Fault architecture, basin structure and evolution of the

Basin Research (2009) doi: 10.1111/j.1365-2117.2009.00401.x

Fault architecture, basin structure and evolution of the Gulf of Corinth Rift, central Greece R. E. Bell, 1 n L. C. McNeill, n J. M. Bull, n T. J. Henstock, n R. E. L. Collier,w and M. R. Leederz n

National Oceanography Centre Southampton, University of Southampton Waterfront Campus, Southampton, UK wInstitute of Geological Sciences, School of Earth and Environment, University of Leeds, Leeds, UK zSchool of Environmental Sciences, University of East Anglia, Norfolk, UK

ABSTRACT The style of extension and strain distribution during the early stages of intra- continental rifting is important for understanding rift-margin development and can provide constraints for lithospheric deformation mechanisms.The Corinth rift in central Greece is one of the few rifts to have experienced a short extensional history without subsequent overprinting.We synthesise existing seismic re£ection data throughout the active o¡shore Gulf of Corinth Basin to investigate fault activity history and the spatio -temporal evolution of the basin, producing for the ¢rst time basement depth and syn-rift sediment isopachs throughout the o¡shore rift. A major basin-wide unconformity surface with an age estimated from sea-level cycles at ca. 0.4 Ma separates distinct seismic stratigraphic units. Assuming that sedimentation rates are on average consistent, the present rift formed at 1^2 Ma, with no clear evidence for along- strike propagation of the rift axis.The rift has undergone major changes in relative fault activity and basin geometry during its short history. The basement depth is greatest in the central rift (maximum 3 km) and decreases to the east and west. In detail however, two separated depocentres 20^50 km long were created controlled by N- and S-dipping faults before 0.4 Ma, while since ca. 0.4 Ma a single depocentre (80 km long) has been controlled by several connected N-dipping faults, with maximum subsidence focused between the two older depocentres.Thus isolated but nearby faults can persist for timescales ca. 1 Ma and form major basins before becoming linked.There is a general evolution towards a dominance of N-dipping faults; however, in the western Gulf strain is distributed across several active N- and S-dipping faults throughout rift history, producing a more complex basin geometry.

INTRODUCTION Basin subsidence and the eventual transition to sea£oor spreading are controlled by the development and interaction of fault systems established in the early stages of continental rifting. However, detailed observations of faulting at the onset and during the ¢rst few millions of years of rifting are rare owing partly to the small number of young active rift zones with a simple history not overprinted by later events.The geometry (Forsyth, 1992), number (Buck, 1991), location and strain rate of faults, together with patterns of subsidence recorded by syn-rift sediments (e.g. Rosendahl, 1987; Cartwright et al., 1995; Stewart & Argent, 2000; Gawthorpe et al., 2003; Le Pourhiet et al., 2003) pro vide valuable clues to deformation processes within both the brittle and ductile part of the crust. Detailed study of syn-rift deformation has provided insights into fault evo lution and rift propagation, for example in the North Sea 1

Present Address: GNS Science, 1 Fairway Drive, Avalon, New Zealand. Correspondence: R. E. Bell, GNS Science, 1 Fairway Drive, Avalon, New Zealand. E-mail: [email protected]

Viking graben (e.g. Cowie et al., 2005), East African rift (e.g. Hayward & Ebinger, 1996) and the Gulf of Suez (e.g. Gupta et al., 1999; Gawthorpe et al., 2003). These results can, and have been used to develop and test numerical and thermo -mechanical models (e.g. Buck et al., 1999) of deformation on a range of scales, from fault displacement and development through to lithospheric deformation. The o5-Myr- old (Ori, 1989) Corinth rift in central Greece is an ideal natural laboratory in which to analyse details of early rift history (Fig. 1). The rift has high strain rates, clearly imaged syn-rift in¢ll (both onshore and o¡shore) and structure uncomplicated by over-printing tectonics. Syn-rift deformation within the rift has been well studied at the scale of individual faults (e.g. McNeill & Collier, 2004; McNeill et al., 2007) and fault networks at the eastern (e.g. Leeder et al., 2005; Sakellariou et al., 2007) and western (e.g. McNeill et al., 2005b; Lykousis et al., 2007; Bell etal., 2008) ends of the rift. However, a systematic analysis of fault evolution across the entire rift system, or even the currently active o¡shore rift, has not previously been attempted. Su⁄cient data have now been collected and published o¡shore to enable an investigation of spatio -

r 2009 The Authors Journal Compilation r Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists

