/sand 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 andO. 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

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, allseparated 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 Tisza-

Dacia 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 alterna-

ting 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 decreasingacross 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 Δ Z 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 

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 20/25

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 21/25

−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 

Δ Z  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 NorthDobrogea 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 deforma-

tion 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 Δ

 Z  rises to almost 200 gamma betweenshot 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 Tisza-

Dacia plates is generally poorly constrained, because it is

concealed by the Sarmatian overthrusted flysch nappesof 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 disconti-

nuity 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 maficor 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

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 EarlyCretaceous and Sarmation deformation, Moesia with

Triassic–Permian and Miocene–Quaternary deforma-

tion, North Dobrogea with Variscan and Triassic–

Jurassic deformation, Scythian/East European platform

with Precambrian deformation) and that the 3D-

orientation of the individual structures can be difficult 

to distinguish in a two-dimensional section. Additional-

ly, 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 pre-

cursors (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 Precam- brian 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 horizontaldisplacements between the Tisza-Dacia and Moesian plate would generally be expected.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 22/25

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 Mioceneand, 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 VRAN-

CEA2001 is only possible in the context of a complex

larger scale geodynamic model using additional geo-

logical 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 kmlong WNW–ESE trending seismic refraction

line was carried out in Romania in order to study the

lithospheric structure. Here we present results from a sub-

section between the Transylvanian Basin across the SE-

Carpathians to the Carpathian foreland areas. The

geophysical and geologic interpretation of the data byforward 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 post-

collisional (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, respec-

tively. Only minor lateral changes in velocity structure

of these blocks were observed. The Tisza-Dacia block isabout 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 Trans-European 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 conti-

nuous effort of many volunteers, in particular students

from the Universities of Amsterdam, Bucharest andKarlsruhe. 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 re-

sponsible 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 ElPaso (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

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 23/25

thankend for their careful and constructive reviews,

which helped to improve and clarify parts of the paper.

References

Airinei, St., 1977. Lithosphere microplates on the Romanian territoryreflected in regional gravimetric anomalies (in Romanian). Stud.

Cercet. Geol. Geofiz. Geogr., Ser. Geofiz. 15 (1), 19 –30.

Airinei, S., Stoenescu, S., Velcescu, G., Romanescu, D., Visarion, M.,

Radan, S., Roth, M., Besutiu, L., si Besutiu, G., 1985. Distributia

anomaliilor magnetice Δ Z a pe teritoriul Romaniei. Stud. Cercet.

Geol. Geofiz. Geogr., Ser. Geofiz. 23, 12–19.

Badescu, D., 1998. Geology of the East Carpathians — an overview.

CERGOP “South Carpathians” monograph, vol. 7 (37), pp. 49–69.

Warszawa.

Bocin, A., Stephenson, R., Tryggvason, A., Panea, I., Mocanu, V.,

Hauser, F., Matenco, L., 2005. 2.5D seismic velocity modelling in

the southeastern Romanian Carpathians Orogen and its foreland.

Tectonophysics 410, 273–291.

Ciulavu, D., Tertiary tectonics of the Transylvanian basin, Ph.D. thesis,154 pp., Vrije Univ., Amsterdam, Netherlands, 1999.

Ciulavu, D., Dinu, C., Szakács, A., Dordea, D., 2000. Neogene

kinematics of the Transylvanian basin (Romania). AAPG Bull. 84

(10), 1589–1615.

Cloetingh, S., Horvath, F., Dinu, C., Stephenson, R.A., Bertotti, G.,

Bada, G., Matenco, L., Garcia-Castellanos, D., and the TECTOP

working group, 2003. Probing tectonic topography in the after-

math of continental convergence in central Europe. EOS 84,

89–93.

Cloetingh, S.A.P.L., Burov, E., Matenco, L., Toussaint, G., Bertotti,

G., Andriessen, P.A.M., Wortel, M.J.R., Spakman, W., 2004.

Thermo-mechanicalcontrols on the mode of continental collision in

the SE Carpathians (Romania). Earth Planet. Sci. Lett. 218, 57–76.

Constantinescu, L., Enescu, D., 1984. A tentative approach to possibly

explaining the occurrence of Vrancea earthquakes. Rev. Roum.

Geol. Geophys. Geogr., Ser. Geophys. 28.

Constantinescu, L., Cornea, S., Lazarescu, D., 1973. An approach to

the seismotectonics of the Romanian Eastern Carpathians. Rev.

