The strain rate field in the eastern Mediterranean region, estimated by repeated GPS measurements

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Tectonophysics 294 (1998) 237–252

The strain rate field in the eastern Mediterranean region, estimated by repeated GPS measurements Hans-Gert Kahle a,Ł , Christian Straub a , Robert Reilinger b , Simon McClusky b , Robert King b , Kenneth Hurst c , George Veis d , Kim Kastens e , Paul Cross f a

Geodesy and Geodynamics Laboratory, Institute fu¨r Geoda¨sie und Photogrammetrie, ETH Zu¨rich, CH-8093 Zu¨rich, Switzerland b Department of Earth, Atmosphere, and Planetary Sciences, M.I.T., Boston, USA c Jet Propulsion Laboratory, Pasadena, USA d Higher Geodesy Laboratory, NTU, Athens, Greece e Lamont Doherty Earth Observatory, Columbia University, New York, USA f University College London, Department of Geomatic Engineering, London WC1E6BT, UK Received 20 May 1997; accepted 26 November 1997

Abstract We use the combined GPS velocity field of the eastern Mediterranean for the period 1988 to 1996 to determine crustal deformation strain rates in a region comprising the Hellenic arc, the Aegean Sea, and western Anatolia. We interpret the velocity field and determine the strain rate tensor by the spatial derivatives of the collocated motion vectors. The region following the line Marmara Sea, North Aegean Trough, northern central Greece, and the central Ionian islands is associated with strong right-lateral shear motion, with maximum shear strain rates of 180 nano-strain=a (180 ð 10 9 =a). In the central Aegean Sea, N–S-oriented extensional processes prevail, reaching 100 nano-strain=a. The southern Aegean is characterized by relatively small strain rates. Maximum extensional components of the strain rate tensor, reaching 150 nano-strain=a in a N–S direction, are found in central Greece. The Hellenic arc is associated with moderate arc-parallel extension and strong compression perpendicular to it. Projections of the strain rates parallel to the major fault zones reveal that the northern Aegean is governed by the westward continuation of the North Anatolian Fault Zone which is associated with strong dextral shearing (maximum 220 nano-strain=a), accompanied by numerous large earthquakes in this century.  1998 Elsevier Science B.V. All rights reserved. Keywords: GPS; strain; crustal deformation; eastern Mediterranean

1. Introduction The Alpine–Mediterranean region is a zone of strong seismotectonic deformation resulting from the interaction between the Eurasian and African=AraŁ Corresponding

geod.ethz.ch

author. Fax: C41 1 633 1066; E-mail: kahle@

bian plates (Fig. 1). Argus et al. (1989) derived rotation rates of Africa relative to Eurasia increasing from Maroc to Tunisia, from 4 mm=a to 7 mm=a. This increase in rate is due to the increased distance from the Euler pole of rotation of Africa relative to Eurasia at 21ºN 21ºW, with Africa rotating anticlockwise at an angular velocity of ! D 0:1º=Ma. For the eastern Mediterranean this model predicts

0040-1951/98/$19.00  1998 Elsevier Science B.V. All rights reserved. PII: S 0 0 4 0 - 1 9 5 1 ( 9 8 ) 0 0 1 0 2 - 4

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Fig. 1. Tectonic pattern of the Mediterranean region and surrounding areas. Motion vectors of the African plate (single arrows) based on rotations around the Euler pole after Argus et al. (1989), relative to Eurasia. Motion along Dead Sea Rift, Red Sea, and Gulf of Aden (double arrows) after Cochran (1983), LaBrecque and Zitellini (1985), relative to Nubia and Somalia, respectively. Double arrow at west Hellenic arc shows maximum motion of the Aegean microplate, relative to the European reference system. AEMP D Aegean microplate; AP D Anatolian microplate; BS D Black Sea; BSZ D Bitlis suture zone; CA D Calabrian arc; CS D Caspian Sea; DSR D Dead Sea Rift; EAF D East Anatolian Fault; GA D Gulf of Aquaba; KTJ D Karliova Triple Junction; LS D Libyan Sea; MS D Marmara Sea; NAFZ D North Anatolian Fault Zone; RP D Russian platform; SP D Sinai Peninsula, WHA D West Hellenic arc.

