First European Conference on Earthquake Engineering and Seismology (a joint event of the 13th ECEE & 30th General Assembly of the ESC)

Geneva, Switzerland, 3–8 September 2006 Paper Number: 1062

MICROSEISMICITY AND FAULTING GEOMETRY IN CENTRAL GREECE

Vassilis KARAKOSTAS1, Christos KARAMANOS2, Eleftheria PAPADIMITRIOU3, Ioannis KASSARAS4 and Constantinos MAKROPOULOS5

SUMMARY

During November 2004 – June 2005 a digital seismological network was deployed in the eastern part of central Greek mainland, in an area seismically active in the 20th century, particularly in the decade of fifties. The strongest earthquake (M=7.0) occurred in 1954, while the last strong one in 1980 (M=6.5). In total 18 Reftek digital loggers (both 72A–07 and R–130) were installed assembled with fifteen guralp CMG40T broadband and three Le_1Hz seismometers. The average spacing between stations was of the order of 20 km to ensure earthquake depth accuracy. Local earthquakes with P– and S– arrivals at four or more stations were located using HYPOINVERSE computer program. Data analysis using the Double Difference technique did not change considerably the spatial distribution of the earthquake foci. The best–recorded earthquakes were used to define a reasonable crustal structure. Lateral variations of the crustal model taken into account, calculating time delays for each station. By this way, earthquakes inside the network or close to its boundaries were located with high accuracy in both the epicenter and focal depth. In addition, focal mechanisms of earthquakes with proper azimuthal coverage were computed. Seismicity covers most of the area and is distributed mainly in clusters along active structures. A magnitude M=3.9 earthquake, was the largest local one recorded by this network (5 December 2004, 17:58 UTC), close to the focal area of the 1980 strong earthquake. Several cross sections striking normal to the trend of the clusters of the epicenters reveal the geometry of the active structures as well as the width of the seismogenic layer. Most of the focal mechanisms exhibit normal faulting and were used along with the microseismicity foci in the cross sections for the definition of the properties of the faults that activated during the experiment.

1. INTRODUCTION

The study area belongs to the extensional back–arc Aegean region, among the most active ones in the Alpine–Himalayan belt, with its most prominent tectonic feature the subduction of the eastern Mediterranean oceanic lithosphere under the Aegean microplate [Papazachos and Comninakis, 1970] along the Hellenic Arc (Fig. 1). Seismicity is very high throughout the arc, which is dominated by thrust faulting with a NE–SW direction of the axis of maximum compression. A belt of thrust faulting runs along the southwestern coasts of Yugoslavia and continues south along the coastal regions of Albania and northwestern Greece, resulted from continental collision between Outer Hellenides and the Adriatic microplate. The direction of the maximum compression axis is almost normal to the direction of the Adriatico–Ionian geological Zone. Between continental

1 Department of Geophysics, School of Geology, Aristotle University of Thessaloniki, GR54124 Thessaloniki, Greece Email : vkarak@geo.auth.gr 2 Department of Geophysics, School of Geology, Aristotle University of Thessaloniki, GR54124 Thessaloniki, Greece Email: chrikar@geo.auth.gr 3 Department of Geophysics, School of Geology, Aristotle University of Thessaloniki, GR54124 Thessaloniki, Greece Email : ritsa@geo.auth.gr 4 Department of Geophysics & Geothermics, Faculty of Geology, School of Science, University of Athens, GR15784 Athens, Greece Email: kassaras@geol.uoa.gr 5 Department of Geophysics & Geothermics, Faculty of Geology, School of Science, University of Athens, GR15784 Athens, Greece Email: kmacrop@geol.uoa.gr

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collision to the north and oceanic subduction to the south, the dextral strike–slip Cephalonia Transform Fault (CTF) is observed [Scordilis et al., 1985], in agreement with the known relative motion of the Aegean and eastern Mediterranean. McKenzie [1978] showed that the northward motion of the Arabian plate pushes the smaller Anatolian plate westwards along the North Anatolian fault, continuing along the North Aegean Trough (NAT) region, which is the boundary between the Eurasian plate and Aegean microplate. Right–lateral strike–slip motion associated with the North Anatolian Fault (NAF) appears to become more distributed in the north Aegean Sea, which is characterized by a combination of right–lateral shear and extension. This motion is transferred into the Aegean but in a southwesterly direction. Our study area constitutes part of the large–scale process zone that encompasses the western end of the North Anatolian Fault, which is characterized by the superposition of two deformation fields (propagating NAF and back–arc extension) now interacting in the Aegean and corresponds to a rapid southwestwards tapering of the NAF slip rate [Flerit et al., 2004].

