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Reflection–refraction seismics in the Gulf of Corinth: hints at deep structure and control of the deep marine basin
Tectonophysics 391 (2004) 97 – 108 www.elsevier.com/locate/tecto
Reflection–refraction seismics in the Gulf of Corinth: hints at deep structure and control of the deep marine basin Christophe Cle´menta, Maria Sachpazib, Philippe Charvisc, David Graindorgec, Mireille Laiglea, Alfred Hirna,*, Giorgios Zafiropoulosd a
Laboratoire de Sismologie Expe´rimentale, De´partement de Sismologie, UMR 7580 CNRS, Institut de Physique du Globe de Paris, 4 Place Jussieu, Tour 14, B89, F-75252 Paris cedex 05, France b Geodynamics Laboratory, National Observatory of Athens, Lofos Nymfon, Athens, Greece c UMR Ge´osciences Azur-IRD, P.O. Box 3, 06235 Villefranche-sur-Mer, France d Hellenic Petroleum, Maroussi, Athens, Greece Accepted 3 June 2004 Available online 11 September 2004
Abstract The Gulf of Corinth is a natural laboratory for the study of seismicity and crustal deformation during continental extension. Seismic profiling along its axis provides a 24-fold normal-incidence seismic reflection profile and wide-angle reflection– refraction profiles recorded by sea-bottom seismometers (OBS) and land seismometers. At wide-angle incidence, the land receivers document the Moho at 40-km depth under the western end of the Gulf north of Aigion, rising to 32-km depth under the northern coast in the east of the Gulf. Both refraction and normal-incidence reflection sections image the basement under the deep marine basin that has formed by recent extension. The depth to the base of the sedimentary basin beneath the Gulf, constrained by both methods, is no more than 2.7 km, with ~1 km of water underlain by no more than ~1.7 km of sediment, less than what was expected from past modeling of uplift of the south coast in the East of the Gulf. Unlike the flat sea-bottom, the basement and sedimentary interfaces show topography along this axial line. Several deeps are identified as depocenters, which suggest that this axial line is not a strike line to the basin. It appears instead to be controlled by several faults, oblique to the S608E overall trend of the south coast of the Gulf, their more easterly strikes being consistent with the instantaneous direction of extension measured by earthquake slip vectors and by GPS. D 2004 Elsevier B.V. All rights reserved. Keywords: Gulf of Corinth; Aegean region; Seismic refraction; Reflection; Crustal structure; Rift; Basin; Extension
1. Introduction * Corresponding author. Tel.: +33 1 44273914; fax: +33 1 44273894. E-mail address: [email protected] (A. Hirn). 0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2004.07.010
The Gulf of Corinth in central Greece (Fig. 1) is located in the back-arc region above the Hellenic subduction zone, in the post-Alpine extensional
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Fig. 1. The Gulf of Corinth. Location map of profiles A, B, and C shot by N/O Nadir and recorded by OBSs, located at numbered triangles and at land stations NAF and KLI (squares). The Moho reflection midpoint lines at these stations from shots on profile A are sketched as the two thick gray lines (offshore north of Aigion and onshore north of the eastern end of the Gulf). Only the axial profile A and its crooked turn in the West, W, could be recorded as a multifold seismic reflection profile because of safety reasons concerning deployment of the ship’s seismic streamer. Geology is adapted from Armijo et al. (1996), with earthquakes from Taymaz et al. (1991), Baker et al. (1997), and Bernard et al. (1997). Inset shows the study region (rectangle) in relation to its surroundings, including the North Anatolian Fault Zone (NAFZ).
domain of the Aegean (e.g., Gautier et al., 1993). This region has also been affected by the North Anatolian strike-slip fault (Fig. 1) since the Pliocene (Armijo et al., 1996). Strong seismic activity characterizes this region, which is extending in a southward direction, at a rate that exceeds 10 mm/year across the Gulf (Clarke et al., 1998). The south coast of the western Gulf of Corinth has experienced destructive earthquakes, such as the Helike event of 373 B.C. and the Aigion event of M s ~7 in 1861. From an early microearthquake study, Melis et al. (1989) proposed that the southern bounding faults of the asymmetrical graben forming the Gulf of Patras and the adjacent western Gulf of Corinth were listric, flattening northward into a midcrustal decollement. In the western Gulf of Corinth, microearthquakes recorded during temporary seismograph deployments (Rigo et al., 1996) appear distributed across the Gulf above a cutoff depth that has been considered as a plane dipping northward at a low angle of 108, from a depth of 8 km under Aigion. This plane has been interpreted as an underlying detachment representing the brittle–ductile transition, the base of the seismogenic layer, or as a fault in the brittle domain that is slipping through small earthquakes (Rietbrock et al., 1996). Focal mechanisms constrained
by waveform modeling of earthquakes of larger magnitude, 5.7 to 6.2 (Taymaz et al., 1991; Baker et al., 1997) have low-angle north-dipping nodal planes. A study of the 1995 M s 6.2 event, which caused damage at Aigion on the south coast but had its focus at depth beneath the northern margin of the Gulf, suggested a 338 northward dip for its fault plane, although aftershocks form a cluster elongated with a smaller apparent dip (Bernard et al., 1997). In the east of the Gulf, the 1981 sequence of three large M sN6 earthquakes provides the main data source. These are typical high-angle normal-fault events. Such a steep fault and a thick elastic plate have been used to model the uplift of the Plio–Quaternary marine terraces on the south coast of the eastern Gulf (e.g., Jackson et al., 1982; Keraudren and Sorel, 1987, Armijo et al., 1996; Westaway, 1996, 2002). The recent geological evolution of the Gulf, of which this earthquake activity is a present instantaneous expression, has accumulated finite extension, forming a morphological rift structure floored by a deep and flat marine basin elongated toward N1208E (i.e., S608E), cutting across the strike of Alpine structures. The mode of seismic energy release, on low-angle or high-angle normal faults, as well as the
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mechanism of extension of the continental crust on the geological time scale are debated for this structure that may be the world’s fastest-extending graben in continental crust. The structure of its marine infill and of the underlying basement are obvious markers of its style of extension. Previous attempts at deciphering this record of deformation from the structure of this basin were based on single-channel shallow-penetration seismics (e.g., Brooks and Ferentinos, 1984; Higgs, 1988). In order to acquire data to help constrain the evolution of this Gulf, we undertook seismic imaging using modern methods in order to provide observations of its structure, to introduce into this discussion a different point of view, complementary to onshore geological observations and their modeling, and studies of seismicity.
2. Methods and survey During the Franco–Greek SEISGRECE seismic survey in January 1997, the oceanographic vessel N/O Le Nadir of IFREMER shot a 2900 cubic inch array of 14 air guns operated in single-bubble mode (Avedik et al., 1996) every 50 m in the part of the Gulf to the east of Aigion (Fig. 1; Sachpazi et al., 1998). This mode of shooting an air gun source provides with maximum efficiency a signal that is peaked in the rather lowfrequency 12–20 Hz band but has a duration short enough for acceptable resolution for normal-incidence seismic reflection. This frequency band is high enough to correspond to that typically observed for reflections returned by the lower crust even in surveys using sources that also generate higher frequencies, like the more usual tuned arrays of air guns. This lowfrequency band also provides signal propagation to the large offset wide-angle reflections, despite attenuation. The shots in the Gulf of Corinth were recorded by six ocean-bottom seismometers (OBSs), operated on the bottom of the Gulf, as well as by seismometers at station sites on land (Fig. 1). The 96-channel, 2400m long seismic streamer could be deployed from the vessel for part of the survey. This allowed recording of the normal-incidence reflection profile A in Fig. 1 with 24-fold coverage, suitable for advanced processing with pre-stack depth migration. However, this was only possible for a 50-km distance along the axis of the Gulf because operation was impeded by the worst
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winter storm in the last 15 years. Acquisition could be maintained for a tie line, W in Fig. 1, at the western end of the Gulf, until crosspoint E (Fig. 1), but with a low signal-to-noise ratio and inadequate geometry as it was acquired along an arc of a circle with the ship turning. Although the strength of this seismic source has allowed us to image at normal incidence the whole crust elsewhere in the Aegean (Sachpazi et al., 1997), here the noise of extremely strong sideways reflection of water waves off the coast of the narrow Gulf dominates at times when reflections that could be returned from deeper than 15 km are expected.