Roum. Geol. Geophys. Geogr., Ser. Geophys. 17 (2), 133–143.

Cornea, I., Radulescu, F., Pompilian, Al., Sova, A.,1981. Deep seismic

soundings in Romania. Pure Appl. Geophys. 119, 1144–1156.

Csontos, L., 1995. Tertiary tectonic evolution of the Intra-Carpathian

area: a review. Acta Vulcanol. 7, 1–13.

Debacker, T., Sintubin, M., Verniers, J. (Eds.), 2005. Avalonia–

Moesia: Early Palaeozoic orogens in the Trans-European Sutue

Zone. Geologica Belgica, vol. 8 (4), pp. 1–192.

Downes, H., Seghedi, I., Szakács, A., Dobosi, G., James, D.E., Vaselli,O.,Rigby, I.J., Ingram, G.A., Rex, D., Pécskay, Z., 1995. Petrology

and geochemistry of late Tertiary/Quaternary mafic alkaline

volcanism in Romania. Lithos 35, 65–81.

Ellouz, N., Roure, F., Sandulescu, M., Badescu, D., 1994. Balanced

cross-sections in the Eastern Carpathians (Romania): a tool to

quantify Neogene dynamics. In: Roure, F., Ellouz, N., Shein, V.S.,

Skvortsov, I. (Eds.), Geodynamic Evolution of Sedimentary

Basins, pp. 305–325.

Enescu, D., Cornea, I., Constantinescu, P., Radulescu, F., si Patrut, S.,

1972. Structura scoartei terestre si a mantalei superioare in zona

curburii Carpatilor. Stud. Cercet. Geol. Geofiz. Geogr., Ser. Geofiz.

10 (1), 23–41.

Enescu, D., Danchiv, D., Bala, A., 1992. Lithosphere structure in

Romania II. Thickness of the Earth crust. Depth-dependent 

 propagation velocity curves for the P and S waves. Stud. Cercet.

Geol. Geofiz. Geogr., Ser. Geofiz. 30, 3–19.

Fielitz, W., Seghedi, I., 2005. Late Neogene to Quaternary tectonic

geomorphology and river drainage evolution in the Eastern

Carpathian Bend area of Romania. Tectonophysics 410, 111–136.

Fuchs, K., Bonjer, K.-P., Bock, G., et al., 1979. The Romanian

earthquake of March 4, 1977; II, aftershocks and migration of seismic activity. Tectonophysics 53, 225–247.

Girbacea, R., Frisch, W., 1998. Slab in the wrong place: lower 

lithospheric mantle delamination in the last stage of the Eastern

Carpathian subduction retreat. Geology 26 (7), 611–614.

Grad, M., Jensen, S.L., Keller, G.R., Guterch, A., Thybo, H., Janik, T.,

Tiira, T., Yliniemi, J., Luosto, U., Motuza, G., Nasedkin, V.,

Czuba, W., Gaczynski, E., Sroda, P., Miller, K.C., Wilde-Piórko,

M., Komminaho, K., Jacyna, J., Korabliova, L., 2002. Crustal

structure of the Trans-European suture zone region along

POLONAISE'97 seismic profile P4. J. Geophys. Res. 108 (B11),

2541. doi:10.1029/2003JB002426.

Hauser, F., Raileanu, V., Fielitz, W., Bala, A., Prodehl, C., Polonic, G.,

Schulze, A., 2001. VRANCEA'99-the crustal structure beneath SE

Carpathians and the Moesian Platform from a seismic refraction profile in Romania. Tectonophysics 340, 233–256.

Hauser, F., Prodehl, C., Landes, M., VRANCEA working group, 2002.

Seismic experiments target earthquake-prone region in Romania.

EOS Trans. AGU 83, 457–463.

Hippolyte, J.C., Sandulescu, M., 1996. Paleostress characterization of 

the “Wallachian” phase in its type area, southeastern Carpathians,

Romania. Tectonophysics 263, 235–249.

Hippolyte, J.C., Badescu, D., Constantin, P., 1999. Evolution of the

transport direction of the Carpathian belt during its collision with

the east European Platform. Tectonics 18, 1120–1138.

Hole, J.A., 1992. Nonlinear high-resolution three-dimensional seismic

travel time tomography. J. Geophys. Res. 97, 6553–6562.