Figure 1: Main seismotectonic properties of the Aegean and surrounding regions. The study area is indicated by

the rectangle. The study area exhibits high seismic activity with several strong earthquakes of M>6.0 during the instrumental era. The last seismic sequence took plane in 1980 with a main shock of M6.5 with its largest aftershock of M6.1 just 24 minutes later. The most severe seismic excitation occurred during 1954–1957 with five events of M>6.0 that produced extensive damage in several urban areas and substantial loss of life. The consecutive occurrence of these events was interpreted by the stress transferred through the associated almost E–W trending normal faults that are positioned along strike and bound the southern part of the study area [Papadimitriou and Karakostas, 2003]. The purpose of the present study is the exploitation of data collected by a temporary network of eighteen (18) digital seismological stations that installed and operated for seven months. Their location as much accurately as possible and fault plane solutions determination, contribute to investigate the local seismic activity and fault kinematics, not only complementing previous investigations [Hatzfeld et al., 1999] but giving more insight in the characteristics of the active structures. Even though microseismicity is not completely representative of the properties of the faults associated with strong earthquake occurrence, it may provide

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valuable information in the cases when accurate and adequate data are not available to constrain in detail the seismotectonic properties of a target area.

2. SEISMOTECTONIC SETTING The study area is characterized by extension on sub–parallel normal faults that strike E–W [Papazachos et al., 2001]. The NW–SE trending normal ones in the Late Miocene – Early Pliocene [Caputo and Pavlides, 1993] control the local morphology and bound two parallel basins and probably the coastline (Fig. 2). A dextral strike slip motion during Miocene was identified suggesting possible initial connection of the western extension of the North Anatolian fault with the southern Thessalia fault zone, which maybe constitutes an active boundary and for this reason larger earthquakes occur there than in its northern margin [Mountrakis et al., 1993]. The contemporary N–S extension is revealed by the fault plane solutions of both strong [Papazachos et al., 1998] and small earthquakes [Hatzfeld et al., 1999]. In the offshore area, north of the North Aegean Trough, a normal fault striking at 100o was imaged in three multichannel profiles [Laigle et al., 2000]. The southern part of the study area comprises the Sperchios basin with two active faults known to be associated with events of M>6.5 during the last five centuries, namely the Lamia fault (M6.8, 1545) and the Skarfeia fault (M6.6, 1740) [Papazachos et al., 2001]. The northern Evoikos basin, with a thinned, stretched continental crust of only 20 km thickness, was developed by the separation of the island of Evia from the Greek mainland, by transtension and stretching of the crust, forming the north Evia and the Sperchios valleys [Makris et al., 2001]. This region contains a set of large coastal faults that have Holocene scarps [Roberts and Jackson, 1991], and which are known to have moved in earthquakes [Papazachos and Papazachou, 2002].

Figure 2: Major faults along with their code names, which are known to be associated with the occurrence of strong (M>6.0) earthquakes [Papazachos et al., 2001 with modifications]. Big yellow stars denote events with M>5.0 since 1911 and smaller white stars events of M>4.0 since 1981 [Papazachos et al., 2005]. The available fault plane solutions of earthquakes occurred during the instrumental era, are also shown as lower–hemisphere

equal area projections.

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The map of Figure 2 shows the known active faults along with instrumental seismicity. It is evident that recent smaller magnitude seismicity is sparse, mainly concentrated offshore at the south western termination of the North Aegean and along the northern part of Paghasitikos gulf, where the 1980 seismic sequence took place. Since this data sample looks inadequate to provide information on the geometry of known active faults in the area, and because we are interested moreover in defining the geometry of active structures that are not known to be associated with strong earthquakes, the installation of a portable seismological network was decided for microseismicity recording. Microseismicity analysis will be then performed in order to reveal the properties of the regional active structures as well as the translation of the strike–slip motion from the Aegean into the Greek mainland. The activation of these faults constitutes a major threat since they are located onshore and close to urban areas.