3. Deep crustal elements from wide-angle reflections Land stations, offset at either end of the shot-line along the axis of the Gulf of Corinth (Fig. 1), recorded clear seismic waves out to the maximum recording distance of 105 km, showing the efficiency of signal generation and reflection. These waves arrive much later, as much as 5–6 s later at 50-km offset, than the first arrival Pg-wave refracted in the basement. With a high-velocity move-out, they thus cannot be interpreted as anything other than the PmP-phase: wideangle reflections from over 30-km depth, that is, from the Moho (Fig. 2). This is a rewarding case of a vertical reflection seismic source allowing one to obtain wide-angle reflections even in single coverage, without stacking for the Moho, and paves the way for the regional mapping of Moho topography using this shooting strategy with numerous receivers. Unfortunately, several other seismograph stations, which were probably slightly less protected from the worst winter storm during acquisition, were dominated by background noise. There was no previous seismic measurement of the Moho depth in this region, only inferences from gravity and regional tomography (Makris and Stobbe, 1984; Tsokas and Hansen, 1997; Papazachos and Nolet, 1997). From a teleseismic tomography experiment resulting in the usual horizontally layered medium with blocks of varying velocity, Tibe´ri et al. (2000) discussed lateral variations in average velocity of their upper blocks in terms of a change in the proportion of crustal and mantle material, that is, a variation in Moho depth. The artificial source wide-
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angle reflection observations in Fig. 2 indeed reveal a strong difference of 8 km in crustal thickness between the western tip of the Gulf west of Aigion and its northeast coast (Sachpazi et al., 1998; Cle´ment, 2000). The water depth and structure of the underlying sedimentary basin, determined from our vertical reflection profiling, are taken into account in the calculation of this variation in Moho, as reported in Fig. 2 and its corresponding caption. This is the simplest possible structural model and has an average crustal velocity typical for Europe, but other solutions are also possible. Propagation paths toward each of the recording stations all sample the crust under the profile of shots but do not sample the same Moho segment at midpoint. The condition that the Moho is continuous and straight, consistent with the average velocity used, could thus be relaxed. With this alternate assumption, Moho depths remain ~8 km larger in the west than the east, but the trade-off between velocity and depth results in a model space ranging from slower crustal velocities giving an overall shallower and convex-upward Moho to faster velocities giving an overall deeper and concaveupward Moho. The fact that Moho reflections are observed at as short an offset as 40 km could lead to prefer slower crustal velocities to shorten the critical distance. However, to achieve the amplitudes observed at small distances, other causes, such as Moho topography focusing the reflection, are necessary, so these observations alone do not provide a tighter constraint on crustal velocity. With data from only two locations, this thickness difference could simply be regarded as inherited from the Alpine phase of deformation, because a crustal root is still present, as attested by the Bouguer gravity low that cuts north–south at right angle to the Gulf (Tsokas and Hansen, 1997) to the west of the offshore area north of Aigion where the reflections to station NAF have their midpoint. After removing the effect of the Ionian slab, Tibe´ri et al. (2001) inverted the residual
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Bouguer anomaly in terms of a variation of crustal thickness that may as well explain the teleseismic tomography results of Tibe´ri et al. (2000). They find a similar difference in Moho depth between the two regions we have sampled with the wide-angle reflections (Sachpazi et al., 1998; Cle´ment, 2000). The Moho topography inverted from gravity shows a Moho with alternating highs and lows, including a low (i.e., thick crust) under the western Gulf and a high (i.e., thin crust) northeast of the Gulf of Corinth, which agree with the wide-angle reflection observations, plus another high south of the eastern end of the Gulf. The crust in this study region thus appears quite thick, somewhat unexpectedly given the common belief associating thin crust with rifts and other extensional domains. Apart from the Moho reflections, the wide-angle intracrustal response is very weak, which cannot be due to a too weak seismic source, as the response of the underlying Moho is strong. That the intracrustal reflectivity would be so weak may also be unexpected in this extensional region, given that strong lower crustal reflectivity has been observed in numerous vertical reflection and wide-angle profiles and attributed to extension (e.g., Allmendinger et al., 1987). This interpretation has also been proposed in western Europe, where this fabric has been acquired in the past, presumably as a result of post-Variscan crustal thinning (e.g., Bois and ECORS Scientific Party, 1990). The profiles that revealed this intracrustal reflectivity showed furthermore that unlike the layer-interface reflectivity well known in sediments, it may not be due to specular reflections on few continuous interfaces between lithological units. It has instead been considered that lower crustal reflectivity is caused by the heterogeneity of the medium and that this heterogeneity is related not only to the nature of the rocks but also to finite deformation or its rate and/or any contribution of magmatism (e.g., Warner, 1990). The strong active extension in this study region has evidently not yet
Fig. 2. Record sections from land stations NAF and KLI (Fig. 1) at variable offset, with 8 km/s reduction velocity, of air gun shots from line A along the axis of the Gulf. Note on both very late, clear waves, interpreted as Moho reflections, and weak upper crustal refracted arrival on NAF, from 3 to 5 s reduced time between 40 and 70 km. West is left and East is right. Superimposed in white are travel-time curves computed for a model including the water and sediment layers under the shot line, as derived from OBSs and vertical reflection profiling, which reveals Moho to be 8 km deeper in the west than in the east. In this model, an average 6.25 km/s velocity was used for the crust under the sediments, and Moho was assumed to dip at 10% (~5.78) from east to west. The Moho thus dips from the 32-km depth sampled at midpoints along the gray line in Fig. 1 under the north coast of the Gulf and east of the shot line (for shots to KLI), to the 40-km depth sampled at midpoints shown as the gray line in Fig. 1 offshore north of Aigion (for shots to NAF).
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achieved the crustal image seen in terranes inferred to have undergone extension long ago. This may indicate that the important factors governing whether extension is imprinted into the lower-crustal fabric are the finite amount of strain and the temperature conditions for flow.