Janik, T., Yliniemi, J., Grad, M., Thybo, H., Tiira, T., and

POLONAISE P2 Working Group, 2002. Crustal structure across

the TESZ along POLONAISE'97 seismic profile P2 in NW

Poland. Tectonophysics 360, 129–152.

Jensen, S.L., Thybo, H., POLONAISE'97 Working Group, 2002.

Moho topography an lower crustal wide-angle reflectivity around

the TESZ in southern Scandinavia and northeastern Europe.

Tectonophysics 360, 187–213.

Knapp, J.H., Knapp, C.C., Raileanu, V., Matenco, L., Mocanu, V.,

Dinu, C.,2005. Crustal constraints on theoriginof mantleseimicity

in the Vrancea Zone, Romania: the case for active continental

lithospheric delamination. Tectonophysics 410, 311–323.

Landes, M., Fielitz, W., Hauser, F., Popa, M., CALIXTO Group, 2004.

3-D upper-crustal tomographic structure across the Vrancea

Seismic Zone, Romania. Tectonophysics 382, 85–102.Martin, M., Ritter, J.R.R., and the CALIXTO Working Group, 2005.

High-resolution teleseismic body-wave tomography beneath SE

Romania – I. Implications for three-dimensional versus one-

dimensional crustal correction strategies with a new crustal

velocity model. Geophys. J. Int. 162, 448–460.

Martin, M., Wenzel, F., and CALIXTO working group, 2006. High-

resolution teleseismic body wave tomography beneath SE-

Romania – II. Imaging of a slab detachment scenario. Geophys.

J. Int. 164, 579–595.

Mason, P.R.D., Seghedi, I., Szakács, A., Downes, 1998. Magmatic

constraints on geodynamic models of subduction in the East 

Carpathians, Romania. Tectonophysics 297, 157–176.

Matenco, L.,Bertotti, G.,2000. Tertiarytectonic evolution of the external

East Carpathians (Romania). Tectonophysics 316, 255–286.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 24/25

Matenco, L., Bertotti, G., Cloetingh, S., Dinu, C., 2003. Subsidence

analysis and tectonic evolution of the external Carpathian–Moesian

Platform region during Neogene times. Sediment. Geol. 156, 71–94.

Morley, C.K., 1996. Models for relative motion of crustal blocks

within the Carpathian region, based on restorations of the outer 

Carpathian thrust sheets. Tectonics 15, 885–904.

 Necea, D., Fielitz, W., Matenco, L., 2005. Late Pliocene–Quaternarytectonic geomorphology along the Putna Valley, East Carpathians,

Romania. Tectonophysics 410, 137–156.

 Neugebauer, J., Greiner, B., Appel, E., 2001. Kinematics of the

Alpine–West Carpathian orogen and palaeogeographic implica-

tions. J. Geol. Soc. (Lond.) 158, 97–110.

Oncescu, M.C., 1984. Deep structure of the Vrancea region,

Roumania, inferred from simultaneous inversion for hypocenters

and 3-D velocity structure. Ann. Geophys. 2 (1), 23–28.

Oncescu, M.C., Bonjer, K.-P., Rizescu, M., 1998. Weak and strong

ground motion of intermediate depth earthquakes from the Vrancea

region. In: Wenzel, F., Lungu, D., Novak, O. (Eds.), Vrancea

Earthquakes: Tectonics, Hazard and Risk Mitigation. Kluwer 

Academic Publishers, Dordrecht, Netherlands, pp. 27–42.

Panea, I., Stephenson, R., Knapp, C., Mocanu, V.v., Drijkonongen, G.,Matenco, L., Knapp, J., Prodehl, C., 2005. Near-vertical seismic

reflection image using a novel technique across the Vrancea zone

and Focsani Basin, Romania. Tectonophysics 410, 293–309.

Panaiotu, C.G., Pécskay, Z., Hambach, U., Seghedi, I., Panaiotu, C.E.,

Itaya, T., Orleanu, M., Szakács, A., 2004. Short-lived Quaternary

volcanism in the Per şani Mountains (Romania) revealed by com-

 bined K –Ar and paleomagnetic data. Geol. Carpath. 55, 333–339.

Pharao, T.C., 1999. Palaeozoic Terranes and their lithospheric

 boundaries within the Trans-European Suture Zone, TESZ, a

review. Tectonophysics 314, 17–41.

Polonic, G., 1996. Structure of the crystalline basement in Romania.