3. TEMPORARY NETWORK OPERATION

A local seismic network was installed and operated in the study area during November 2004 – May 2005. The network consisted of eighteen 3–component stations equipped with fifteen broad–band and three short period sensors, all connected to GPS antenna for timing. The first choice was to place the network in the area shown in the left part of figure 3 (November 2004–March 2005) in order to define as much as possible more accurately the spatial distribution of microseismicity associated with faults that were activated in the 20th century. During the second phase of the experiment (March 2005–June 2005), four stations that were placed in the northern part were moved to the south in order to closing up the network at the northern margin, as well as at the south part, which exhibited higher seismic activity (Fig. 3, right part). The stations operated in continuous mode with a sampling rate of 125 samples/sec, allowing the detection and location of 323 events both on land and offshore. In addition, the recordings of the permanent seismological stations shown in figure 3, which are operated by the Geophysics Department of Aristotle University of Thessaloniki, were embodied to the local network data. The recordings at the permanent stations are digitized at a rate of 100 samples/sec.

Figure 3: Locations of the regional network (rectangles) and temporary network (stars) stations. The left part of the figure depicts the locations during the first period, while the right one the locations during the second period

of the experiment.

4. DATA AND PROCEDURES

We analyzed the P and S wave arrival times and P wave first motions from the local network for hypocenter location and fault plane solutions determination. During the field program explosions were recorded and were used for a local velocity model constraint. The explosions were recorded up to a distance of 45 km and a one–layer velocity model was defined with P–wave velocity equal to 5.73 km·sec-1. The events with the larger

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number of recordings (up to 180 km) and best location by the use of the local model (46 events in total) were used to construct travel time curves, which were used to obtain a local velocity model (v1=5.73 km·sec-1, d1=4.07 km, v2=6.11 km·sec-1, d2=8.83 km, v3=6.76 km·sec-1, d3=8.27 km, v4=6.89 km·sec-1, d4=∞) and a best vp/vs ratio equal to 1.76. In order to obtain an integrated velocity model and since our recordings did not exceed the distance of 180 km, we added a layer that starts at the depth of 32 km, and thus now d4=10.89 km, with a velocity v5=7.85 km·sec-1. This was derived from the recordings of the permanent network for the study area and is in excellent agreement with the regional velocity model (7.9 km·sec-1 at a depth of 31 km) proposed by Panagiotopoulos [1984]. After the location, the events were grouped in five spatial clusters and in each subregion the residuals at each station were calculated in order to take into account lateral heterogeneities. This set of station delays was used in addition with the local velocity model as input to HYPOINVERSE [Klein, 1989] for improved hypocenters of 868 events. The histograms in Figure 4 show that the final locations had root–mean–square residual (rms) less than 0.30 s and for the majority of the data less than 0.20 s (left part of Figure 4). The epicenter error was for most of the locations less than 3 km and for the majority of them less than 2 km (central part of Figure 4). Most calculated depths exhibited errors less than 3 km (right side of Figure 4).

Figure 4: Histograms of the root–mean square of the time errors, RMS (in seconds) (left part of the figure), the

horizontal, ERH (in km) (central part of the figure), and vertical, ERZ (in km) (right part of the figure), uncertainties of the hypocentral determination.

Because the station spacing was adequate, we were able to compute the individual microearthquake focal mechanisms for the study area. All the single–event, lower–hemisphere focal mechanisms were determined, when more than 6 arrivals were available, using the FPFIT computer program [Reasenberg and Oppenheimer, 1985]. In order to avoid misleading determination and to ensure the nodal plane constraints, the station azimuthal coverage as well as the distribution of the compressions and dilatations, were taken into account. The above criteria deterred the focal mechanisms determination of all the recorded events, but the ones located inside the network. A total of 33 mechanisms were computed with a gap smaller than 180o sampling a minimum of three quadrants. Most of the reliable solutions in the study area exhibit either strike–slip or normal faulting. 4.1 Spatial distribution of the seismicity The epicentral distribution of the located events is shown in the map in the left side of Fig. 5, where spatially distinct clusters can be observed with different epicentral alignment, evidencing the existence of a diverse

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faulting pattern. An E–W trending cluster is located along the northern coast of Paghasitikos gulf, probably associated with the fault that activated during the 1980 seismic sequence. In the offshore area to the east, a second cluster is observed, where the North Aegean Trough is terminated. A more dense cluster is aligned along the northeastern coast of Evia Island, trending NW–SE, evidencing an active structure there which is not known to be associated with the occurrence of a strong earthquake. The most dense cluster is located in the south central part of the study area, exhibiting an epicentral line up from NE to SW. North and close to Lamia fault an E–W epicentral alignment is present, probably associated with the aforementioned fault. The seismicity spatial distribution will be discussed in a later section together with fault plane solutions and regional seismotectonic properties.