relative northward motion of the African plate at a rate of 7 mm=a in the Ionian Sea and 8 mm=a in the Libyan Sea (Fig. 1). At the northeastern end of the African plate, the Arabian plate is separating from Africa (Nubia) in a NNE direction across the Red Sea spreading center (Drake and Girdler, 1964), and from Somalia in a NE direction across the Gulf of Aden spreading center (Laughton et al., 1970). Based on the spreading rates of these young oceanic basins, derived from magnetic seafloor-spreading anomalies, the rate of northward motion of the Arabian plate is on the order of 10 mm=a along the Gulf of Aquaba (GA) and the

Dead Sea Rift (DSR) (Cochran, 1983; LaBrecque and Zitellini, 1985), and 15 mm=a in the southeastern part of the Arabian plate around Oman (Cochran, 1981; Jestin et al., 1994). The NUVEL 1A model (DeMets et al., 1994) predicts a north motion of about 22 mm=a and a west motion of 9 mm=a of Arabia relative to Eurasia just south of the Bitlis suture zone (BSZ). These higher rates in the NUVEL 1A model can be explained if the rotation of Africa relative to Eurasia is added to the rotation of Arabia relative to Africa (Somalia). The low rate of shortening between Africa=Arabia and Eurasia contrasts with the Aegean microplate,

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Fig. 2. Seismicity of the Mediterranean between 1979 and 1995 with Ms ½4, after NEIC (1994). This map shows that the Aegean Sea region is the most active area in the eastern Mediterranean. A D Apennines; Al D Albania; Ap D Apulia; AS D Aegean Sea; C D island of Crete; Cyp D island of Cyprus; D D Dinarides; E D island of Evia; Ep D Epirus; Io D Ionian islands; GC D Gulf of Corinthos; K D Kozani; MS D Marmara Sea; NAFZ D North Anatolian Fault Zone; NA D North Aegean Trough; R D island of Rhodes; S D Sicily; WHA D West Hellenic arc.

which is moving up to 35 mm=a to the southwest with respect to Eurasia. The motion of the Aegean is accompanied by a seismic belt encircling the Hellenic–Aegean domain (Fig. 2). From the Dinarides it follows the Hellenic arc=trench system, the North Aegean Trough (NAT) and the North Anatolian Fault Zone (NAFZ) in the Marmara Sea region (MS), NW Anatolia. The early division of McKenzie (1972) of the eastern Mediterranean into several microplates drew attention to the dominant role of the Anatolian– Aegean microplates. A detailed tectonic stress field was compiled by Rebai et al. (1992), using seismic focal plane mechanisms, in-situ stress measurements, and neotectonic data. Motion in the Alpine=Mediterranean region was initially measured by Satellite Laser Ranging (SLR) using permanent

and mobile tracking equipment (Smith et al., 1994; Noomen et al., 1995; Robbins et al., 1995). Since 1988, Global Positioning System (GPS) networks have been remeasured several times to assess the fine structure of the kinematic field. From west to east these networks are: the Calabrian arc (CA) (Zerbini et al., 1994; Kaniuth et al., 1995), the West Hellenic arc (WHA) (Kahle et al., 1993), central Greece (CG) (Denys et al., 1995), the Aegean Sea (AS) (Gilbert et al., 1994; Kastens et al., 1998), the Marmara Sea (MS) (Straub and Kahle, 1995; Straub, 1996) and the Anatolian microplate (AP) (Oral, 1994; Reilinger et al., 1997). In this paper we present the combined velocity field derived from the WHA, CG, AS, MS and the western part of the AP GPS networks, calculate regional crustal strain rates and compare the results with seismic data.

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2. The geodetic networks 2.1. West Hellenic Arc (WHA) network The WHA network comprises southeast Apulia (Italy), the Ionian islands, the mainland of western Greece, the Peloponnesus and the island of Crete. A few sites have also been observed in the AS network in order to connect the WHA and AS networks. The characteristics of the WHA surveys and the rates of motion vectors are summarized in Kahle et al. (1995); the trajectories of crustal deformation between 1989 and 1994 are compiled in Kahle et al. (1996). The kinematic field of western Greece is characterized by negligible motion north of the Kephalonia Fault Zone (KFZ), relative to southeast Italy, a right-lateral transitional shear zone to the south of the KFZ, and a homogeneous SW-oriented motion of southwest Greece reaching rates of 35 mm=a. 2.2. The central Greece (CG) network The CG network was established in 1988 to study the crustal motion in the region around the Gulf of Corinth and the island of Evia (Denys et al., 1995). A few sites of the CG network were also occupied in the WHA surveys to tie both networks together. The CG network was occupied in 1989, 1991, and 1993. Subsets of it were also observed in 1995 and 1996. Since the 15 June 1995 Egion earthquake in the Gulf of Corinth, M D 6:1, has caused local co-seismic displacements we have included only results of the first three surveys in this paper (Mu¨ller, 1996). Calculations of interseismic velocities including observations in 1996, with the co-seismic displacements removed, are in progress (Clarke et al., 1998). 2.3. Aegean Sea (AS) network The AS network consists of 30 stations in Greece, on islands, and in the periphery of the Aegean Sea. Several sites in the western part of the network were also occupied in the WHA surveys. The details of the data, analysis, and the derived velocity field are presented and discussed in Kastens et al. (1998). The southern half of the network was first measured in the fall of 1988. The northern half was first measured