1

2

22°00′

22°30′

Domvrena

23°00′

–1000

0

1000

2000m

Pleistocene subsidence 10 mm/yr

Pleistocene uplift

Holocene subsidence

Holocene uplift

Terraces

Major active faults Intermediate active faults Inactive/buried faults Uncertain faults

Fig. 1. Tectonic framework of the Corinth rift showing major faults interpreted in this study and derived from Armijo et al. (1996), Sakellariou et al. (2001), Stefatos et al. (2002), Leeder et al. (2005), McNeill et al. (2005b), Palyvos et al. (2005), Rohais et al. (2007) and Bell et al. (2008). Major active faults have throws 4500 m and intermediate active faults throws o500 m. Inactive faults are de¢ned as having no reported displacement during the Quaternary. Minor faults which sole out onto the basement surface are not included.Topography is from the Shuttle RadarTopography Mission (http:// srtm.usgs.gov), bathymetry data from M.V.Vasilios 2003 (McNeill et al., 2005b) and R/V Maurice Ewing 2001 (Zelt et al., 2004) cruises. Uplift and subsidence measurements at the coastline (numbered circles) are explained in Tables 1 and 2. Arrows are north coast GPS velocity vectors (relative to the southern coastline), from Clarke et al. (1998) Fig. 16. Inset: Summary map of the Aegean region.

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R. E. Bell et al.


Evolution of the Gulf of Corinth Rift, Central Greece temporal patterns of basin subsidence and deposition within the rift in order to improve our understanding of fault and rift evolution. This study also provides observations that can be compared with numerical rift models which attempt to constrain fault mechanical properties and lithospheric processes. In this contribution, we build on the work of Armijo et al. (1996), Stefatos et al. (2002), McNeill et al. (2005b), Sakellariou et al. (2007) and Bell et al. (2008) to produce an updated fault framework for the Corinth rift and to extend stratigraphic correlation across the whole o¡shore rift to investigate basin subsidence patterns through rift history. This is accomplished through compilation of available seismic re£ection data (published and unpublished). We present new basement depth grids and sediment isopachs for the rift illustrating its complex and temporally variable depocentre geometry. This study also provides relative timings of fault activation and relative magnitude of activity, as well as a compilation of estimated individual fault slip rates for the rift. Our analyses provide the ¢rst integrated fault evolution model for the Corinth rift and o¡ers an additional rifting ‘case study’ which can be compared with other rifts worldwide and numerical fault evolution models.

TECTONIC AND GEOLOGICAL FRAMEWORK OF THE CORINTH RIFT The Corinth rift is a 100 30 km high- strain band in central Greece that experiences N^S extension at rates of 5^15 mm year 1 (e.g. McKenzie, 1972; Roberts & Jackson, 1991; Avallone et al., 2004;Fig. 1) and is younger and more focused than most other areas of contemporary continental extension, such as the East African rift system (Davidson & Rex, 1980; Hayward & Ebinger, 1996), the Baikal rift (Mats, 1993) and the Basin and Range province (Hamilton, 1987; Thatcher et al., 1999). Extension is thought to be related to some combination of: back-arc extension due to subduction of the African plate at the Hellenic Trench (McKenzie, 1972, 1978; Doutsos et al., 1988); westward propagation of the North Anatolian fault (Dewey & Sengor, 1979; Armijo et al., 1996); and gravitational collapse of lithosphere thickened in the Hellenide orogeny (Jolivet, 2001; Le Pourhiet et al., 2003). The Corinth rift is currently deforming due to activity on E^W to NW^SE orientated coastal south margin and o¡shore faults which have formed the Gulf of Corinth and Alkyonides Gulf (Fig. 1). The active modern rift probably also includes regions east of the Alkyonides Gulf (Tsodoulos etal., 2008) and the Saronic Gulf, but deformation here, where extension rates are reduced, is not considered further in this contribution. Inactive faults on the southern margin are responsible for early syn-rift braid fan and £uvio -lacustrine sedimentation in what is now the NW Peloponnese (Ori, 1989; Sorel, 2000; Collier & Jones, 2003; Rohais et al., 2007) and a potential 42 km of sedi-