Figure 5: Spatial distribution of the located seismicity (left part) and fault plane solutions (right part).

4.2 Focal mechanisms of the earthquakes The most reliable fault plane solutions with first–motion polarities for 37 events are shown in the right part of Fig. 5. Nearly all of the focal mechanisms show transtensional or normal faulting. Nearby located events exhibit similarities in their fault plane solutions, and in particular, the events that seem to be associated with the same active structure. Mechanisms related to small–magnitude earthquakes are not usually associated with the finite slip on major faults and their slip vectors could be associated with the motion of randomly distributed preexisting faults, but P or T axes should be generally oriented as the main stress directions. The information then derived from this analysis can complement the information coming from stronger events occurrence and in some cases reveal unknown or temporarily inactive faults. Near the northern coast of Paghasitikos gulf, pure normal to oblique normal faulting is shown, generally in agreement with the strike of the fault plane solution of the 1980 mainshock (strike=81o, dip=40o, rake=–90o, [Papazachos et al., 1983]). The fault plane solutions offshore in the Aegean and along the eastern coasts of Evia Island exhibit mainly strike slip and secondary normal faulting, along either NE–SW or NW–SE striking faults. The Northern Aegean area is dominated by dextral NE–SW trending strike–slip faults in several parallel branches, because of the continuation of the North Anatolian Fault westwards [Papazachos et al., 1998]. The coexistence of left–lateral and right–lateral strike–slip faulting or possibly conjugate faulting has also been observed in the western part of the Aegean Sea, where such activation was observed during the 2001 Skyros Island sequence [Karakostas et al., 2003]. Two fault plane solutions in the Maliakos gulf are in full agreement with the normal movement assigned for the coastal faults in that area [Papazachos et al., 2001]. The tight south cluster provides evidence for a structure that does not accord with what is known in that area and it will be discussed in more detail in the following sections. 4.3 Geometry of the active structures In order to investigate further the geometry of the active structures we divided the region into six rectangles comprising the seismicity clusters, shown in Fig. 6 (upper part) along with seismicity and the locations of cross sections. The sections, which are perpendicular to the mean trend of the seismicity and are deliberately chosen to

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be narrow to reveal details, were performed with front projections of the focal mechanisms that were determined in each rectangle. The thickness of the seismogenic layer comes to 15 km for almost all the cases with the vast majority of the hypocenters being between 10 and 15 km. The dips of the nodal planes are in good agreement with the mean dips of the clusters defined by the hypocenters alone.

Figure 6: Epicenter map and projections of the hypocenters onto vertical cross sections within the six boxes shown in the upper part of the figure. The mechanisms that were plotted in figure 4 are shown as equal area

projections of the front hemisphere.

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Along the northern coast of Paghasitikos gulf (section A1A2 in Fig. 6) the microseismicity seems to be associated with the Aghialos fault, which has an ENE–WSW strike and dips to south and which is associated with the 1980 main shock (M6.5). At the bottom of the seismogenic layer a ‘flattening’ of the hypocentral distribution is observed, an indication for termination in a detachment near the brittle–ductile transition. The section B1BB2 is located at the termination of the North Aegean Trough, along which dextral strike–slip faulting on NE–SW striking faults takes place. The section reveals two parallel active structures striking NE–SW and dipping to the SE with rather steep angles (more than 45 ). This geometry and the oblique normal faulting exhibited by the three fault plane solutions, imply that these structures are part of the dominant dextral strike slip faulting but with the N–S extension being more prevalent.

