3D crustal structure in the neotectonic basin of the Gulf of ... - GEIN-NOA

Tectonophysics 400 (2005) 55 – 65 www.elsevier.com/locate/tecto

3D crustal structure in the neotectonic basin of the Gulf of Saronikos (Greece) G. Drakatosa,T, V. Karastathisa, J. Makrisb,1, J. Papouliac,2, G. Stavrakakisa a

National Observatory of Athens, Institute of Geodynamics, P.O. Box 20048, GR 118 10 Athens, Greece b GEOPRO GmbH, St. Annenufer 2, 20457, Hamburg, Germany c Hellenic Center of Marine Research, Institute of Oceanography, P.O. Box 712, 19013 Anavissos, Attiki, Greece Received 18 March 2003; accepted 9 February 2005 Available online 29 March 2005

Abstract An on-/offshore seismic network consisting of 36 three-component stand-alone digital stations was deployed in the area of the Saronikos Gulf, in the vicinity of Athens (Greece), in the fall of 2001. In the present study, from an initial set of more than 1000 micro-earthquakes, 374 were selected and 6666 P- and S-wave arrivals were inverted, based on a 3D linearized tomography algorithm, in order to determine the 3D velocity structure of the region. The resulting 3D velocity distribution, in agreement to the micro-seismicity distribution, reflects the Saronikos structure down to a depth of 12 km. So, the neotectonic basin of the Saronikos Gulf is divided in two parts by a central platform, which implies the existence of a NNE–SSW-trending rupture zone. This zone is probably the offshore extension of a large thrust belt dominating the adjacent onshore areas. Due to their different structure, the two basins are dominated by different velocity values in comparison to the central platform. The western part is characterised by higher seismic activity than the eastern one. Furthermore, the western Saronikos Gulf is divided in a northern and a southern part by a well-defined rupture zone trending E–W. This seems to be the extension of the Corinthiakos Gulf fault zone. At the depth of 17 km, the velocity increases considerably and the crustal thickness is restricted down to 20 km. This dunexpectedT low thickness in the region of Saronikos Gulf seems to be the result of the extensional stress field, which dominates the region, as well as of the emergence of the mantle material along the volcanic arc, which clearly appears at the depth of 12 km. Yet the lack of deep events and, hence, the poor resolution below the depth of 17 km does not support a definite conclusion about the crust–mantle boundary in this region. D 2005 Elsevier B.V. All rights reserved. Keywords: Velocity structure; Tomography; Saronikos Gulf; Greece

T Corresponding author. Tel.: +30 10 3490164; fax: +30 10 3490180. E-mail addresses: [email protected] (G. Drakatos), [email protected] (J. Makris), [email protected] (J. Papoulia). 1 Tel.: +49 40 30399576; fax: +49 40 30399578. 2 Tel.: +30 22910 76370; fax: +30 22910 76323. 0040-1951/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2005.02.004

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G. Drakatos et al. / Tectonophysics 400 (2005) 55–65

1. Introduction The Saronikos Gulf is situated in the northwestern part of the Hellenic volcanic arc, between the Pliocene volcano of Aegina and the Pleistocene volcanoes of Methana and Sousaki (Fig. 1). The area is in generally characterised by low seismicity. However, at its northern and western borders, strong historical as well as recent earthquakes have occurred (Fig. 1), all associated with a roughly N–S extensional tectonic field. In the vicinity of the Saronikos Gulf, the existence of the highly populated Attiki region and particularly the city of Athens gives to the seismicity of the area a major significance, from the social and economic point of view (see also Papadopoulos et al., 2000). During the instrumental observation period of Greece, in the broader area of Athens and particularly in the Saronikos Gulf area, no considerable seismic activity had been recorded prior to the event of September 7, 1999 (Ms=5.9). Therefore, it is not surprising that no systematic micro-seismicity study