4. Refraction seismic constraints on basement depth and sedimentary fill The velocity layering of the upper crust is clearly illustrated by the conspicuous branches of the traveltime curve in the wavefield recorded by OBS 4 located in the eastern widest and deepest part of the basin, in the hanging wall of the Xylocastro fault (Fig. 3a). The velocity–depth structure is derived from the split-spread profile recorded by this OBS for shots on line B, as follows. Branch 1 in yellow is a refractiondiving wave that constrains a velocity–depth gradient layer just beneath the sea bottom of sediment with a low seismic velocity V 1 of average 2 km/s. Branch 2 in blue is a precritical reflection on and a diving refraction into the underlying layer of more compact sediment with a smaller gradient and V 2 ~3 km/s. Branch 3 in brown is the precritical reflection on top of and refraction-head wave into a V 3N5 km/s medium, interpreted as the pre-rift basement, the top of the Hellenic nappes after the Alpine orogeny. The sediment thickness under OBS 4 is tightly constrained to 2 km by the slopes and intercept times of the refracted travel-time curves. On the split-spread profile reduced with the 4.5 km/s velocity of the basement (Fig. 3a), the intercept or delay time of the basement refraction only slightly increases towards south; hence, the maximum depth of the base of the sedimentary basin at the foot of this fault is no more
than 2.7 km. Further south, the basement travel-time curve shows a significant kink, visible in Fig. 3b around 3.5-km offset, with early arrivals due to traversing the master fault into the higher footwall block. Data from OBS 2 (not shown) on the same line reverse this profile, for which the 2D model is controlled by ray-tracing in Fig. 3c. The record section of a parallel line of shots, C, through OBS 1, located closest to the offshore continuation of the Xylocastro fault farther east is shown in Fig. 3b. The basement wave on its split-spread profile is correspondingly very asymmetric. It yields a value of 2.5–3 km for the maximum basement depth. Such in situ refraction measurements are considered to give the best geophysical constraint on basement depth. The value thus measured for the sediment thickness is unexpected. Indeed, Armijo et al. (1996) used a general model of a thick plate and a single high-angle fault, which accounted for the footwall uplift they constrained from the glacioeustatic markers of the Corinth marine terraces. The corresponding sediment thickness on the subsiding hanging wall resulted as 5 km from this modeling. The much smaller value that we measure may enter as a constraint in further modeling attempts. Westaway (2002) has assumed such a small value of sediment thickness and discussed rheological concepts that allow him to model that realistic value.
5. Imaging of sediments, basement topography, and intra-basement structure by multichannel reflection The 24-fold stack of multichannel reflection profile A along the axis of the Gulf (Fig. 4a) allows reduction of the strong sea-bottom multiples that, along with the
Fig. 3. Examples of OBS data and modeling, NW is to the left, SE to the right. Xylocastro fault is assumed to continue eastward forming the deep basin edge just south of OBS 1 and 2. (a) OBS 4 in the eastern part of the basin, in the hanging wall of the Xylocastro fault. Record section with 4.5 km/s reduction velocity shown as a split-spread profile across the fault for the shot line B. Travel time branches correlated are: in yellow, diving wave in upper sediment layer with velocity increasing with depth from 2.0 to 2.2 km/s; in blue, lower sediments with velocity increasing from 3 to 3.5 km/s; in brown, basement with velocity increasing from 5.3 to 6 km/s. (b) OBS 1 located just basinward of the Xylocastro fault. Record section with 4.5 km/s reduction velocity of data recorded for the parallel shot line C. Note the difference in the arrival time of basement refractions as brown first arrivals on either side, which resolves the greater depth of the basement under the deep marine basin. (c) Line B velocity modeling and ray-tracing for OBS 2 and OBS 4. Lower frame displays the arrival times picked for the observations and the computed travel-time curves as black lines, represented with 6 km/s reduction velocity. Upper frame shows the corresponding two-dimensional velocity model containing a water layer, then upper and lower sedimentary layers at average 2 and 3 km/s with velocity–depth gradients and basement with 5 km/s velocity at the top and velocity–depth gradient.
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different strengths of the seismic sources, limited single-channel penetration. The western, crooked part is only a partial brute stack due to the effect on streamer noise and the geometry of the ship’s turning and is shown for the sake of completeness. Fig. 4a shows a series of strong reflections with echo times of up to 2–3 s, indicating a basin with stratified sediments. In order to preserve the characteristic signal frequency response of interfaces with depth, this profile is shown as a time section (restretched to time from the result of pre-stack depth migration) rather than the depth section of Fig. 4b, which stretches the waveforms within the high-velocity basement and thus gives a longer apparent signal period, so that the true longer-period character of these seismic waves could not be assessed. In the time section in Fig. 4a, the clear signature in the timedomain signal waveform is real, marking the usual low-frequency response of the top of basement that represents a former land surface. This high-velocity layer can thus be interpreted here as the pre-rift basement, indicating the subaerial land surface that developed at the top of the sequence of nappes that were emplaced during the Alpine orogenic evolution of the Hellenides. Unexpectedly, in contrast to the flat sea bottom along this axial line, this reflector at the base of this basin fill and the internal sediment layering both show significant variations in topography. They should instead be flat and horizontal if this axial line were a strike line to the extensional basin. The OBS refraction observations on profile A are consistent with this topography and the corresponding velocity and depth estimates of the pre-stack depth migration, for which a depth section is displayed in Fig. 4b. Along this reflection seismic section, three distinct basement
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deeps are revealed. They can each be identified as causing a local depocentre because there is neither a succession of horizontal nor of equal thickness layers over these basement deeps that would indicate filling or draping of a preexisting depression. The fact that there are three depocenters as imaged here requires that extensional evolution of this basin has been controlled by several distinct normal faults in the basement. Overall extension has thus occurred at an oblique strike to the axis of the Gulf, with segmentation at a shorter scale than the length spanned by this axial seismic profile. This is consistent with the observed geometry of faults to the south of the Gulf, which pass onshore to offshore from west to east (Fig. 1). This image does not support the view that the subsidence of this basin was controlled by a single normal fault striking parallel to the average N1208E trends of the south coast and basin axis. However, profile A is thus not perpendicular to the faults controlling these localized depocenters either, so it does not allow a straightforward analysis of the geometry of sedimentation and deformation. Inferences can nonetheless be drawn on the number and approximate location of controlling faults, as well as on the possible interpretation of the reflectors that are suggested in the basement in Fig. 4a. The southeastern depocenter in Fig. 4a is obviously in the hanging wall of the Xylocastro fault. Although these sediments are controlled by this fault, they do not show it, nor does the basement surface. This is because this profile is oblique to this fault and remains in its hanging wall. This profile does not cross the trace of this fault, which was instead imaged by line B (Fig. 3a). The western depocenter on profile A is in the hanging wall of faults that could correspond to the prolongation 15 km eastward into the basin of the
Fig. 4. (a) Right-hand part: profile A, shot with a 2900 cubic inch array of 14 guns in single-bubble mode (Avedik et al., 1996) at 50 m spacing, into a 96-channel, 2.4-km-long streamer, giving 24-fold coverage. This time section was obtained from a pre-stack depth-migrated section, restretched to time to illustrate the reflective character of the basement (which is outlined). Possible features in the basement (sketched) have dips opposite to that expected for Alpine nappes and could be younger faults. To the left, profile W (Fig. 1) is the crooked line, including a semicircular turn, forming the westward continuation of profile A. It is a lower-quality brute stack time section of partial data, due to noise and the ship’s turning. For orientation, letters indicate positions in Fig. 1; in particular, point E labeled on both parts of this figure corresponds to the crosspoint of profiles A and W in Fig. 1. The outline shaded in gray may be the base of the sediments. (b) Depth section of profile A obtained by pre-stack depth migration. Vertical exaggeration is approximately 6:1. Sediment layers are coloured, above the basement left black and white, in which imaging artefacts due to noise dominate as a result of increasing depth relative to streamer length. Lower sedimentary unit in brown, with subdued reflectivity and velocity 3.3 km/s, upper unit subdivided between main reflectors, layer velocities from 1.7 to 2. 2 km/s. Faults are tentatively sketched as black lines along disruptions of layers.