ments beneath the Corinth Isthmus (King et al., 1988; Fig. 1). The age of the early syn-rift succession in the NW Peloponnese is poorly constrained but is thought to be Middle to Late Pliocene (Kontopoulos & Doutsos, 1985; Frydas, 1987, 1989). Throughout this paper, we use ‘Gulf of Corinth Basin’ to refer to the modern active rift controlled by the faults identi¢ed in bold in Fig.1, i.e. the Gulf of Corinth and the Alkyonides Gulf, and ‘Corinth rift’ to refer to both the Gulf of Corinth Basin and onshore regions in£uenced by earlier extension (dashed and bold faults in Fig. 1). A general northward migration of faulting across the southern margin has resulted in the progressive uplift of hangingwall sediments including Gilbert fan deltas (e.g. Kerauden & Sorel, 1987; Ori, 1989; Dart et al., 1994; Gawthorpe et al., 1994; Armijo et al., 1996; De Martini et al., 2004; Malartre et al., 2004; McNeill & Collier, 2004; Ford et al., 2007; Rohais et al., 2007; Fig. 1, Tables 1 and 2). Late Pleistocene and Holocene uplifted terraces and notches along the southern coastline of the Gulf of Corinth (e.g. Pirazzoli et al., 1994; Stewart, 1996) and current seismicity con¢rm that the south margin faults are active (Fig. 1 and Table 1). Estimates of surface uplift and hence slip rate for south margin faults have been deduced from these late Quaternary geomorphic features, which are summarised in Tables 1 and 2. Late Quaternary averaged uplift rates of 1^1.5 mm year 1 for the Gulf of Corinth south margin have led to estimates of fault slip rate based on several di¡erent methods in the range 3^ 11mm year 1 (Armijo et al., 1996; De Martini et al., 2004; McNeill & Collier, 2004; McNeill et al., 2005a). Uplift rates in the Holocene are generally higher, around 1.3^ 2.2 mm year 1, suggesting an increase in slip rates in the latest Quaternary (Stewart, 1996; Stewart & Vita-Finzi, 1996; Pirazzoli et al., 2004). Uplift rates in the Alkyonides Gulf are lower, averaging 0.3 mm year 1 over the late Quaternary (Leeder et al., 1991) and 0.5 mm year 1 for the Holocene (Pirazzoli et al., 1994; Kershaw & Guo, 2001). The northern coastline is more sinuous than the southern coast, suggesting subsidence (Stefatos et al., 2002), and this is con¢rmed in the Trizonia area by rates of vertical motion as recorded by GPS (Bernard et al., 2006). Observations from this study indicate subsidence along parts of the northern coastline (Fig.1; points 12, 13 and 14; Table 1); however, localised uplift is also observed (Fig. 1, points 15, 16 and 19; Table 1; Moretti et al., 2003). Short-term regional geodetic extension rates increase from o5 mm year 1 in the east to 410^15 mm year 1 in the west (Davies et al., 1997; Clarke et al., 1998; Briole et al., 2000; Avallone et al., 2004). This high geodetic rate in the western Gulf can be explained by combined activity on the N-dipping onshore faults and major active o¡shore faults (that dip both N and S) identi¢ed in seismic re£ection surveys (McNeill et al., 2005b; Bell et al., 2008). The distribution and polarity of these faults produce a complex basin structure rather than simple half graben (Stefatos et al., 2002; Moretti et al., 2003; Sachpazi et al., 2003;


3

R. E. Bell et al. Table 1. Summary of Holocene coastline vertical movement studies along the Gulf of Corinth shoreline categorized by the fault probably responsible for the movement Site number and location East Eliki fault 1. Diakofto 2. Paralia Platanou 3. Ladopotamos fan Derveni fault 4. Aegira/Mavra Litharia

Perachora fault: 5. Heraion 6. Mylokopi Skinos fault 7. Strava 8. Skinos 9. Mavrolimni East Alkyonides fault 10. Alepochori Psatha fault 11. Psatha

Study

Neo -tectonic indicator

Quantitative measurements

Pirazzoli et al. (1994), Stewart & Vita-Finzi (1996), Stewart (1996) Pirazzoli et al. (1994), Stewart & Vita-Finzi (1996), Stewart (1996) McNeill & Collier (2004)

Uplifted notches

Uplift 0.9^1.5 mm year 1

Uplifted notches

Uplift 0.8^2.1mm year 1

Uplifted beach and fauna

Uplift 1.4^2 mm year 1

Uplift of ancient harbour Uplifted notches

2 m of uplift since 2000 BP Uplift 1.3^2.2 mm year 1

Uplifted fauna.