in the fall of 1989. A third survey covering both the northern and southern sub-networks was carried out in 1992, providing velocities for all stations based on at least two sets of measurements. The stations in the southern part of the Aegean are essentially moving together at about 30–35 mm=a to the southwest relative to the Eurasian plate. 2.4. Marmara Sea (MS) network The MS network covers the northwestern part of the boundary between the Eurasia and Anatolian plates (AP). The continuation of the single fault zone of the NAFZ is splaying into a complex fault system (Barka and Kadinsky-Cade, 1988). The results of the MS network for the periods 1990, 1992 and 1994 are summarized in Straub and Kahle (1995). The average westward motion of the Anatolian plate in the Marmara Sea region is 22 mm=a relative to the Eurasian plate. West of 27.5ºE the motion turns to ENE–WSW. Most of the deformation occurs along the northern fault zone, which passes through the Marmara Sea. This corresponds well with the distribution of major historical earthquakes in this area (Ambraseys and Finkel, 1991). A detailed description of the data analysis, the crustal motion derived from it, and the models constructed, is given in Straub (1996). Comparisons of the GPS results, including data of the 1996 survey, with geologic estimates of crustal motion are made in Straub et al. (1997). 2.5. Anatolian microplate (AP) network The results of the AP network are summarized and discussed by Reilinger et al. (1997). The velocities for the sites located south of the Bitlis suture zone (BSZ) in eastern Anatolia indicate NNW- to NW-oriented motion relative to Eurasia. Stations located north of the BSZ in eastern and central Turkey show velocities similar to those to the south of it, suggesting that the motion of Arabia is being transferred directly to Anatolia. The stations bridging the Caucasus mountains indicate a relative shortening on the order of 10 mm=a across the Caucasus. Sites north of the NAFZ and north of the Greater Caucasus show insignificant motion in the Europe-fixed reference frame.

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West of the Karliova Triple Junction (KTJ), the motion of Anatolia is characterized by coherent counterclockwise rotation about a pole located near the northern Sinai Peninsula (SP). Internal deformation of the central part of the Anatolian plate is less than 2 mm=a. The Anatolian microplate is decoupled from Eurasia by the right-lateral strike-slip NAFZ and from Arabia by the left-lateral East Anatolian Fault (EAF). Significant deviations from coherent plate motion occur in western Turkey where N–S extension at a rate of about 15 mm=a is superimposed on the counterclockwise rotation. 2.6. SLR results in the eastern Mediterranean The SLR motion solutions give a good quantitative pattern of the regional deformation taking place in the eastern Mediterranean (Noomen et al., 1993, 1995; Smith et al., 1994; Robbins et al., 1995). The northward motion of a site south of the Bitlis suture zone (BSZ in Fig. 1) reflects the movement of the Arabian plate. The westward motion of the Anatolian block is seen in the solution for the motion of stations in central Anatolia. Three stations located on the leading edge of the overriding Aegean microplate show SW-oriented motion with rates of nearly 35 mm=a, relative to a Europe-fixed reference system. The solutions for the sites in the Italian part of the network suggest that the motion here is small relative to the Europe-fixed sites.