has been carried out in the past. After the event of 1999, the need of mapping the active faults in the Attiki offshore area was reconsidered. Specifically, the understanding of the active deformation of the Saronikos Gulf area is a must in order to estimate the seismic potential and seismic hazard of the Attiki region. The acquisition of a local velocity model is also judged essential for understanding the seismogenic processes. The 2D crustal structure of the Saronikos–eastern Corinthiakos basins was investigated in the past, suggesting an intense crustal thinning below the volcanic area of the Saronikos Gulf (Makris et al., 2004a). Following the passive seismic observations of the fall of 2001, a 3D active seismic experiment was also conducted, the results of which are under evaluation (Makris, personal communication). Some work has been also done in the adjacent onshore area of Attiki, after September 7, 1999 earthquake, with 3D passive seismic tomography based on local seismological networks (Drakatos et al., 2002). The existing information

Fig. 1. A simplified tectonic map of the study region is shown. The black lines represent faults, whereas the black broken line represents the thrust belt zone dominating Attiki region. The diamonds represent the historical earthquakes from 450 B.C. to 1900 (M N6.0, after Papadopoulos et al., 2000). The stars show the epicentre distribution from 1901 to 1964 (M N5.0, after Makropoulos et al., 1989; Papazachos and Papazachou, 1997). The black circles show the epicentre distribution from 1965 to 2000 (M N6.0, Monthly Bulletins of Institute of Geodynamics, NOA).


for the velocity structure, provided by large-scale regional passive seismic tomography studies for the whole Greek territory, is not of the required resolution to be effectively used for a detailed study of the active deformation (e.g., Spakman, 1986; Spakman et al., 1988; Drakatos and Drakopoulos, 1991; Ligdas and Main, 1991; Papazachos et al., 1995; Alessandrini et al., 1997; Drakatos et al., 1997). The aim of the present study is to derive the 3D velocity structure of the Saronikos Gulf area, based on the recorded micro-seismic activity during the field experiment of 2001, that involved the installation of eight ocean bottom seismographs (OBS) and 28 stand-alone land stations (Fig. 2). Both marine and land stations were three-component digital stations, used in the above passive experiment the SEDIS III seismic recorder of GeoPro, Hamburg (Makris and Moeller, 1990).

2. Regional geological and tectonic setting The proposed geodynamic models for the construction of the Aegean neotectonic basins consider, in general, a tensional regime with the formation of

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tectonic grabens by normal faulting in the back-arc area of the Hellenic trench (McKenzie, 1978; Dewey and Sengor, 1979; Le Pichon and Angelier, 1979; Brooks et al., 1988). Within this process, the Saronikos Gulf lies along the Hellenic volcanic arc, within the Pliocene volcano of Aegina and the Pleistocene volcanoes of Methana and Sousaki (Papanikolaou et al., 1988). It is divided into a western and an eastern part by a very shallow N–Strending platform, part of which emerges at the islands of Methana, Angistri, Aegina and Salamina (Fig. 1). To the west, the Gulf includes two basins, the WNW– ESE-oriented Epidaurus basin to the south, with a depth greater than 400 m and the E–W-oriented Megara basin to the north, which is relatively shallow (less than 250 m). The WNW–ESE marginal faults of the Epidaurus basin in the southwest have about 350 m of throw and create a rather symmetric tectonic graben with significant volcanic intrusions within its eastern part. The Megara basin is a graben formed by E–W to ENE–WSW marginal faults having a throw of 400–500 m (Papanikolaou et al., 1988). In the eastern Saronikos Gulf, NW–SE faults with relatively smaller throw control the alternation of basins and plateaux of the sea bottom morphology.

Fig. 2. The distribution of the on-/offshore seismic network is shown (Makris et al., 2004b).