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The elongated cluster that is extended along the northeastern coasts of Evia Island (section C1C2 in Fig. 6) signifies a NW–SE trending active structure, dipping to the NE at a high angle. The fault plane solutions agree with this geometry, giving more insight for sinistral strike–slip faulting at this location, which is in agreement with the faulting type of the 2001 Skyros main shock [Karakostas et al., 2003], located more southeasterly. The fault plane solution of this mainshock, however, according to the CMT solution determined by Harvard (strike=148o, dip=71o, slip=–1o), exhibits more steep fault plane and pure sinistral strike slip motion. In the present case the oblique faulting, which is in better agreement with the mechanism of the 1967 M6.6 earthquake (strike=313o, dip=43o, slip=–56o) [Taymaz et al., 1991], provides evidence for transtensional motion in this part of our study area. The small cluster in the straight north of Evia Island (section D1D2 in Fig. 6) denotes a south dipping normal faulting structure that is parallel to the coastlines. Although regional seismicity (Fig. 2) or microseismicity from previous local experiments [Hatzfeld et al., 1999] were not intense in this location, our data may outline an active structure there, for which further investigation is needed by the use of larger data sample. The cross section E1E2 comprises the hypocenters of an E–W trending cluster and manifests the activity associated with the Lamia normal fault that dips to the north. The southernmost tight cluster denotes a south dipping active structure, with a NE–SW strike, of oblique normal faulting with a considerable dextral strike–slip component (section F1F2 in Fig. 6). The located microseismicity exhibited its association with faults that are known to have generated strong earthquakes in the past (Aghialos and Lamia faults), brings out the activity on structures for which no strong event has been assigned, revealed the transtensional faulting character in the study area and the continuation of the dextral strike–slip faulting farther to the Greek mainland, and the current dormancy of well known active faults (Skarfeia fault).

5. DISCUSSION AND CONCLUSIONS Resolving the geometry of active structures in a seismically active region, and in particular in the cases where these structures are on shore and close to urban areas, has been a long–standing goal since the derived information can directly be used in seismic hazard assessment. The experience shows that an active fault and its slip direction can be determined by combining accurately located earthquakes and fault plane solutions. For this reason microseismicity and focal mechanism solutions are used to determine the current tectonic activity in the area of southern Thessalia and northern Evoikos gulf (central Greece). Microseismicity during November 2004 and June 2005 mainly occurred on a few reasonably well–defined fault zones, substantially agreed with historical and instrumental regional seismicity. It suggests present activity on parallel E–W trending normal faults (north Paghasitikos gulf and Lamia fault), and on oblique normal faulting structures that are either the continuation of the dextral strike–slip motion that dominates in the North Aegean, or the manifestation of the conjugate sinistral strike–slip faults. The fault plane solutions of the microearthquakes that determined in the present study are consistent with the mechanisms of stronger earthquakes [Papazachos et al., 1998; Taymaz et al., 1991] as well as with smaller ones [Hatzfeld et al., 1999] and therefore can be considered as representing the dominant seismotectonic pattern. Neither strike–slip faulting delineated by the microseismicity has been mapped previously in the area of the most southern cluster. It is difficult to interpret this fault in terms of a regional pattern since its extent, and thus its relation to other faults, is unknown. A possible explanation is that right lateral shear stress produced along the North Aegean Trough and the parallel dextral strike–slip faults may penetrate as far as this location of the Greek mainland, producing this faulting type. The right–lateral strike–slip movements superimposed on the normal faulting along this area may result from this process.

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Faults that had accommodated intense activity in the past few centuries as the historical information reveals, as well as during the instrumental period exhibit quiescence even in the small magnitude events occurrence. This is not an artifact since the network geometry permitted the detection of such activity, but it is rather due to their current placidity. Changing patterns of large faults activity in central Greece was discussed by Jackson [1999] and include faults becoming permanently or temporarily inactive, spatial migration of activity between different faults or fault sets, and faults apparently going through episodes of intense activity separated by long quiescent periods. During the 17th and 18th century, a west–to–east migration of seismic activity (M>6.0) took place three times along the northern margin of Thessalia basin [Papazachos et al., 1994] and the southern margin of the basin, where only two events of M>6.2 occurred in the last five centuries, was struck in a series of strong events during the second half of the 20th century. The episodic occurrence of strong events in the Thessalia area was interpreted in terms of stress transfer through neighboring faults of fault segments [Papadimitriou and Karakostas, 2003]. Identification of changing activity patterns along with fault geometry and kinematics has significant implications in the seismic hazard assessment of a study area. Acknowledgements. The GMT system [Wessel and Smith, 1998] was used to plot the figures. This study was supported by the research project between Greece and USA EPAN–M.4.3.6.1 funded by the General Secretariat of Research and Technology of Greece. Geophysics Department contribution 667.

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