3. The kinematic and strain rate fields The velocity field shown in Fig. 3 was generated by combining rates estimated separately by the different analysis groups for the AS, WHA, MS, and AP networks (Reilinger et al., 1995). Each solution was performed nominally in an Eurasian-fixed reference frame, defined by including in the analysis of each survey data from tracking stations in northern Europe. We then further aligned the frames by minimizing the rates with respect to Europe of four stations in the Aegean, determined by Kahle et al. (1996) and Kastens et al. (1998), and three stations near the Black Sea, determined by Reilinger et al. (1997). Comparison of pairs of closely located stations in the AP and AS networks indicates that at

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least these two are aligned within their current uncertainties. We included the MS network with no additional translation and found that the Eurasiafixed velocities determined by Straub (1996) are compatible with those of Reilinger et al. (1997) within their uncertainties. Four stations at SLR pads, measured during the surveys indicated in Table 1, showed an average deviation of 2.2 mm=a in the N–S component and 3.3 mm=a in the E–W component. A rigorous treatment of the data from all of the surveys, from the west Hellenic arc to the Caucasus region and including measurements acquired in 1996 and 1997, is in preparation (McClusky et al., 1998). The strain rate field is based on the horizontal displacement rates shown in Fig. 3. For the African=Arabian side of the eastern Mediterranean, where no GPS data are available, we used rates calculated from rigid rotations (Africa: Euler rotation pole at 21ºN 21ºW; angular velocity ! D 0:1º=Ma (Argus et al., 1989); Arabia: 33ºN 24ºE; ! D 0:42º=Ma (Jestin et al., 1994)). The rates were interpolated using the method of 2dimensional least-squares collocation. Collocation is a general method of least squares adjustment including parameter estimation, filtering and prediction. The displacements, the signal-to-noise ratio of the displacements, and a covariance function are used as input. The covariance function correlates the interpolated displacement values with the actual measurements. Details are given in Wirth (1990), Danuser et al. (1993), and Kahle et al. (1995). We apply this method for the quasi-spherical case (Straub, 1996). We use a covariance function which is reciprocal to the square of the distance between the stations. The correlation length of the covariance function is set to 250 km (signal-to-noise ratio of 12.3) because it corresponds nearly with the average distances between the stations. The consequence of using much smaller correlation lengths is a strongly varying strain field with an exaggeration of effects of the noise, whereas a much larger correlation length than the average distance of the stations tends to smooth out signals, thereby loosing valuable information. Our heuristic approach in defining the correlation length is described in Straub (1996). The strain tensor components žilh at a station i are given by:

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Fig. 3. Combined velocity field in the eastern Mediterranean relative to a Europe-fixed reference frame, based on GPS results from Mu¨ller (1995, 1996), Straub (1996), Reilinger et al. (1997) and Kastens et al. (1998).

žilh D

0

1

arccos.#i j / @#i j 2 2 Ð r0l2 Ð rsphere Ð q Ð B C 2 @ x hj n B C 1 # X i j B 2 C Ð kjC B¦l   2 B C 2 jD1 @ r0l2 C rsphere Ð arccos2 .#i j / A where ¦l2 D a-priori sigma of the signal, l D component; r0l D correlation length; rsphere D radius of the sphere; #i j D cosines of the centric angle between the points i and j; .@#i j /=.@ x hj / D derivative at position j in direction of the component h must be expanded separately for the latitude and longitude; k j D correlation vector. The displacements represent the motion vectors averaged over the time interval between the observa-

tion surveys. As such, the ž-values are strain rates. The principal values and axes of deformation are calculated by solving the eigenvalue problem (Fig. 4). To account for the neo- and seismotectonic processes forming the major fault systems, the strain rate tensor components were projected onto the known major faults in the eastern Mediterranean. The projection of strain rates onto the major fault systems is made using the method described by Straub (1996) (Figs. 5 and 6). The scales are given in µ-strain=a; they equal a relative motion of 1 mm=a over a distance of 1 km, normal to the fault. If the width of the deformation zone is multiplied with the strain rates the total relative motion is obtained. The fault system applied was assembled using data from Le Pichon and Angelier (1979); Finetti (1982); Taymaz et al. (1991a,b); Barka (1992); Reuther et al.