Uplift 3.2 0.3 mm year 1

Pirazzoli et al. (1994) Kershaw & Guo (2001) Pirazzoli et al. (1994) Kershaw & Guo (2001)

Uplifted notches and fauna


Uplifted notches and fauna


Jackson et al. (1982),Vita-Finzi & King (1985) Jackson et al. (1982),Vita-Finzi & King (1985) Jackson et al. (1982),Vita-Finzi & King (1985)

Observed submergence

Subsided during 1981 earthquakes Subsided during 1981 earthquakes Subsided during 1981 earthquakes

Jackson et al. (1982),Vita-Finzi & King (1985)

Observed emergence

Uplifted during 1981 earthquakes

Jackson et al. (1982),Vita-Finzi & King (1985) Leeder et al. (1991)

Observed emergence

0.2 m uplift in 4 years since 1981 earthquakes Uplift 0.3 mm year 1

Pirazzoli et al. (1994) Mouyaris et al. (1992) Stewart & Vita-Finzi (1996) Pirazzoli et al. (2004)

General Alkyonides Gulf location 12. Porto Germeno Vita-Finzi & King (1985) and this study 13. Aliki Vita-Finzi & King (1985) and this study 14. Livadostros This study General North Gulf coast 15. Pangalos/ Mikalanos peninsula 16. Pangalos/ Mikalanos peninsula 17. Pangalos/Mikalanos peninsula 18.Velanidia peninsula 19. Galaxidi

This study Moretti et al. (2003) Papatheoderou & Ferentinos (1995) Papatheoderou & Ferentinos (1995) This study

Observed submergence Observed submergence

Uplifted notch Submerged 2400 BP archaeological remains Submerged 2400 BP Archaeology Possible submerged multiple notches

Subsidence 0.68 mm year 1

Possible uplifted wave cut platform Uplifted beach rock

None

Submerged wave cut platforms Submerged wave cut platforms Possible uplifted wave cut platforms

Subsidence 0.83 mm year 1 None

Beach rock under dating None None None

Numbered locations are shown in Fig. 1.

McNeill et al., 2005b; Bell et al., 2008). An alternative model involving dominance of a deep, low-angle Ndipping detachment zone has been proposed (Rigo et al., 1996; Sorel, 2000; Gautier et al., 2006), but is not compatible with onshore footwall uplift and topography (Armijo et al., 1996), o¡shore basin stratigraphy (Bell et al., 2008) or earthquake focal mechanisms with fault plane dips stee-

4

per than the detachment plane (Hatzfeld et al., 1996; Bernard et al., 1997, 2006). Cores sampling shallow subsurface sediments in the Gulf of Corinth Basin suggest the basin ¢ll includes hemipelagic, turbidite and debris £ow sediments, with fan delta deposits common along the southern margin (Brooks & Ferentinos, 1984; Moretti et al., 2004). However, the more


Evolution of the Gulf of Corinth Rift, Central Greece Table 2. Summary of Late Quaternary vertical movement rates from studies along the Gulf of Corinth shoreline Site number and location Psathopyrgos fault 20. Aigion fault 21. West Eliki fault 22. East Eliki fault 23.

Xylokastro fault 24. Perachora peninsula 25. Perachora peninsula 26. SE Alkyonides fault, Alepochori

Study

Neo -tectonic indicator

Quantitative measurements

Houghton et al. (2003)

Uplifted fauna


De Martini et al. (2004), McNeill et al. (2007)

Uplifted terraces Displaced stratigraphy

Uplift 1^1.2 mm year 1

De Martini et al. (2004)

Uplifted terraces

Uplift 1.25 mm year 1

McNeill & Collier (2004), De Martini et al. (2004), McNeill et al. (2005a)

Uplifted terraces

Uplift 0.9^1.1mm year 1

Kerauden & Sorel (1987), Armijo et al. (1996)

Uplifted terraces

Uplift 1.3 mm year 1

Leeder et al. (2005) Leeder et al. (1991)

Uplifted terraces Uplifted terraces

Uplift 0.2^0.3 mm year 1 Uplift 0.2^0.3 mm year 1

Numbered locations are shown in Fig. 1.