Rev. roum. Géophysique 40, 57–70.

Polonic, G., 1998. The structure and morphology of the crystalline

 basement in Romania. CERGOP “South Carpathians” monograph,

vol. 7 (37), pp. 127–131. Warszawa.

Pompilian, A., Radulescu, F., Diaconescu, M., Bitter, M., Bala, A.,

1993. Refraction seismic data in the eastern side of Romania. Rom.

Rep. Phys. 7–8, 613–621.

Popescu, M.N., Dragoescu, I., 1987. Maps of recent vertical crustal

movements in Romania: similarities an differences. J. Geodyn. 8,

123–136.

Radulescu, F., 1988. Seismic models of the crustal structure in Romania.

Rev. Roum. Geol. Geophys. Geogr., Ser. Geophys. 32, 13–17.

Radulescu, F., Diaconescu, M., 1998. Deep seismic data in Romania.

CERGOP “South Carpathians” monograph, vol. 7 (37), pp.

177–192. Warszawa.

Radulescu, D.P., Sandulescu, M., 1973. The plate-tectonics concept and the geological structure of the Carpathians. Tectonophysics 16,

155–161.

Radulescu, D.P., Cornea, I., Sandulescu, M., Constantinescu, P.,

Radulescu, F., Pompilian, A., 1976. Structure de la croute terrestre

en Roumanie, Essai d'interpretation des etudes sismiques

 profondes. Anu. Inst. Geol. Geofiz. L (50), 5–36.

Radulescu, F., Nacu, V., Bitter, M., 1998. Geodetic and geodynamic

data on the recent crustal movements. CERGOP “South

Carpathians” monograph, vol. 7 (37), pp. 231–242. Warszava.

Raileanu, V., 1998. Studiul unor parametri fizici ai litosferei in unele

zone din Romania,. Ph.D. Thesis, Inst.Fizica Atomica, Bucharest,

232 pp.

Raileanu, V., Diaconescu, C., 1998. Seismic signature in Romanian

crust. Tectonophysics 288, 127–136.

Raileanu, V., Diaconescu, C.C., Radulescu, F., 1994. Characteristics of 

Romanian lithosphere from deep seismic reflection profiling.

Tectonophysics 239, 165–185.

Raileanu, V., Bala, A., Hauser, F., Prodehl, C., Fielitz, W., 2005.

Crustal properties from S-wave and gravity data along a seismic

refraction profile in Romania. Tectonophysics 410, 251–272.

Royden, L.H., 1988. Late Cenozoic tectonics of the Pannonian basinsystem. In: Royden, L.H., Horvath, F. (Eds.), The Pannonian

Basin, a study in basin evolution. AAPG Memoir, vol. 45, pp.

27–48.

Sanders, C.A.E., Andriessen, P.A.M., Cloetingh, S.A.P.L., 1999. Life

cycle of the East Carpathian orogen: erosion history of a doubly

vergent critical wedge assessed by fission track thermochronology.

J. Geophys. Res. 104, 29095–29112.

Sandulescu, M., 1984. Geotectonics of Romania. Technical Publishing

House, Bucharest. 336 pp. (in Romanian).

Sandulescu, M., 1988. Cenozoic tectonic history of the Carpathians.

In: Royden, L.H., Horvath, F. (Eds.), The Pannonian Basin, a

Study in Basin Evolution. AAPG Mem., vol. 45, pp. 17 –25.

Seghedi, A., 1998. The Romanian Carpathian foreland. CERGOP

“South Carpathians” monograph, vol. 7 (37), pp. 21–48. Warszawa.Seghedi, I., Szakács, A., 1994. The Upper Pliocene–Pleistocene

effusive and explosive basaltic volcanism from the Per şani

Mountains. Rom. J. Petrol. 76, 101–107 Bucureşti.

Seghedi, I., Balintoni, I., Szakacs, A., 1998. Interplay of tectonics and

neogene post-collisional magmatism in the intracarpathian region.

Lithos 45, 483–497.

Sperner, B., Lorenz, F., Bonjer, K., Hettel, S., Mueller, B., Wenzel, F.,

2001. Slab break-off-abrupt cut or gradual detachment? New

insights from the Vrancea region (SE Carpathians, Romania). Terra

 Nova 13, 172–179.