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Table 1 Comparison of the motions at SLR sites and unweighted deviations Name Name

Latitude (N)

Longitude (E)

KARI

39º440

20º390

XRIS

36º470

21º520

DION

38º040

23º550

ROUM

35º240

24º410

VN (mm=a)

VE (mm=a)

1N (mm=a)

1E (mm=a)

Network

0.7 3.6 0.5 3.8 21.9 18.8 21.0 22.9 21.7 23.0 21.7 22.6 26.9 33.1 26.0

3.5 6.3 2.3 2.3 21.7 25.2 22.0 20.6 19.8 19.7 19.8 16.8 29.6 22.6 15.7

0.6 3.7 0.6 3.7 0.8 2.4 0.2 1.8 0.6 0.8 0.6 0.4 1.8 4.4 2.7 2.2

1.05 3.85 0.15 4.75 0.7 2.8 0.4 1.8 0.8 0.7 0.8 2.2 7.0 0.0 6.9 3.3

WHA Aegean CG SLR WHA Aegean CG SLR WHA Aegean CG SLR WHA Aegean SLR

Average deviation

WHA D West Hellenic arc; CG D central Greece; SLR D satellite laser ranging after Noomen et al., 1995; Lat D Latitude; Lon D Longitude. VN D north component, VE D east component, 1N , 1E D unweighted deviations from other campaigns indicated.

(1993); Royden (1993); Rebai et al. (1993); Jackson (1994); Caputo (1995) and Pavlides et al. (1995).

4. Discussion 4.1. Comparison with stress maps The stress data for the Aegean Sea and western Anatolia compiled by Mu¨ller et al. (1992) and Zoback (1992) come mainly from neotectonics and seismology (Mercier et al., 1987; McKenzie, 1978; Anderson and Jackson, 1987). These data show dominant N– S extension in western Anatolia and in the Aegean. Data east of 30ºE display more scatter in compressive horizontal stress .SHmax / orientations, and indicate a change in deformational style to a more strike-slip type. The Hellenic arc is associated with SHmax orientation, consistent with SSW convergence direction between the Aegean Sea region and Africa. Despite the scatter of the data and the uneven distribution of sampling sites, the overall pattern is in general agreement with the results derived from the GPS based strain rate tensor field (Fig. 4) showing strong compression perpendicular to the Hellenic arc. This result is based on the GPS data combined with the assumed counterclockwise rotation for the African plate.

Relatively small normal strain rates are seen in central Anatolia. As shown by Reilinger et al. (1997), these are also the areas where residuals of the modeled rigid block rotations for the eastern Mediterranean are smallest. In addition, our results indicate a relatively strain-free region in the central southern Aegean (AS in Fig. 4), between the volcanic (37º to 38ºN) and the non-volcanic Hellenic arc (35º to 36ºN). High extensional strain rates are confined to the northern Aegean Sea, as well as in western Anatolia, consistent with the stress map of Mu¨ller et al. (1992). Maps of in-situ stress measurements compiled by Rebai et al. (1992) show mainly N–S extensional horizontal stress .SHmin / in central Greece, which is also seen in the GPS results. 4.2. Comparison with fault plane solutions Seismicity and fault-plane solutions for southern Italy show active deformation which varies between N–S shortening and NE–SW extension on normal faults along the Apennine Chain (Anderson and Jackson, 1987). The Dinaric coast region is deforming on strike-slip and thrust faults. A belt of NE– SW-shortening continues into northwestern Greece with numerous transcurrent fault systems (Mantovani et al., 1992). NE–SW-oriented compressional

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Fig. 4. Principal values and axes of strain rates calculated from the velocity field shown in Fig. 3. White arrows indicate extension; black arrows show compression. Points without arrows indicate low strain rates. AS D Aegean Sea.

strain is also seen in the GPS-based strain rates in northwestern Greece (Fig. 4). There is a lack of GPS stations in this region and to the north of it for one to go into further detail. The southwestern margin of Greece is dominated mainly by the subduction of the African plate along the Hellenic trench. The most important fault zone taking up this motion is the right-lateral KFZ (Kahle et al., 1996). It separates GPS stations to the north of it, which show negligible motion relative to Europe, from sites to the south of it, which are rapidly moving to the SW. The dextral nature of the KFZ is clearly witnessed by the shear strain rates shown in Fig. 6. This area shows intense seismicity as well (Fig. 2): Devastating earthquakes with magnitudes of M > 7, associated with dextral type of FPS (Fig. 7), occurred on the central Ionian islands in 1945, 1953

(Stiros et al., 1994), 1972 (Anderson and Jackson, 1987), and in 1983 (Kiratzi and Langston, 1991). Taymaz et al. (1991a) showed that, based on fault plane solutions, the northern and central Aegean area is governed by distributed right lateral strike-slip faulting trending NE to ENE. Normal faulting created several prominent basins at the eastern end of the NAT. The GPS strain rate analysis reveals that these features are dominated by right-lateral sense of shear rate components (Fig. 6) accompanied by extension, scattering between NNE–SSW to NNW– SSE directions (Fig. 5). The dextral sense of motion along the NAFZ, clearly visible in the GPS results (Fig. 6), is also witnessed in the FPS (Fig. 7). Sinistral shear-strain is dominant in the area beginning east of the island of Crete (C), running along the island of Rhodes (R) and into southwest Turkey,