deeply buried basin ¢ll has never been sampled and its sedimentary facies cannot be con¢rmed. The Gulf of Corinth is connected to the Mediterranean at its western end by the Rion sill (Straits of Rion, Fig. 1), which lies 60^ 70 m below sea level and produces an isolated lake during marine lowstands (Collier et al., 2000; Perissoratis et al., 2000). Given uplift rates in the central Corinth Isthmus area of 0.3 mm year 1 (Collier et al., 1992; Dia et al., 1997), a marine connection at the eastern end of the rift would have existed before about 300 ka and during the 200 ka highstand.The latter is supported by the presence of oolitic carbonate deposits in the southern Isthmus which would have accumulated in a shallow, high- energy seaway just before emergence of the southern Isthmus (Collier & Dart, 1991). The presence of lacustrine conditions within the Gulf of Corinth Basin during subsequent glacio - eustatic lowstands is con¢rmed by o¡shore sediment cores (Collier et al., 2000; Moretti et al., 2004). This glacialînterglacial variation in depositional environment produces a strong signature in seismic stratigraphy which has allowed interpretation of an approximate chronology (Perissoratis et al., 2000; Sachpazi et al., 2003; Leeder et al., 2005; McNeill et al., 2005b; Lykousis et al., 2007; Sakellariou et al., 2007; Bell et al., 2008).

METHODOLOGY Datasets We have compiled seismic re£ection and swath bathymetry data (both original data and published pro¢les) from a number of o¡shore studies, summarised in Table 3 and Fig. 2. High-resolution multichannel seismic re£ection pro¢les from the 2003 M.V. Vasilios survey (Table 3 and Fig. 2, McNeill et al., 2005b; Bell et al., 2008) are our pri-

mary data source in the western Gulf of Corinth.The area to the west of Aigion (i.e. theTrizonia basin) is not studied in detail here. Pro¢les in the central Gulf of Corinth are sparser but are orientated such that they give good general coverage at a range of resolutions and penetration depths. These include pro¢les from the RRS Shackleton 1982 dataset, described by Brooks & Ferentinos (1984), Higgs (1988) and Stefatos et al. (2002) and the three deep penetrating lines published by Clement (2000), Sachpazi et al. (2003) and Clement et al. (2004). Few published pro¢les exist in the area of the Lechaion Gulf and this depocentre is not included in our analysis. In the Alkyonides Gulf, we have used sparker and pinger data from the 1996 M.V.Vasilios survey (Collier et al., 2000; Leeder et al., 2002, 2005) and published pro¢les from Sakellariou et al. (2007). Finally, pro¢les from the R/V Maurice Ewing 2001 survey throughout the Gulf (Goodli¡e et al., 2003; Zelt et al., 2004 and lower resolution images of the dataset from http://www.ig.utexas.edu/sdc) providing broad- scale structure (i.e. basement depth) at key locations have been incorporated.

Depth conversion Depth conversion to quantify fault activity was based on typical interval velocities from stacking of 2003 M.V.Vasilios multichannel data, velocity estimates from cores (Collier et al., 2000; Moretti et al., 2004) and generalised sediment velocity curves (Hamilton, 1979, 1980). We estimate that sediments with a two -way travel time (TWTT) of between 0^0.5 and 0.5^1s below the sea£oor have average velocities of 1.5^2.0 and 2.0^2.5 km s 1, respectively. Our depth conversions are consistent with published depth sections in the Gulf also used in our analysis (e.g. Clement et al., 2004).


5

6

Alkyonides 1996 survey (Sakellariou et al., 2007)

R/V AEGAEO

R/V Maurice Ewing 2001- higher res. images R/V Maurice Ewing 2001- lower res. images

SEIS−GRECE 1997

M.V. Vasilios 1996 Hellenic Petroleum 1985

M.V. Vasilios 2003 RRS Shackleton 1982

22°00′

Gul f of

22°30′

Cor int h

Fig. 2. Map showing seismic pro¢les used in this study.Table 3 gives further details of each survey. Letters refer to pro¢les shown in Figs 10 and 11.

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R. E. Bell et al.


Evolution of the Gulf of Corinth Rift, Central Greece Table 3. Summary of seismic data compiled for use in this study

Survey

Source

Maximum TWTT from sea- surface (s)

RRS Shackleton 1982n

Single- channel air-gun

2

DEP-Hellenic Petroleum 1985n M.V.Vasilios 1996n SEIS-GRECE 1997n

Maxipulse explosive source Multi- channel Single- channel sparker Air gun

6 0.85 4

R/V Maurice Ewing 2001

Multi- channel air-gun

6

M.V.Vasilios 2003w

Multi- channel 150^ 2000 Hz Sparker Single channel airgun Single channel air gun

1.5

Brooks & Ferentinos (1984), Higgs (1988), Stefatos et al. (2002) Sachpazi et al. (2003), Moretti et al. (2003), Clement (2000) Leeder et al. (2002, 2005) Collier et al. (2000) Clement (2000), Clement et al. (2004), Sachpazi et al. (2003) Goodli¡e et al. (2003), Zelt et al. (2004) UTIG Marine Seismic Data Centre McNeill et al. (2005b), Bell et al. (2008)

1.8 0.9

Lykousis et al. (2007) Sakellariou et al. (2007)

R/VAEGAEO Gulf of Alkyonides 1996

Publications

n

Surveys for which we have access to some unpublished pro¢les or higher quality images to supplement published data. wData collected by us. No stars: Seismic pro¢les viewed from publications or data archive websites only. TWTT, two-way travel time.