Sperner, B., Ratschbacher, L., Nemčok, M., 2002. Interplay between

subduction retreat and lateral extrusion: tectonics of the Western

Carpathians. Tectonics 21 (6), 1051. doi:10.1029/TC901028.

Stammler, K., 1994. SeismicHandler  — programmable multichannel

data handler for interactive and automatic processing of seismo-

logical data. Comput. Geosci. 19 (2), 135–140.

Stampfli, G.M., Mosar, J., Marquer, D., Marchant, R., Baudin, T.,

Borel, G., 1998. Subduction and obduction processes in the Swiss

Alps. Tectonophysics 296, 159–204.

Stanica, D., Stanica, M., 1998. 2D modelling of the geoelectric

structure in the area of the deep-focus Vrancea earthquakes.

CERGOP “South Carpathians” monograph, vol. 7 (37), pp.

193–203. Warszawa.

Stefanescu, M., and Working Group,1985. Geological cross sections at 

scale 1 : 200,000 A14-A20 and D1. Inst. Geol. Geofiz., Bucharest.

Szakács, A., Seghedi, I., 1995. The Calimani–Gurghiu–Harghita

volcanic chain, East Carpathians, Romania: volcanologicalfeatures. Acta Vulcanol. 7 (2), 145–155.

Szakács, A., Seghedi, I., 1996. Volcaniclastic sequences around

andesitic stratovolcanoes, East Carpathians, Romania. Rom. J.

Petrol. 77 (Suppl. 1), 1–55.

Tari, G., Dicea, O., Faulkerson, J., Georgiev, G., Popov, S., Stefanescu,

M., Weir, G., 1997. Cimmerian and Alpine Stratigraphy and

Structural Evolution of the Moesian Platform (Romania/Bulgaria).

In: Robinson, A.G. (Ed.), Regional and petroleum geology of the

Black Sea and surrounding region. AAPG Memoir, vol. 68, pp.

63–90.

Tarapoanca, M., Bertotti, G., Matenco, L., Dinu, C., Cloetingh, S.A.P.L.,

2003. Architecture of the Focsani Depression: a 13 km deep basin in

the Carpathians bend zone (Romania). Tectonics 22/6, 1074.

doi:10.1029/2002TC001486.

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

7/27/2019 Hauser Et Al Tectonophysics2007

http://slidepdf.com/reader/full/hauser-et-al-tectonophysics2007 25/25

Visarion, M., 1998. Gravity anomalies on the Romanian territory.

CERGOP “South Carpathians” monograph, vol. 7 (37), pp.

133–138. Warszawa.

Visarion, M., Sandulescu, M., Stanica, D., Veliciu, S., 1988.

Contributions a la connaissance de la structure profonde de la

 platforme Moesienne en Roumanie. Stud. Teh. Econ. - Inst. Geol.,

Ser. D Prospect. Geofiz. 15, 211–222.Wenzel, F., 1997. Strong Earthquakes: a challenge for geosciences and

civil engineering — a new Collaborative Research Center in

Germany. Seismol. Res. Lett. 68, 438–443.

Wenzel, F., Lungu, D., Novak, O. (Eds.), 1998a. Vrancea Earthquakes:

Tectonics, Hazard and Risk Mitigation. Selected papers of the First 

International Workshop on Vrancea Earthquakes, Bucharest,

 November 1–4, 1997. Kluwer Academic Publishers, Dordrecht,

 Netherlands. 374 pp.

Wenzel, F., Achauer, U., Enescu, D., Kissling, E., Russo, R., Mocanu,

V., Mussachio, G., 1998b. The final stage of plate detachment;

International tomographic experiment in Romania aims to a high-

resolution snapshot of this process. EOS 79 (589), 592–594.

Wortel, R., Spakman, W., 2000. Subduction and slab detachment in the

Mediterranean–Carpathian region. Science 290, 1910–1917.

Zelt, C.A., Smith, R.B., 1992. Seismic traveltime inversion for 2-Dcrustal velocity structure. Geophys. J. Int. 108, 16–34.

Zugr ăvescu, D., Polonic, G., Horomnea, M., Dragomir, V., 1998.

Recent vertical crustal movements on the Romanian territory,

major tectonic compartments and their relative dynamics. Rev.

roum. Géophysique 42, 3–14.

Zweigel, P., Ratschbacher, L., Frisch, W., 1998. Kinematics of an

arcuate fold-thrust belt: the southern East Carpathians (Romania).

Tectonophysics 297, 177–207.

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