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Fig. 5. Scaled normal strain rates projected onto major fault structures derived from space geodetic observations. Extension governs western Anatolia, the Aegean Sea and central Greece. The WHA prevails by strong compression. AP D Anatolian microplate; CA D Calabrian arc; KFZ D Kephalonia Fault Zone; CG D central Greece; WHA D West Hellenic arc.

identifying the Strabo and Pliny troughs (ST–P), which may act as transcurrent faults, separating the moving Aegean plate from the Herodotus (HB)– Antalya (AB) and Levant Basins (LB) (see Fig. 6). The southern Aegean region and central Anatolia appear to be shear-free regions. Northwest Anatolia is dominated by both extensional and compressional rates. The SW-trending portions of the NAFZ are associated with compressional rates, whereas the NW-oriented segments show mostly extensional activity (Fig. 5). This can be related to the restraining and releasing stress behavior along the NAFZ which changes its strike on several well-defined fault segments (Barka, 1992). The western part of Anatolia is characterized by N– S-oriented extensional strain rates, coincident with W–E-trending graben structures, which are also as-

sociated with large-size earthquakes of normal faulting type in the recent past. Examples are the 1971 Burdur earthquake sequence (Taymaz and Price, 1992) or the 1995 Dinar earthquake (Eyidogan and Barka, 1996). The latter one was also accompanied by a left-lateral strike-slip component. Small sinistral strike-slip strain rates are also seen in the GPS derived strain field (Fig. 6). We surmise that this part of southwestern Anatolia forms a triangle which is pulled out to the west and at the same time possibly rotated counter-clockwise (compare also Barka and Reilinger, 1997). 4.3. Comparison with seismic moment tensors The assessment of seismic strain requires an estimate of the moment tensor for each earthquake,

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Fig. 6. Scaled shear strain rates projected onto major fault structures derived from space geodetic observations. The NAFZ and its westward continuation along the NAT, passing central Greece and the KFZ are the dominant shearing structures of the eastern Mediterranean. KFZ D Kephalonia Fault Zone; WHA D West Hellenic arc; HB D Herodotus Basin; LB D Levant Basin; AB D Antalya Basin; AS D Aegean Sea; CY D Cyclades; R D Rhodes; C D Crete; NAT D North Aegean Trough; NAFZ D North Anatolian Fault Zone; ST–P D Strabo–Pliny trench.

constructed from the strike, dip and rake combined with the seismic moment. The seismic strain tensor is commonly estimated by summing the seismic moment tensors in a volume of known dimensions (Kostrov, 1974). 4.3.1. Hellenic Arc Components of strain rate tensors based on seismic moment tensor summation have recently been re-calculated by Papazachos and Kiratzi (1996) and converted to velocities using 63 seismogenic sources in the Aegean and the surrounding area. The deformation along the coastal regions of Albania and northwestern Greece down to the KFZ, is taken up by compression in a direction perpendicular to the

coast line, at a rate of 4 mm=a in a direction N49ºE. The right-lateral strike-slip KFZ demarcates the beginning of the Hellenic subduction zone. Nearly all seismic events that have occurred south of it, along the Hellenic arc, have dominantly dip-slip thrust mechanisms. The increasing importance of compressional strain accumulation along the Hellenic arc is also clearly seen in the GPS results (Fig. 5). 4.3.2. Central Greece Ambraseys and Jackson (1990) have examined the seismicity of central Greece between 1890 and 1988. They found that the distribution of the 100-year interval is comparable to the last 25-year interval. The most striking feature of the seismicity of the longer

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Fig. 7. Compilation of fault plane solutions (FPS) of earthquakes in the eastern Mediterranean region (1908–1996) published by McKenzie (1972, 1978); Anderson and Jackson (1987); Eyidogan (1988); Jackson and McKenzie (1988); Taymaz et al. (1991a,b); Papazachos and Kiratzi (1992); Kiratzi (1993); Zanchi and Angelier (1993); Jackson (1994); Kiratzi and Papazachos (1995); Pavlides et al. (1995), and Harvard CMT determinations (www-address: ‘http:==tempo.harvard.edu=CMT.html’).