FAULT ARCHITECTURE IN THE GULF OF CORINTH BASIN We present a new interpretation of faulting (Fig. 1) within the Gulf of Corinth Basin, improving the earlier fault geometries for the central basin presented by Brooks & Ferentinos (1984), Higgs (1988), Armijo et al. (1996) and Stefatos et al. (2002), and building on detailed interpretations in the western Gulf of Corinth (McNeill et al., 2005b; Bell et al., 2008) and Alkyonides Gulf (Leeder et al., 2002; Sakellariou et al., 2007). Below we discuss the major faults within each region of the o¡shore Gulf of Corinth and brie£y describe the character of major faults onshore from the literature.

Western Gulf of Corinth (Aigion to Akrata) We have previously described the complex fault structure of the western Gulf of Corinth (McNeill et al., 2005b; Bell et al., 2008) and present only a summary here (Fig. 1). O¡shore, the North and South Eratini faults each have a length of 15 km and overlap completely, resulting in the uplift of a prominent basement horst (McNeill et al., 2005b; Bell et al., 2008). A prominent axial channel draining the western rift is controlled by the S-dipping West Channel fault. At the eastern tip of the West Channel fault, the basin £oor widens and the northern margin becomes controlled by the S-dipping East Channel fault which steps 2 km to the north. Along the southern coastline, the East and West Eliki and Aigion faults control the position of the coastline and onshore topography (Fig. 1). The Akrata fault may be a splay of the major East Eliki fault, 4 km to its north. We have not examined structures to the west of Aigion, but published faults are indicated in Fig. 1. A series of largely inactive faults landward of the south coast (e.g. the Mamoussia^Pirgaki to Kalavrita faults) are thought

to have accommodated an earlier phase of extension (Ori, 1989; Sorel, 2000; Collier & Jones, 2003; Rohais et al., 2007).

Central Gulf of Corinth (Akrata to Xylokastro) The southern margin boundary fault between Akrata and Xylokastro has commonly been interpreted as a single N-dipping fault with a length of 30^40 km, named the Derveni or Corinth fault (Brooks & Ferentinos, 1984; Higgs, 1988; Stefatos et al., 2002). Others have split this fault into two segments (e.g. Moretti et al., 2003). We support the view that there are two fault segments bordering this part of the basin and interpret the break between the two segments to lie at the cusp in the coastline at Likoporia, based on available seismic data and coastal morphology (Fig. 1). We name the western segment the Derveni fault (interpreted originally by Brooks & Ferentinos, 1984; Higgs, 1988) and the eastern segment the Likoporia fault. The northern basin boundary is the S-dipping West Antikyra fault and the eastern extension of the East Channel fault (western end described by Bell et al., 2008). From the few seismic pro¢les that traverse the region south of Antikyra bay (Fig. 1), we have interpreted a number of additional short, minor active faults. The inactive onshore Xylokastro fault now lies in the footwall of the o¡shore Likoporia fault (Fig. 1; Armijo et al., 1996).

Eastern Gulf of Corinth (Xylokastro to Perachora peninsula) The south margin of the eastern Gulf of Corinth Basin is controlled by the N-dipping Xylokastro fault and NWdipping Perachora fault each with lengths of 12 km. Charalampakis et al. (2007) suggest that the surface trace of the Xylokastro fault forms three smaller segments. The connectivity between the inactive onshore Xylokastro


7

R. E. Bell et al. fault, Xylokastro fault and Perachora fault, from west to east, is unclear. Some authors have connected the Xylokastro and Perachora faults (e.g. Armijo et al., 1996), whereas others have not (e.g. Stefatos et al., 2002; Moretti et al., 2003).This study found that data availability prevents con¢rmation of either hypothesis. The coherent uplift of terraces on the south coast around Corinth could be evidence for one connected fault; however, uplift in this region may be partly driven by regional deformation (e.g. Collier & Dart, 1991; Westaway, 2002; Leeder et al., 2003, 2005; Leeder & Mack, 2007). The northern basin margin is de¢ned by shorter, less signi¢cant fault segments, such as the S-dipping East Antikyra fault (Fig. 1). A series of bur-

(a) 400

ied inactive faults produce additional localised basement subsidence (Fig. 1, the North and South Corinth Buried faults). The Kenchreai fault and a number of other faults form the southern boundary to the Pliocene^Quaternary sediments of the Corinth Isthmus (south of Fig. 1, discussed by Collier & Dart, 1991; Roberts, 1996; Goldsworthy & Jackson, 2001; Ghisetti & Vezzani, 2004).