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Judging by the fault displacements and sediment distribution, we concluded that the western part is obviously much more active than the eastern one. The presence of recent volcanoes in the central part leads to an even more complicated structure and it separates the active western part from the relatively inactive eastern one. The onshore areas of Attiki (as mapped and published in the 1:50,000 scale geological maps of the Institute of Geology and Mineral Exploration of Greece, IGME, 1989) belong to an autochthonous metamorphic basement of Palaeozoic–Mesozoic age. It consists of marbles and schists (Papanastassiou et al., 2000), with remnants of a tectonic cover identified in the same localities. A major tectonic boundary, trending ENE, separates the mountains of Penteli and Himmitos to the southeast, from the mountains of Parnis and Aegaleo to the northwest (broken line, Fig. 1). Normal faults along the Parnis Mountain are considered as the eastward extension of the Corinthiakos Gulf fault system of almost E–W orientation.

3. Data–method–modelling More than 1000 micro-earthquakes were recorded during the passive seismic experiment of 2001 (Makris et al., 2004b). Four hundred forty five (445) earthquakes (0.3 b ML b 3.8) were determined in the investigated region (37.48N–38.28N, 22.88E–24.28E), recorded at least at five stations. The locations of the hypocenters were obtained using the HYPOINVERSE software (Klein, 1989), which allows the application of spatially varying local velocity models for the hypocenter location. Magnitudes were defined by the coda length of the seismograms, calibrated by using earthquakes also recorded by the Seismograph Network of the National Observatory of Athens. As far as the magnitude estimates are concerned, a mean function from four stations having similar site response was derived, and this was applied to the full temporary network. More details about RMS, ERH, ERZ, magnitude and depth distribution are given in Makris et al., 2004b. The inverse problem of three-dimensional local earthquake tomography is formulated as a linear approximation of a non-linear function (Pavlis and Booker, 1983). Solutions are generally obtained by

linearization with respect to a reference earth model (Aki and Lee, 1976; Nolet, 1978). The solutions obtained and the reliability estimates depend thus on the initial velocity model. Unrealistic initial conditions may result in artefacts of significant amplitude. The used tomography technique was initially introduced by Thurber (1983) and further developed by Eberhart-Phillips (1986, 1990). The method performs an iterative simultaneous inversion for 3D velocity structure and hypocenter parameters using travel time residuals from local earthquakes. The velocity of the medium is parameterised by assigning velocity values at the intersections of a non-uniform, three-dimensional grid. The spacing within the grid is selected in such a way as to have enough ray paths near each grid point so that its velocity may be adequately resolved. The spacing does not need to be uniform throughout the study area. The velocity for a point along a ray path and the velocity partial derivatives are computed by linear interpolation between the surrounding grid points. Thus, the velocity solution will show progressive changes in velocity rather than the sharp discontinuities shown in typical wide-angle reflection models or block-type parameterisations. The ray-tracing technique used is the bART+Pseudo-bendingQ (Approximate Ray Tracing, proposed by Thurber, 1983). As far as the modelling stage is concerned, we set the grid of the nodes at the specified area from 37.408N to 38.208N and 22.808E–24.208E. The nodes are well distributed in seven horizontal layers (slices) at depths of 1, 3.5, 7, 12, 17, 22 and 30 km, respectively. Each layer includes 12 nodes in the E– W direction and 9 nodes in the N–S direction. The distance between two consequent nodes varies from 10 km to 20 km in both directions. In order to constrain better the 3D inversion, initially based on the velocity model (Table 1) proposed by Makris et al. (2004a), we used the Vp and Vs models shown in Table 2. Vp values between two consequent nodes Table 1 Velocity model in Saronikos Gulf (Makris et al., 2004a) Velocity (km/s)

Depth (km)

4.70 5.70 6.80 8.10

0.0 3.0 7.0 17.0

G. Drakatos et al. / Tectonophysics 400 (2005) 55–65 Table 2 Grid geometry Depth (km)

Nodes (E–W)

Nodes (N–S)

Initial P-velocity (kms 1)

Initial S-velocity (kms 1)