period is the intense activity in the Gulf of Corinth and between the island of Evia and the Greek mainland. Our GPS results indicate that this region is characterized by large extensional strain rates. Maximum extension occurs near the region of the graben-type plains in northern central Greece, which in 1954 was the site of several strong normal faulting earthquakes (Papastamatiou and Mouyaris, 1986), in 1957, and in 1980 (Papazachos et al., 1993). The most recent example of normal faulting events in northern central Greece is the Kozani earthquake .M D 6:6/ of 1995 (Pavlides et al., 1995) which ruptured along a normal fault system, the T-axis oriented in a NW–SE direction. Average rates based on seismic moment release for northern central Greece were estimated at 10 mm=a in the N6ºW

direction, a rate that is considerably smaller than the one obtained from GPS data. Maximum strain rates, found by Ambraseys and Jackson (1990), are 70 nano-strain=a as compared to 150 nano-strain=a we obtained from the GPS observations. Whether this discrepancy is due to aseismic strain release or to an apparent seismic strain deficit because of the short time interval considered for the moment summation, remains an open question. 4.3.3. Aegean Sea The new inversion of seismic events of Papazachos and Kiratzi (1996) confirms that the Aegean is characterized by extensional deformation. Its direction shows an counterclockwise rotation in the north (Marmara–NAT), from east to west. The extension

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has a NNE direction in northwestern Anatolia and NNW direction in central and northern Greece, as well as along the southern volcanic arc, and its extension into southwestern Anatolia. Jackson et al. (1992, 1994) determined the horizontal velocity field from the spatial distribution of seismic moment tensors of earthquakes in the Aegean region. They found E–W right-lateral shear in northwestern Anatolia related to motion on the NAFZ, becoming distributed as it enters the northern Aegean Sea. Further south, towards the volcanic chain, they identified N–S extension. In the eastern Aegean the seismicity does not account for the geodetic strain rates observed. It was, therefore, postulated by Jackson et al. (1994) that a seismic strain deficit or aseismic slip may exist in the eastern Aegean, as well as in central Greece. Our data confirm this observation for central and northern Greece in particular. For seismic hazard assessment it is important to assess which of the two reasons, strain deficit or aseismic slip, is responsible for the discrepancy clearly revealed. On the other hand the southeastern Aegean seems to be almost strain-free, based on the GPS data. 4.4. Comments on driving mechanisms The spatial acceleration of motion from the Anatolian to the Aegean region and across the northern and southern Aegean cannot be solely related to the convergence between the Arabian and Eurasian plates and to the tectonic escape of the Anatolian plate to the west (Kasapoglu and Tokso¨z, 1983). In fact, the push of the Arabian microplate, in terms of rates, is about half of the rates or less as compared to the rapid motion of the Aegean region. To account for the seismicity pattern and for the pronounced spatial acceleration of motion of the Aegean microplate additional driving mechanisms such as slab roll-back and gravitational collapse must be invoked, in addition to tectonic escape (Meijer, 1995; Meijer and Wortel, 1996). The boundaries of the deforming and moving Aegean region have been the subject of debate. While it is generally agreed that the northern Aegean trough and the northern margin of the Marmara Sea form the boundary in the northeast, considerable uncertainty exists as to where the boundary passes through north-

ern Greece, how it connects with the West Hellenic subduction zone and what the relative displacement rates are. Reuther et al. (1993) have argued that central and northern Greece neotectonics are dominated by a transform pull-apart mechanism. Older faults were reactivated as E–W-trending tensional and oblique strike slip faults. These E–W-trending faults connect the transtensional dextral NAT with the active dextral KFZ. The pull-apart mechanism additionally supports the overall tensional processes in central Greece. Late Pliocene to Holocene extension has also been reported for northern Greece where several basins, strike-slip faults and detachment zones form an extensional system which is kinematically linked to the dextral-slip displacements along the NAT (Dinter and Royden, 1993). Stations on the mainland of Africa will shed light on the amount of shortening, as for example between Crete and Africa. Motions obtained from a DORIS station in Niger, Africa, is 6 mm=a relative to Wettzell, Germany, and 5 mm=a relative to Madrid, Spain (Kolenkiewicz et al., 1996). Results of surveys conducted recently in the eastern Mediterranean, including Egypt, will eventually constrain the kinematic boundary conditions for models of driving mechanisms.