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Helike fault documented on land (Fig. 1) and whose footwall at Cape Akrata has uplifted marine terraces deposited during Pleistocene interglacial marine highstands (e.g., Armijo et al., 1996; Westaway, 1996, 2002). The middle depocenter provides evidence for control by another fault in between. We suggest that such a fault, extending down into the basement, corresponds to the deep expression of the feature previously mapped from the surface into the upper kilometer of the basin fill and interpreted as a listric growth fault in the sediments from a grid of singlechannel seismic lines (Brooks and Ferentinos, 1984; Higgs, 1988). From its strike mapped at shallow depth, this deep-reaching fault appears to be the eastward prolongation of a fault along the locally north-facing coast segment north of Derveni (Sorel, 2000), which passes onshore south of Cape Akrata (Fig. 1). Our reflection line also suggests the presence of intra-basement reflective structures. However, with a single profile, their identification and interpretation remain speculative. Nonetheless, westward-dipping reflectors that crop out at the basement–sediment interface, as sketched in a preliminary way on Fig. 4a, cannot be inherited Alpine nappe structures, as these would dip in the opposite direction, but they may be candidates for low-angle normal faults active during the recent evolution of the basin. Tectonic control of these basement deeps and depocenters thus occurs by more than a single normal fault, striking at an oblique angle, 308 counterclockwise, to the axis of the Gulf and the average orientation of its coastline but consistent with the present N–S direction of extension across the Gulf from GPS measurements and earthquake slip vectors (e.g., Baker et al., 1997; Clarke et al., 1998).
6. Discussion and conclusions The structure of the seismically active and rapidly extending continental rift in the Gulf of Corinth has been seismically imaged into the crust using new marine normal-incidence reflection and OBS wideangle reflection–refraction seismics. Although the middle and lower crustal response cannot be seen on the vertical reflection profile, as it was hidden in the very strong and complex water waves reflected off the shelves and coastlines along either side, the signal
energy was sufficient to penetrate to the Moho and to be recorded as wide-angle reflections by fixed stations on land. However, the extreme noise level due to storm conditions resulted in clear data at only two stations, beyond the western and northeastern ends of the shot line along the axis of the Gulf. The Moho depth is 40 km under the western Gulf north of Aigion and 32 km under its north coast, north of Corinth. Surprisingly, there appears to be no significant lowercrustal reflectivity, as is expected to result from extension (e.g., Bois and ECORS Scientific Party, 1990). This could suggest that large finite deformation, or lower-crustal flow, or a time lag after its occurrence may be needed for such a fabric to be seen. However, this deduction remains speculative because the effect of extension on lower-crustal reflectivity has been generally discussed using the vertical reflection response, which is here unattainable because of the extreme amplitude of basin-side water waves. However, the fact that intracrustal reflections are not even seen at wide angle may instead be because they are hidden by noise, as even the Moho reflection has a low signal-to-noise ratio. Using OBSs, which recorded our lines of shots as refraction profiles, the maximum depth of pre-rift basement in the hanging wall of the Xylocastro fault is measured as 2.7 km. This is significantly less than the 5 km expected from applying to the whole time period of sedimentation the model of a single steep fault in a thick elastic plate used to account for the uplift of the Plio–Quaternary Corinth marine terraces (Armijo et al., 1996) but consistently interpreted in the model of Westaway (2002). Our multichannel vertical reflection seismic image reveals that a sedimentary basin, consisting of several depocenters, underlies the flat sea floor. We suggest that these depocenters have been controlled by at least three normal faults, striking oblique to the seismic profile, which follows the axis of the Gulf. These faults would thus be oblique to the overall Gulf axis and the trend of its south coast and would have roughly east–west strikes, consistent with the present north–south direction of extension across the Gulf from GPS and earthquake slip vectors. These faults could be the along-strike eastward prolongations of the active faults that crop out on the south coast of the Gulf. The segmentation of these structures may relate to the low-angle dipping fault
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plane of the recent earthquake at Aigion in 1995, to the distribution of its aftershocks (Bernard et al., 1997), and to other possible low-angle normal faults in this western part of the Gulf of Corinth (Baker et al., 1997; Rietbrock et al., 1996).
Acknowledgments N/O Nadir and its multichannel seismic facility, operated by IFREMER, and R/V Filia (for OBS deployment) participated in this SEISGRECE cruise. We acknowledge the support of their masters and crews. This multichannel processing was initiated at Centre de Traitement Sismique at the Institut de Physique du Globe, Strasbourg. Pre-stack depth migration was facilitated by C. Ranero, GEOMAR, Kiel, through the Training and Mobility in Research Program of the European Union, under grant ERBFMGECT98-0108. R. Nicolich, L. Jolivet, and anonymous reviewers provided constructive criticism. We thank R. Westaway for helpful suggestions and editorial assistance.
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Aigion earthquake (Greece): evidence for low-angle normal faulting in the Corinth Rift. J. Seismol. 1, 131 – 150. Bois, C., 1990. Major geodynamic processes studied from the ECORS deep seismic profiles in France and adjacent areas. Tectonophysics 173, 397 – 410. Brooks, M., Ferentinos, G., 1984. Tectonics and sedimentation in the Gulf of Corinth and the Zakynthos and Kefallinia channels, western Greece. Tectonophysics 101, 25 – 54. Clarke, P., Davies, R., England, P., Parsons, B., Billiris, H., Paradissis, D., Veis, G., Cross, P., Denys, P., Ashkenazi, V., Bingley, R., Kahle, H.G, Briole, P., 1998. Crustal strain in central Greece from repeated GPS measurements in the interval 1989–1997. Geophys. J. Int. 135, 195 – 214. Cle´ment, C., 2000. Imagerie sismique crustale de la subduction He´lle´nique et du golfe de Corinthe. PhD thesis, Universite´ Denis Diderot–Paris 7, pp. 178. Gautier, P., Brun, J.-P., Jolivet, L., 1993. Structure and kinematics of the Upper Cenozoic extensional detachment on Naxos and Paros islands (Cyclades Islands, Greece). Tectonics 12, 1180 – 1194. Higgs, B., 1988. Syn-sedimentary structural controls on basin deformation in the Gulf of Corinth, Greece. Basin Res. 1, 155 – 165. Jackson, J.A., Gagnepain, J., Houseman, G., King, P., Papadimitriou, P., Soufleris, C., Virieux, J., 1982. Seismicity, normalfaulting and the geomorphological development of the Gulf of Corinth (Greece): the Corinth earthquakes. Earth Planet. Sci. Lett. 57, 377 – 397. Keraudren, B., Sorel, D., 1987. The terraces of Corinth (Greece)—a detailed record of eustatic sea-level variations during the last 500,000 years. Mar. Geol. 77, 99 – 107. Makris, J., Stobbe, C., 1984. Physical properties and state of the crust and upper mantle of the eastern Mediterranean sea deduced from geophysical data. Mar. Geol. 55, 347 – 363. Melis, N.S., Brooks, M., Pearce, R.G., 1989. A microearthquake study in the Gulf of Patras region, western Greece, and its seismotectonic implications. Geophys. J. Int. 98, 515 – 524. Papazachos, C., Nolet, G., 1997. P and S deep velocity structure of the Hellenic area obtained by robust nonlinear inversion of travel times. J. Geophys. Res. 102, 8349 – 8367. Rietbrock, A., Tiberi, C., Scherbaum, F., Lyon-Caen, H., 1996. Seismic-slip on a low-angle normal fault in the Gulf of Corinth: evidence from high-resolution cluster analysis of microearthquakes. Geophys. Res. Lett. 14, 1817 – 1820. Rigo, A., Lyon-Caen, H., Armijo, R., Deschamps, A., Hatzfeld, D., Makropoulos, K., Papadimitriou, P., Kassaras, I., 1996. A microseismic study in the western part of the Gulf of Corinth (Greece): implications for large-scale normal faulting mechanisms. Geophys. J. Int. 126, 663 – 688. Sachpazi, M., Hirn, A., Nercessian, A., Avedik, F., McBride, J., Loucoyannakis, M., the STREAMERS-PROFILES group, 1997. A first coincident normal-incidence and wide-angle approach to studying the extending Aegean crust. Tectonophysics 270, 301 – 312. Sachpazi, M., Hirn, A., Cle´ment, C., Charvis, P., Graindorge, D., Hello, Y., Daignie`res, M., 1998. New reflection and refraction investigations in the region of the Gulf of Corinth. Proceedings
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of the 1st Workshop on the Development of a Multiborehole Observatory at the Gulf of Corinth, Athens, October 26–27, 1997, chairman F.H. Cornet. Sorel, D., 2000. A Pleistocene and still-active detachment fault and the origin of the Corinth–Patras rift (Greece). Geology 28, 83 – 86. Taymaz, T., Jackson, J., McKenzie, D., 1991. Active tectonics of the north and central Aegean Sea. Geophys. J. Int. 106, 433 – 490. Tibe´ri, C., Lyon-Caen, H., Hatzfeld, D., Achauer, U., Karagianni, E., Louvari, E., Panagiotopoulos, D., Kassaras, I., Kaviris, G., Makropoulos, K., Papadimitriou, P., 2000. Crustal and upper mantle structure beneath the Corinth Rift (Greece) from a teleseismic tomography study. J. Geophys. Res. 105, 28159 – 28172.