1.0 3.5 7.0 12.0 17.0 22.0 30.0

12 12 12 12 12 12 12

9 9 9 9 9 9 9

4.0 5.6 5.9 6.0 7.0 7.5 8.1

2.2 3.1 3.3 3.4 3.9 4.2 4.5

were computed by linear interpolation. From the initial set of earthquakes, 374 events were selected (1.6 bM b 2.6), which gave a good coverage distribution. Some of the events were located outside the modelled area. The inclusion of earthquakes and/or stations outside the modelled area generally improves

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the ray path distribution within it. In order to optimize the best depth determination, we included in the relocation process not only P- but also S-wave arrivals. In total, 6666 arrivals (3402 P-wave and 3264 S-wave readings) were considered. The data set has been inverted five times to get a stable solution. The final epicentre distribution is shown in Fig. 3. The final calculation of the RMS values was improved significantly and most of the events have RMS values varying from 0.1 to 0.3 s. In Fig. 4, the depth distribution of the events is shown. The results of the inversion, in terms of velocity variation, are shown in Fig. 5. Due to the lack of deep events, the two bottom layers (22 and 30 km) were poorly resolved. Therefore, the last one was discarded and not included in the presentation of the final velocity structure. Fig. 6 shows the velocity distribution along selected cross sections.

Fig. 3. Epicentral distribution is shown, after the 3D simultaneous inversion.

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Fig. 4. A 3D view of the depth distribution is shown.

Several parameters, such as the number of hits at each block, the sum of derivatives, etc., can be used to verify the reliability and stability of the solution. The resolution and, hence, the reliability and the stability of the solution strongly depend on the degree of intersection of crossing rays. For the latter, a dense station network of homogeneous distribution over the area under investigation is usually required. In Fig. 7, we present the resolution matrix of the diagonal elements. It is obvious that the solution we obtained is stable for the major part of the velocity model.

4. Discussion and conclusions The micro-seismicity study of the Saronikos Gulf in 2001 (Makris et al., 2004b) produced reliable and accurately located seismic data because of the homogeneous and densely spaced on-/offshore seismic array. A 3D seismic passive tomography based on this data set evaluated for the first time the 3D velocity model for the complete Saronikos area of 6060 km2. As the first evaluation showed (Makris et al., 2004b), the microearthquake activity is associated

with the tectonic regime rather than the volcanic activity of the area. The relocation of the foci after the simultaneous inversion was improved and exhibits now an RMS of 0.1–0.3 s. This, however, had no significant influence on the results already published in 2004, where the main active faults were located north of the Aegina island striking E–W and continuing onshore west Peloponnese and to the east of Aegina striking NNE–SSW (see Fig. 3). From west to east, the micro-seismicity is restricted along a narrow zone trending NNE–SSW (Aegina and Salamina islands) that divides the Saronikos in two regions (Figs. 1 and 3). The shallow micro-seismicity underlines the existence of a shallow platform, part of which emerges at the islands of Methana, Angistri, Aegina and Salamina (Papanikolaou et al., 1988). This feature, as obtained from the velocity structure, seems to continued up to the depth of 12 km (Figs. 5 and 6). This is probably an offshore continuation of the thrust belt, which is dominant in the Attiki region (Fig. 1). This result is also supported by the onshore epicentre distribution of the aftershock sequence of the September 7, 1999 earthquake, which was restricted to the west of the thrust belt (Papanastassiou et al., 2000; Drakatos et al., 2002; Pavlides et al.,


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Fig. 5. The P-wave velocity distribution is shown, after the 3D simultaneous inversion. Depth denotes the layer (slice), where the grid nodes (white crosses) were applied.

2002). This could explain the similarity of the seismotectonic regime of western Saronikos to that of the Corinthiakos Gulf, while the eastern Saronikos has a more simple structure.