5. Conclusions In this paper the current kinematic field of crustal motion, based on repeated GPS observations, is analyzed in terms of strain rates and compared with seismological data. From these results a first assessment of the geophysical implications is given in light of the critical boundaries between the rapidly moving Aegean–Anatolian microplates and Eurasia. The following major conclusions are reached. (1) The principal axes and the values of compressional and extensional strain rates (Fig. 4) calculated from the GPS velocity field (Fig. 3), yield insights into the major features subjected to this stress pattern: in the western part of Anatolia as well as in central and southern Greece, N–S-oriented extensional strain rates dominate, accompanied by normal faulting earthquake mechanisms. We propose that they are due to transtensional types of active processes in the Aegean graben system, possibly generated by

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back-arc spreading and roll-back behind the Hellenic arc. (2) The calculation of extensional, compressional and shear strain components projected onto the major fault systems show that three main types of strain distribution dominate: extension in the Aegean region and central northern Greece, compression perpendicular to the Hellenic arc (Fig. 5), and shear along the line NAFZ, NAT and KFZ (Fig. 6). They are manifestations of the westward motion of Anatolia as well as the rapidly moving Aegean microplate. Our strain rate results corroborate previous studies. The stress trajectories determined from marine geophysical data (Lybe´ris, 1984) agree well with the combination of extensional and shear components of strain rate components found in the GPS data. (3) Papazachos and Kiratzi’s (1996) new results of inversion of seismic events for the central Ionian islands are now in better agreement with the GPSbased rates. In an earlier paper (Papazachos et al., 1992) their results indicated almost 50% lower rates for the KFZ region. The reason for the higher values they find now might be due to the longer time span (1928–1992), the integration of more recent seismic events, and change of some parameters, such as the thickness of the seismogenic layer. In the southern Ionian islands compression occurs in a N34ºE direction across the Hellenic arc. The direction of shortening is almost constant from the Ionian islands to the island of Rhodes. The GPS data show that the maximum shortening trends N27ºE, which is in fairly good agreement with the seismic results. (4) In light of the critical boundaries of the moving Aegean–Anatolian microplates we concluded that maximum dextral shear strain rates occur along the NAFZ–NAT–KFZ zone, clearly delineating the northern boundary of the Aegean extensional domain (Fig. 6). Maximum GPS-deduced sinistral shear strain rates are observed along the Strabo-Pliny trench system (ST–P), clearly defining the southeastern boundary of the Aegean microplate. The ST–P seems to act as a transform fault system, at least in terms of GPS data. Its role as possible backstop for the subducting African plate remains unclear. (5) One important application of geodetically determined strain rates and seismically estimated moment tensors is seismic risk assessment. Jackson and McKenzie (1988) discussed the relationship between

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seismic moment tensors and rates of deformation. Based on velocity measurements for only 6 SLR sites, apparent deficits in seismically released strain were already localized in the eastern Mediterranean (Jackson et al., 1994). From the GPS analysis we conclude that this holds true for central and northern Greece in particular. The southeastern Aegean region is almost strain-free, based on the GPS data. In summary, we conclude that the strain rate field in the eastern Mediterranean is further constrained by the GPS data presented. The results are considered as a first preliminary step towards a better understanding of the complex present-day dynamics of the eastern Mediterranean–Alpine region. An important part of the regional deformation processes, the height components of recent plate tectonic movements and deformations in particular, are as yet widely unknown. Another lack of information is the absence of stations on the African plate. New GPS surveys were recently conducted in the eastern Mediterranean area including the Balkan region and northern Africa. In addition, continuous GPS networks (CGPS) are being installed, which are designed to monitor the ongoing deformation in the frame of earthquake hazard assessment. One example is the CGPS network across the Kephalonia Fault Zone (Peter et al., 1998). These new data will also help to connect all the networks and fill remaining gaps in the kinematic pattern of the eastern Mediterranean.

Acknowledgements This study benefited greatly from discussions with A. Geiger and M.V. Mu¨ller, ETH, on GPS analysis, collocation and strain field calculation. We thank B. Ambrosius, University of Delft, and T. Dixon, University of Miami, for critically reading the manuscript and helpful suggestions for its improvement. ETH’s contribution was funded by ETH grant 41-2647.5. MIT’s participation was supported in part by NSF Grant EAR-9304554.

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