Tibe´ri, C., Diament, M., Lyon-Caen, H., King, T., 2001. Moho topography beneath the Corinth Rift area (Greece) from inversion of gravity. Geophys. J. Int. 145, 797 – 808. Tsokas, G., Hansen, R., 1997. Study of the crustal thickness and subducting lithosphere in Greece from gravity data. J. Geophys. Res. 102, 20585 – 20597. Warner, M., 1990. Basalts, water, or shear zones in the lower crust. Tectonophysics 173, 163 – 174. Westaway, R., 1996. Quaternary elevation change in the Gulf of Corinth of central Greece. Philos. Trans. R. Soc. Lond. Ser. A 354, 1125 – 1164. Westaway, R., 2002. The Quaternary evolution of the Gulf of Corinth, central Greece: coupling between surface processes and flow in the lower continental crust. Tectonophysics 348, 269 – 318.
Reflection–refraction seismics in the Gulf of Corinth: hints at deep structure and control of the deep marine basin Christophe Cle´menta, Maria Sachpazib, Philippe Charvisc, David Graindorgec, Mireille Laiglea, Alfred Hirna,*, Giorgios Zafiropoulosd a
Laboratoire de Sismologie Expe´rimentale, De´partement de Sismologie, UMR 7580 CNRS, Institut de Physique du Globe de Paris, 4 Place Jussieu, Tour 14, B89, F-75252 Paris cedex 05, France b Geodynamics Laboratory, National Observatory of Athens, Lofos Nymfon, Athens, Greece c UMR Ge´osciences Azur-IRD, P.O. Box 3, 06235 Villefranche-sur-Mer, France d Hellenic Petroleum, Maroussi, Athens, Greece Accepted 3 June 2004 Available online 11 September 2004
Abstract The Gulf of Corinth is a natural laboratory for the study of seismicity and crustal deformation during continental extension. Seismic profiling along its axis provides a 24-fold normal-incidence seismic reflection profile and wide-angle reflection– refraction profiles recorded by sea-bottom seismometers (OBS) and land seismometers. At wide-angle incidence, the land receivers document the Moho at 40-km depth under the western end of the Gulf north of Aigion, rising to 32-km depth under the northern coast in the east of the Gulf. Both refraction and normal-incidence reflection sections image the basement under the deep marine basin that has formed by recent extension. The depth to the base of the sedimentary basin beneath the Gulf, constrained by both methods, is no more than 2.7 km, with ~1 km of water underlain by no more than ~1.7 km of sediment, less than what was expected from past modeling of uplift of the south coast in the East of the Gulf. Unlike the flat sea-bottom, the basement and sedimentary interfaces show topography along this axial line. Several deeps are identified as depocenters, which suggest that this axial line is not a strike line to the basin. It appears instead to be controlled by several faults, oblique to the S608E overall trend of the south coast of the Gulf, their more easterly strikes being consistent with the instantaneous direction of extension measured by earthquake slip vectors and by GPS. D 2004 Elsevier B.V. All rights reserved. Keywords: Gulf of Corinth; Aegean region; Seismic refraction; Reflection; Crustal structure; Rift; Basin; Extension
1. Introduction * Corresponding author. Tel.: +33 1 44273914; fax: +33 1 44273894. E-mail address: [email protected] (A. Hirn). 0040-1951/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2004.07.010
The Gulf of Corinth in central Greece (Fig. 1) is located in the back-arc region above the Hellenic subduction zone, in the post-Alpine extensional
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Fig. 1. The Gulf of Corinth. Location map of profiles A, B, and C shot by N/O Nadir and recorded by OBSs, located at numbered triangles and at land stations NAF and KLI (squares). The Moho reflection midpoint lines at these stations from shots on profile A are sketched as the two thick gray lines (offshore north of Aigion and onshore north of the eastern end of the Gulf). Only the axial profile A and its crooked turn in the West, W, could be recorded as a multifold seismic reflection profile because of safety reasons concerning deployment of the ship’s seismic streamer. Geology is adapted from Armijo et al. (1996), with earthquakes from Taymaz et al. (1991), Baker et al. (1997), and Bernard et al. (1997). Inset shows the study region (rectangle) in relation to its surroundings, including the North Anatolian Fault Zone (NAFZ).