The western Saronikos Gulf is divided in two parts, a northern and a southern one, by a well-defined seismic zone of E–W direction (Fig. 3), which is shown between the northeastern part of Peloponnisos

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Location map of cross-sections

A

23.0

23.5

H 38.0

24.0

K

M

A

B

C

D

E

F

38.0

37.5

37.5

G 23.0

23.5

24.0

Vertical section (LAT 37.98)

Vertical section (LON-23.1)

Distance (degrees) B 22.8 23.0 23.2 23.4 23.6 23.8 24.0 24.2

A

G Depth (km)

Depth (km)

B

L

J

-10 -20 3

4

5

6

7

37.4

-20 3

-10 -20 7

37.4

8

3

4

5

6

7

8

Vertical Section (LON 23.67)

-10 -20 7

8

Vp (km/s)

L 37.4 Depth (km)

Depth (km)

K 38.2

Vp (km/s)

Distance (degrees) D 22.8 23.0 23.2 23.4 23.6 23.8 24.0 24.2

6

8

-20

C

5

7

-10


4

6

Distance (degrees) 37.6 37.8 38.0

Vp (km/s)

3

5

Vertical Section (LON 23.33) J Depth (km)

Depth (km)

Distance (degrees) F 22.8 23.0 23.2 23.4 23.6 23.8 24.0 24.2

6

4

Vp (km/s)


5

H 38.2

8

E

4

Distance (degrees) 37.8 38.0

-10

Vp (km/s)

3

37.6

37.6

Distance (degrees) 37.8 38.0

M 38.2

-10 -20 3

4

5

6

7

8

Vp (km/s)

Fig. 6. The P-wave velocity distribution is shown (B), after the 3D simultaneous inversion, along the cross sections shown in (A).

and Aegina Island. These two parts are actually the Megara and Epidaurus basins, respectively, but they are not well defined in terms of velocity variations (Figs. 5 and 6).

An important feature is shown at the depth of 12 km. In general, we could say that higher velocities appear underneath the old volcanic regions of Methana, Sousaki and Salamina. This dring of high


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0.3

0.5

0.8

1.0

0.0

0.3

0.5

0.8

1.0

24 .2

0.0

0.3

0.5

0.8

1.0

0.0

0.3

0.5

0.8

1.0

.8 22

37

.5

23

23

.2

38

23

.4

23

.6

23

.8

24

24

.2

0.0

24

Depth 1.0 km

.8 22

37

.5

23

23 .2

38

23 .4

23

.6

23

.8

24

.2

Depth 3.5 km

.8 22

37

.5

23

23 .2

38

23 .4

23

.6

23 .8

24

Depth 7 km

.8

0.0

0.3

0.5

0.8

0.0

0.3

0.5

0.8

.8 22 .8

37

.5

23

23 .2

38

23 .4

23 .6

23 .8

24

24 .2

22

Depth 22 km

37

.5

23

23

.2

38

23

.4

23

.6

23

.8

24

24

.2

22

Depth 17 km

37

.5

23

23 .2

38

23

.4

23

.6

23 .8

24

24

.2

Depth 12 km

Fig. 7. The diagonal elements of the Resolution matrix are shown, corresponding to the layers where the grid was applied.

velocitiesT indicates upraised mantle material due to the palaeo-volcanic activity. This result is in agreement with the very low resistivity values determined at depths below 10 km during a magnetotelluric exploration of Sousaki geothermal area (Tzanis and Lagios, 1993). The presence of this mantle material

and the extensional stress field are probably the main causes of the thin crustal thickness in the Saronikos region observed by Makris et al. (2004a). At the depth of 17 km, the velocity increases considerably. As it is shown in the cross sections of Fig. 6, the thickness of the crust is restricted down

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to 20 km. Despite the slight influence of the initial velocity model in the results of the tomography, the low thickness in the region of Saronikos Gulf is also supported by active seismic observations. This could be the result of the extensional stress field, which dominates in the region at present, as well as of the intrusion of mantle material in the volcanic arc.

Acknowledgements The authors are grateful to Dr. F. Haslinger and Dr. C. Chiarabba for their constructive comments and suggestions. This study was partially supported by the Earthquake Planning and Protection Organisation and the General Secretariat for Research and Technology of Greece.

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