Alkyonides Gulf Our interpretation of fault architecture closely follows Collier et al. (2000), Stefatos et al. (2002), Leeder et al. (2005) and Sakellariou et al. (2007). The most signi¢cant

S

N

1 km

TWTT (ms)

600

800

1000

VE ~ 4

1200 100 m

(b)

Long piston core tie

400 0

Marine (highstand) mud/silt Lacustrine (lowstand) /debris flow

50

Unit 600

Marine mud/silt ?

100 150

Axial channel sedimentation

200

Unit A

800

250

U

300

Unit B

350

1000

Basement

mu Gulf of Corinth enters lacustrine conditions

ltip

le

West Channel fault

1200

Fig. 3. (a) Uninterpreted and (b) Interpreted multichannel seismic pro¢le (M.V.Vasilios 2003 survey) from the western Gulf of Corinth, location in Fig. 2 (after Bell etal., 2008). Stratigraphy has been correlated with the eustatic sea-level curve of Siddall etal. (2003), modi¢ed for Gulf of Corinth lowstand depths. Horizon numbers relate to Oxygen Isotope Stage correlation. U is the unconformity between wellstrati¢ed unit A and poorly strati¢ed unit B.

8


Evolution of the Gulf of Corinth Rift, Central Greece W

E

Figure 6

Figure 5

5 km

800

TWTT (msec)

1000

1200

1400

1600

800

U

TWTT (msec)

1000 1

1200

U 5

1400

1600

7

9

U

Fig. 4. (a) Uninterpreted and (b) Interpreted E^W multichannel seismic pro¢le (M.V.Vasilios 2003 survey) across the western to central Gulf of Corinth, location shown in Fig. 2 (after Bell et al., 2008). Stratigraphy has been correlated with the eustatic sea-level curve of Siddall et al. (2003), modi¢ed for Gulf of Corinth lowstand depths. Horizon numbers relate to Oxygen Isotope Stage correlation. U is the unconformity between wellstrati¢ed unit A and poorly strati¢ed unit B.Vertical lines indicate intersection with seismic pro¢les shown in Figs 5 and 6.

faults are the N-dipping West and East Alkyonides o¡shore and the N-dipping Pisia and Skinos onshore (Fig. 1, described as the South Alkyonides fault segment by Morewood & Roberts, 2001). Although there are limitations in data coverage, the West and East Alkyonides faults do not appear to be linked at the surface and the East Alkyonides fault likely overlaps the western end of the West Alkyonides fault (Fig. 1). The eastern end of the East Alkyonides fault splays and overlaps with the Psatha fault to the east (Leeder et al., 2005; Sakellariou et al., 2007). Both of these faults trend E^Wand cut across the NW trend of the older Megara basin to the south (Fig. 1, Leeder et al., 1991; Collier et al., 1992). In the westernmost Alkyonides basin, faults control the Strava graben and uplifted Alkyonides islands (Papatheodorou & Ferentinos, 1993). Onshore, the Ndipping Pisia and Skinos faults are major structures controlling the signi¢cant topography (Jackson et al., 1982; Leeder et al., 1991, 2005; Collier et al., 1998). In the Alkyonides Gulf, S-dipping faults are generally less signi¢cant. For example, the Kaparelli fault on the north side of the rift was activated during the 1981 sequence of earthquakes (Jackson et al., 1982) but palaeoseismological studies suggest infrequent slip and low slip rates (Jackson et al., 1982; Benedetti et al., 2003; Kokkalas et al., 2007). The degree

of activity of other north margin faults (e.g. Domvrena, Livadostros, Germeno, Fig. 1) is unclear.

SYN-RIFT STRATIGRAPHY The syn-rift stratigraphy of the o¡shore Gulf of Corinth can be generally divided into two main units with contrasting seismic character (e.g. Moretti et al., 2003; Sachpazi et al., 2003; Bell et al., 2008). A major widespread unconformity separates the two units and is a key correlatable horizon. In the western Gulf, the stratigraphy has been interpreted in detail by Bell et al. (2008). Here, we extend this stratigraphic interpretation to the wider rift system whilst incorporating other published interpretations in order to constrain rift history.