domain of the Aegean (e.g., Gautier et al., 1993). This region has also been affected by the North Anatolian strike-slip fault (Fig. 1) since the Pliocene (Armijo et al., 1996). Strong seismic activity characterizes this region, which is extending in a southward direction, at a rate that exceeds 10 mm/year across the Gulf (Clarke et al., 1998). The south coast of the western Gulf of Corinth has experienced destructive earthquakes, such as the Helike event of 373 B.C. and the Aigion event of M s ~7 in 1861. From an early microearthquake study, Melis et al. (1989) proposed that the southern bounding faults of the asymmetrical graben forming the Gulf of Patras and the adjacent western Gulf of Corinth were listric, flattening northward into a midcrustal decollement. In the western Gulf of Corinth, microearthquakes recorded during temporary seismograph deployments (Rigo et al., 1996) appear distributed across the Gulf above a cutoff depth that has been considered as a plane dipping northward at a low angle of 108, from a depth of 8 km under Aigion. This plane has been interpreted as an underlying detachment representing the brittle–ductile transition, the base of the seismogenic layer, or as a fault in the brittle domain that is slipping through small earthquakes (Rietbrock et al., 1996). Focal mechanisms constrained
by waveform modeling of earthquakes of larger magnitude, 5.7 to 6.2 (Taymaz et al., 1991; Baker et al., 1997) have low-angle north-dipping nodal planes. A study of the 1995 M s 6.2 event, which caused damage at Aigion on the south coast but had its focus at depth beneath the northern margin of the Gulf, suggested a 338 northward dip for its fault plane, although aftershocks form a cluster elongated with a smaller apparent dip (Bernard et al., 1997). In the east of the Gulf, the 1981 sequence of three large M sN6 earthquakes provides the main data source. These are typical high-angle normal-fault events. Such a steep fault and a thick elastic plate have been used to model the uplift of the Plio–Quaternary marine terraces on the south coast of the eastern Gulf (e.g., Jackson et al., 1982; Keraudren and Sorel, 1987, Armijo et al., 1996; Westaway, 1996, 2002). The recent geological evolution of the Gulf, of which this earthquake activity is a present instantaneous expression, has accumulated finite extension, forming a morphological rift structure floored by a deep and flat marine basin elongated toward N1208E (i.e., S608E), cutting across the strike of Alpine structures. The mode of seismic energy release, on low-angle or high-angle normal faults, as well as the
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mechanism of extension of the continental crust on the geological time scale are debated for this structure that may be the world’s fastest-extending graben in continental crust. The structure of its marine infill and of the underlying basement are obvious markers of its style of extension. Previous attempts at deciphering this record of deformation from the structure of this basin were based on single-channel shallow-penetration seismics (e.g., Brooks and Ferentinos, 1984; Higgs, 1988). In order to acquire data to help constrain the evolution of this Gulf, we undertook seismic imaging using modern methods in order to provide observations of its structure, to introduce into this discussion a different point of view, complementary to onshore geological observations and their modeling, and studies of seismicity.
2. Methods and survey During the Franco–Greek SEISGRECE seismic survey in January 1997, the oceanographic vessel N/O Le Nadir of IFREMER shot a 2900 cubic inch array of 14 air guns operated in single-bubble mode (Avedik et al., 1996) every 50 m in the part of the Gulf to the east of Aigion (Fig. 1; Sachpazi et al., 1998). This mode of shooting an air gun source provides with maximum efficiency a signal that is peaked in the rather lowfrequency 12–20 Hz band but has a duration short enough for acceptable resolution for normal-incidence seismic reflection. This frequency band is high enough to correspond to that typically observed for reflections returned by the lower crust even in surveys using sources that also generate higher frequencies, like the more usual tuned arrays of air guns. This lowfrequency band also provides signal propagation to the large offset wide-angle reflections, despite attenuation. The shots in the Gulf of Corinth were recorded by six ocean-bottom seismometers (OBSs), operated on the bottom of the Gulf, as well as by seismometers at station sites on land (Fig. 1). The 96-channel, 2400m long seismic streamer could be deployed from the vessel for part of the survey. This allowed recording of the normal-incidence reflection profile A in Fig. 1 with 24-fold coverage, suitable for advanced processing with pre-stack depth migration. However, this was only possible for a 50-km distance along the axis of the Gulf because operation was impeded by the worst
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winter storm in the last 15 years. Acquisition could be maintained for a tie line, W in Fig. 1, at the western end of the Gulf, until crosspoint E (Fig. 1), but with a low signal-to-noise ratio and inadequate geometry as it was acquired along an arc of a circle with the ship turning. Although the strength of this seismic source has allowed us to image at normal incidence the whole crust elsewhere in the Aegean (Sachpazi et al., 1997), here the noise of extremely strong sideways reflection of water waves off the coast of the narrow Gulf dominates at times when reflections that could be returned from deeper than 15 km are expected.
3. Deep crustal elements from wide-angle reflections Land stations, offset at either end of the shot-line along the axis of the Gulf of Corinth (Fig. 1), recorded clear seismic waves out to the maximum recording distance of 105 km, showing the efficiency of signal generation and reflection. These waves arrive much later, as much as 5–6 s later at 50-km offset, than the first arrival Pg-wave refracted in the basement. With a high-velocity move-out, they thus cannot be interpreted as anything other than the PmP-phase: wideangle reflections from over 30-km depth, that is, from the Moho (Fig. 2). This is a rewarding case of a vertical reflection seismic source allowing one to obtain wide-angle reflections even in single coverage, without stacking for the Moho, and paves the way for the regional mapping of Moho topography using this shooting strategy with numerous receivers. Unfortunately, several other seismograph stations, which were probably slightly less protected from the worst winter storm during acquisition, were dominated by background noise. There was no previous seismic measurement of the Moho depth in this region, only inferences from gravity and regional tomography (Makris and Stobbe, 1984; Tsokas and Hansen, 1997; Papazachos and Nolet, 1997). From a teleseismic tomography experiment resulting in the usual horizontally layered medium with blocks of varying velocity, Tibe´ri et al. (2000) discussed lateral variations in average velocity of their upper blocks in terms of a change in the proportion of crustal and mantle material, that is, a variation in Moho depth. The artificial source wide-
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angle reflection observations in Fig. 2 indeed reveal a strong difference of 8 km in crustal thickness between the western tip of the Gulf west of Aigion and its northeast coast (Sachpazi et al., 1998; Cle´ment, 2000). The water depth and structure of the underlying sedimentary basin, determined from our vertical reflection profiling, are taken into account in the calculation of this variation in Moho, as reported in Fig. 2 and its corresponding caption. This is the simplest possible structural model and has an average crustal velocity typical for Europe, but other solutions are also possible. Propagation paths toward each of the recording stations all sample the crust under the profile of shots but do not sample the same Moho segment at midpoint. The condition that the Moho is continuous and straight, consistent with the average velocity used, could thus be relaxed. With this alternate assumption, Moho depths remain ~8 km larger in the west than the east, but the trade-off between velocity and depth results in a model space ranging from slower crustal velocities giving an overall shallower and convex-upward Moho to faster velocities giving an overall deeper and concaveupward Moho. The fact that Moho reflections are observed at as short an offset as 40 km could lead to prefer slower crustal velocities to shorten the critical distance. However, to achieve the amplitudes observed at small distances, other causes, such as Moho topography focusing the reflection, are necessary, so these observations alone do not provide a tighter constraint on crustal velocity. With data from only two locations, this thickness difference could simply be regarded as inherited from the Alpine phase of deformation, because a crustal root is still present, as attested by the Bouguer gravity low that cuts north–south at right angle to the Gulf (Tsokas and Hansen, 1997) to the west of the offshore area north of Aigion where the reflections to station NAF have their midpoint. After removing the effect of the Ionian slab, Tibe´ri et al. (2001) inverted the residual
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Bouguer anomaly in terms of a variation of crustal thickness that may as well explain the teleseismic tomography results of Tibe´ri et al. (2000). They find a similar difference in Moho depth between the two regions we have sampled with the wide-angle reflections (Sachpazi et al., 1998; Cle´ment, 2000). The Moho topography inverted from gravity shows a Moho with alternating highs and lows, including a low (i.e., thick crust) under the western Gulf and a high (i.e., thin crust) northeast of the Gulf of Corinth, which agree with the wide-angle reflection observations, plus another high south of the eastern end of the Gulf. The crust in this study region thus appears quite thick, somewhat unexpectedly given the common belief associating thin crust with rifts and other extensional domains. Apart from the Moho reflections, the wide-angle intracrustal response is very weak, which cannot be due to a too weak seismic source, as the response of the underlying Moho is strong. That the intracrustal reflectivity would be so weak may also be unexpected in this extensional region, given that strong lower crustal reflectivity has been observed in numerous vertical reflection and wide-angle profiles and attributed to extension (e.g., Allmendinger et al., 1987). This interpretation has also been proposed in western Europe, where this fabric has been acquired in the past, presumably as a result of post-Variscan crustal thinning (e.g., Bois and ECORS Scientific Party, 1990). The profiles that revealed this intracrustal reflectivity showed furthermore that unlike the layer-interface reflectivity well known in sediments, it may not be due to specular reflections on few continuous interfaces between lithological units. It has instead been considered that lower crustal reflectivity is caused by the heterogeneity of the medium and that this heterogeneity is related not only to the nature of the rocks but also to finite deformation or its rate and/or any contribution of magmatism (e.g., Warner, 1990). The strong active extension in this study region has evidently not yet
Fig. 2. Record sections from land stations NAF and KLI (Fig. 1) at variable offset, with 8 km/s reduction velocity, of air gun shots from line A along the axis of the Gulf. Note on both very late, clear waves, interpreted as Moho reflections, and weak upper crustal refracted arrival on NAF, from 3 to 5 s reduced time between 40 and 70 km. West is left and East is right. Superimposed in white are travel-time curves computed for a model including the water and sediment layers under the shot line, as derived from OBSs and vertical reflection profiling, which reveals Moho to be 8 km deeper in the west than in the east. In this model, an average 6.25 km/s velocity was used for the crust under the sediments, and Moho was assumed to dip at 10% (~5.78) from east to west. The Moho thus dips from the 32-km depth sampled at midpoints along the gray line in Fig. 1 under the north coast of the Gulf and east of the shot line (for shots to KLI), to the 40-km depth sampled at midpoints shown as the gray line in Fig. 1 offshore north of Aigion (for shots to NAF).
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achieved the crustal image seen in terranes inferred to have undergone extension long ago. This may indicate that the important factors governing whether extension is imprinted into the lower-crustal fabric are the finite amount of strain and the temperature conditions for flow.
4. Refraction seismic constraints on basement depth and sedimentary fill The velocity layering of the upper crust is clearly illustrated by the conspicuous branches of the traveltime curve in the wavefield recorded by OBS 4 located in the eastern widest and deepest part of the basin, in the hanging wall of the Xylocastro fault (Fig. 3a). The velocity–depth structure is derived from the split-spread profile recorded by this OBS for shots on line B, as follows. Branch 1 in yellow is a refractiondiving wave that constrains a velocity–depth gradient layer just beneath the sea bottom of sediment with a low seismic velocity V 1 of average 2 km/s. Branch 2 in blue is a precritical reflection on and a diving refraction into the underlying layer of more compact sediment with a smaller gradient and V 2 ~3 km/s. Branch 3 in brown is the precritical reflection on top of and refraction-head wave into a V 3N5 km/s medium, interpreted as the pre-rift basement, the top of the Hellenic nappes after the Alpine orogeny. The sediment thickness under OBS 4 is tightly constrained to 2 km by the slopes and intercept times of the refracted travel-time curves. On the split-spread profile reduced with the 4.5 km/s velocity of the basement (Fig. 3a), the intercept or delay time of the basement refraction only slightly increases towards south; hence, the maximum depth of the base of the sedimentary basin at the foot of this fault is no more
than 2.7 km. Further south, the basement travel-time curve shows a significant kink, visible in Fig. 3b around 3.5-km offset, with early arrivals due to traversing the master fault into the higher footwall block. Data from OBS 2 (not shown) on the same line reverse this profile, for which the 2D model is controlled by ray-tracing in Fig. 3c. The record section of a parallel line of shots, C, through OBS 1, located closest to the offshore continuation of the Xylocastro fault farther east is shown in Fig. 3b. The basement wave on its split-spread profile is correspondingly very asymmetric. It yields a value of 2.5–3 km for the maximum basement depth. Such in situ refraction measurements are considered to give the best geophysical constraint on basement depth. The value thus measured for the sediment thickness is unexpected. Indeed, Armijo et al. (1996) used a general model of a thick plate and a single high-angle fault, which accounted for the footwall uplift they constrained from the glacioeustatic markers of the Corinth marine terraces. The corresponding sediment thickness on the subsiding hanging wall resulted as 5 km from this modeling. The much smaller value that we measure may enter as a constraint in further modeling attempts. Westaway (2002) has assumed such a small value of sediment thickness and discussed rheological concepts that allow him to model that realistic value.
5. Imaging of sediments, basement topography, and intra-basement structure by multichannel reflection The 24-fold stack of multichannel reflection profile A along the axis of the Gulf (Fig. 4a) allows reduction of the strong sea-bottom multiples that, along with the
Fig. 3. Examples of OBS data and modeling, NW is to the left, SE to the right. Xylocastro fault is assumed to continue eastward forming the deep basin edge just south of OBS 1 and 2. (a) OBS 4 in the eastern part of the basin, in the hanging wall of the Xylocastro fault. Record section with 4.5 km/s reduction velocity shown as a split-spread profile across the fault for the shot line B. Travel time branches correlated are: in yellow, diving wave in upper sediment layer with velocity increasing with depth from 2.0 to 2.2 km/s; in blue, lower sediments with velocity increasing from 3 to 3.5 km/s; in brown, basement with velocity increasing from 5.3 to 6 km/s. (b) OBS 1 located just basinward of the Xylocastro fault. Record section with 4.5 km/s reduction velocity of data recorded for the parallel shot line C. Note the difference in the arrival time of basement refractions as brown first arrivals on either side, which resolves the greater depth of the basement under the deep marine basin. (c) Line B velocity modeling and ray-tracing for OBS 2 and OBS 4. Lower frame displays the arrival times picked for the observations and the computed travel-time curves as black lines, represented with 6 km/s reduction velocity. Upper frame shows the corresponding two-dimensional velocity model containing a water layer, then upper and lower sedimentary layers at average 2 and 3 km/s with velocity–depth gradients and basement with 5 km/s velocity at the top and velocity–depth gradient.
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different strengths of the seismic sources, limited single-channel penetration. The western, crooked part is only a partial brute stack due to the effect on streamer noise and the geometry of the ship’s turning and is shown for the sake of completeness. Fig. 4a shows a series of strong reflections with echo times of up to 2–3 s, indicating a basin with stratified sediments. In order to preserve the characteristic signal frequency response of interfaces with depth, this profile is shown as a time section (restretched to time from the result of pre-stack depth migration) rather than the depth section of Fig. 4b, which stretches the waveforms within the high-velocity basement and thus gives a longer apparent signal period, so that the true longer-period character of these seismic waves could not be assessed. In the time section in Fig. 4a, the clear signature in the timedomain signal waveform is real, marking the usual low-frequency response of the top of basement that represents a former land surface. This high-velocity layer can thus be interpreted here as the pre-rift basement, indicating the subaerial land surface that developed at the top of the sequence of nappes that were emplaced during the Alpine orogenic evolution of the Hellenides. Unexpectedly, in contrast to the flat sea bottom along this axial line, this reflector at the base of this basin fill and the internal sediment layering both show significant variations in topography. They should instead be flat and horizontal if this axial line were a strike line to the extensional basin. The OBS refraction observations on profile A are consistent with this topography and the corresponding velocity and depth estimates of the pre-stack depth migration, for which a depth section is displayed in Fig. 4b. Along this reflection seismic section, three distinct basement
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deeps are revealed. They can each be identified as causing a local depocentre because there is neither a succession of horizontal nor of equal thickness layers over these basement deeps that would indicate filling or draping of a preexisting depression. The fact that there are three depocenters as imaged here requires that extensional evolution of this basin has been controlled by several distinct normal faults in the basement. Overall extension has thus occurred at an oblique strike to the axis of the Gulf, with segmentation at a shorter scale than the length spanned by this axial seismic profile. This is consistent with the observed geometry of faults to the south of the Gulf, which pass onshore to offshore from west to east (Fig. 1). This image does not support the view that the subsidence of this basin was controlled by a single normal fault striking parallel to the average N1208E trends of the south coast and basin axis. However, profile A is thus not perpendicular to the faults controlling these localized depocenters either, so it does not allow a straightforward analysis of the geometry of sedimentation and deformation. Inferences can nonetheless be drawn on the number and approximate location of controlling faults, as well as on the possible interpretation of the reflectors that are suggested in the basement in Fig. 4a. The southeastern depocenter in Fig. 4a is obviously in the hanging wall of the Xylocastro fault. Although these sediments are controlled by this fault, they do not show it, nor does the basement surface. This is because this profile is oblique to this fault and remains in its hanging wall. This profile does not cross the trace of this fault, which was instead imaged by line B (Fig. 3a). The western depocenter on profile A is in the hanging wall of faults that could correspond to the prolongation 15 km eastward into the basin of the
Fig. 4. (a) Right-hand part: profile A, shot with a 2900 cubic inch array of 14 guns in single-bubble mode (Avedik et al., 1996) at 50 m spacing, into a 96-channel, 2.4-km-long streamer, giving 24-fold coverage. This time section was obtained from a pre-stack depth-migrated section, restretched to time to illustrate the reflective character of the basement (which is outlined). Possible features in the basement (sketched) have dips opposite to that expected for Alpine nappes and could be younger faults. To the left, profile W (Fig. 1) is the crooked line, including a semicircular turn, forming the westward continuation of profile A. It is a lower-quality brute stack time section of partial data, due to noise and the ship’s turning. For orientation, letters indicate positions in Fig. 1; in particular, point E labeled on both parts of this figure corresponds to the crosspoint of profiles A and W in Fig. 1. The outline shaded in gray may be the base of the sediments. (b) Depth section of profile A obtained by pre-stack depth migration. Vertical exaggeration is approximately 6:1. Sediment layers are coloured, above the basement left black and white, in which imaging artefacts due to noise dominate as a result of increasing depth relative to streamer length. Lower sedimentary unit in brown, with subdued reflectivity and velocity 3.3 km/s, upper unit subdivided between main reflectors, layer velocities from 1.7 to 2. 2 km/s. Faults are tentatively sketched as black lines along disruptions of layers.
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Helike fault documented on land (Fig. 1) and whose footwall at Cape Akrata has uplifted marine terraces deposited during Pleistocene interglacial marine highstands (e.g., Armijo et al., 1996; Westaway, 1996, 2002). The middle depocenter provides evidence for control by another fault in between. We suggest that such a fault, extending down into the basement, corresponds to the deep expression of the feature previously mapped from the surface into the upper kilometer of the basin fill and interpreted as a listric growth fault in the sediments from a grid of singlechannel seismic lines (Brooks and Ferentinos, 1984; Higgs, 1988). From its strike mapped at shallow depth, this deep-reaching fault appears to be the eastward prolongation of a fault along the locally north-facing coast segment north of Derveni (Sorel, 2000), which passes onshore south of Cape Akrata (Fig. 1). Our reflection line also suggests the presence of intra-basement reflective structures. However, with a single profile, their identification and interpretation remain speculative. Nonetheless, westward-dipping reflectors that crop out at the basement–sediment interface, as sketched in a preliminary way on Fig. 4a, cannot be inherited Alpine nappe structures, as these would dip in the opposite direction, but they may be candidates for low-angle normal faults active during the recent evolution of the basin. Tectonic control of these basement deeps and depocenters thus occurs by more than a single normal fault, striking at an oblique angle, 308 counterclockwise, to the axis of the Gulf and the average orientation of its coastline but consistent with the present N–S direction of extension across the Gulf from GPS measurements and earthquake slip vectors (e.g., Baker et al., 1997; Clarke et al., 1998).
6. Discussion and conclusions The structure of the seismically active and rapidly extending continental rift in the Gulf of Corinth has been seismically imaged into the crust using new marine normal-incidence reflection and OBS wideangle reflection–refraction seismics. Although the middle and lower crustal response cannot be seen on the vertical reflection profile, as it was hidden in the very strong and complex water waves reflected off the shelves and coastlines along either side, the signal
energy was sufficient to penetrate to the Moho and to be recorded as wide-angle reflections by fixed stations on land. However, the extreme noise level due to storm conditions resulted in clear data at only two stations, beyond the western and northeastern ends of the shot line along the axis of the Gulf. The Moho depth is 40 km under the western Gulf north of Aigion and 32 km under its north coast, north of Corinth. Surprisingly, there appears to be no significant lowercrustal reflectivity, as is expected to result from extension (e.g., Bois and ECORS Scientific Party, 1990). This could suggest that large finite deformation, or lower-crustal flow, or a time lag after its occurrence may be needed for such a fabric to be seen. However, this deduction remains speculative because the effect of extension on lower-crustal reflectivity has been generally discussed using the vertical reflection response, which is here unattainable because of the extreme amplitude of basin-side water waves. However, the fact that intracrustal reflections are not even seen at wide angle may instead be because they are hidden by noise, as even the Moho reflection has a low signal-to-noise ratio. Using OBSs, which recorded our lines of shots as refraction profiles, the maximum depth of pre-rift basement in the hanging wall of the Xylocastro fault is measured as 2.7 km. This is significantly less than the 5 km expected from applying to the whole time period of sedimentation the model of a single steep fault in a thick elastic plate used to account for the uplift of the Plio–Quaternary Corinth marine terraces (Armijo et al., 1996) but consistently interpreted in the model of Westaway (2002). Our multichannel vertical reflection seismic image reveals that a sedimentary basin, consisting of several depocenters, underlies the flat sea floor. We suggest that these depocenters have been controlled by at least three normal faults, striking oblique to the seismic profile, which follows the axis of the Gulf. These faults would thus be oblique to the overall Gulf axis and the trend of its south coast and would have roughly east–west strikes, consistent with the present north–south direction of extension across the Gulf from GPS and earthquake slip vectors. These faults could be the along-strike eastward prolongations of the active faults that crop out on the south coast of the Gulf. The segmentation of these structures may relate to the low-angle dipping fault
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plane of the recent earthquake at Aigion in 1995, to the distribution of its aftershocks (Bernard et al., 1997), and to other possible low-angle normal faults in this western part of the Gulf of Corinth (Baker et al., 1997; Rietbrock et al., 1996).
Acknowledgments N/O Nadir and its multichannel seismic facility, operated by IFREMER, and R/V Filia (for OBS deployment) participated in this SEISGRECE cruise. We acknowledge the support of their masters and crews. This multichannel processing was initiated at Centre de Traitement Sismique at the Institut de Physique du Globe, Strasbourg. Pre-stack depth migration was facilitated by C. Ranero, GEOMAR, Kiel, through the Training and Mobility in Research Program of the European Union, under grant ERBFMGECT98-0108. R. Nicolich, L. Jolivet, and anonymous reviewers provided constructive criticism. We thank R. Westaway for helpful suggestions and editorial assistance.
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