Western Gulf of Corinth The stratigraphy of the western Gulf is fully described in Bell et al. (2008) and McNeill et al. (2005b) and is brie£y summarised here. The two stratigraphic units (A and B) can be clearly distinguished in the deep central rift basin separated by an unconformity. The older unit (unit B, Bell et al., 2008) is generally non-re£ective and poorly strati-


9

R. E. Bell et al. Cross over with Figure 4

Cross over with Sachpazi et al. (2003) Profile C (our Fig. 6)

WNW

ESE

0

TWTT (s)

1

2

3

10 km

4

VE ~ x5

N

S DER

ECF- fault zone

VE = x 1

¢ed, whereas the younger unit (unit A, Bell et al., 2008) is well strati¢ed and cyclical (Figs 3 and 4). Stratigraphic cyclicity in unit A is proposed to correlate with 100 ka climate-driven cycles, £uctuating between highstand marine and lowstand lacustrine conditions (Perissoratis et al., 2000; Sachpazi et al., 2003; Bell et al., 2008). Speci¢cally, unit A shows thin high-amplitude, low-frequency bands (marine highstand) alternating with thicker layers of lowamplitude and higher frequency (lacustrine lowstand) (Bell et al., 2008). Applying this chronostratigraphic interpretation produces an estimated age of the unit A/B unconformity of ca. 0.4 Ma (Bell et al., 2008). Bell et al. (2008) extended unit A sedimentation rates to unit B (which cannot be directly correlated with sea-level curves) to derive an age estimate of the oldest o¡shore sediments in the western Gulf at ca. 1.5 Ma (Bell et al., 2008). The change in sediment character o¡shore, between low re£ectivity monotonous unit B sediments and unit A cyclical sediments, may be related to changes in the intensity of glacial and inter-glacial cycles post 400 ka (OIS 11 e.g. Droxler & Farrell, 2000). Unit A’s strong cyclicity may be related to higher amplitude inter-glacial and glacial cycles

10

Fig. 5. Our interpretation of stratigraphy along the WNWÊSE along-axis central Gulf of Corinth SEIS^GRECE pro¢le (time section) following Clement (2000) and Clement et al. (2004). Location in Fig. 2. Interpretations are based on correlation with higher resolution data (e.g. seismic pro¢le Fig. 4) and seismic facies recognition. Numbered horizons correspond to Oxygen Isotope Stages as in Figs 3 and 4.

Fig. 6. Our interpretation of stratigraphy along the Hellenic Petroleum line, Pro¢le C (depth section) after Sachpazi et al. (2003) and Moretti et al. (2003). Locations in Fig. 2. Interpretations are based on correlation with higher resolution data (e.g. seismic pro¢le Fig. 4) and seismic facies recognition. Numbered horizons correspond to Oxygen Isotope Stages as in Figs 3 and 4. Note the detachment surface has been interpreted by Sachpazi et al. (2003), and we do not necessarily support the interpretation of this structure. DER and ECFare the Derveni and East Channel faults, respectively.

with 100 kyr periodicity after the Mid-Pleistocene climate transition, which would have a greater in£uence on stratigraphy than the more muted 41kyr oscillations prior (Mudelsee & Schulz, 1997). On the margins of the basin, well-imaged highresolution seismic stratigraphy, including clear lowstand deltaic clinoform packages, can be directly correlated with eustatic £uctuations, such as in the N Eratini fault hangingwall basin (McNeill et al., 2005b; Lykousis et al., 2007; Bell et al., 2008; Fig. 3).

Central and east Gulf of Corinth Data from the central Gulf presented by Clement et al. (2004) clearly demonstrate a distinction between younger wellstrati¢ed sediments and older poorly strati¢ed sediments separated by an unconformity (Clement, 2000; Clement et al., 2004; Fig. 5), as observed in the western Gulf by Bell et al. (2008). Our linkage of stratigraphy between the west and central basin, con¢rms that the unconformity is a common surface and can be correlated throughout. Sachpazi et al. (2003) tentatively estimate this


Evolution of the Gulf of Corinth Rift, Central Greece c1

(a)

c2

c3

c4

1 km

N

S

(b) III

1 km

IV V VI I

West Alkyonides Fault

II

III

Fig. 7. (a) Interpretation of Alkyonides Gulf seismic pro¢les from the 1996 M.V.Vasilios cruise, location in Fig. 2. Stratigraphic interpretation of the northern margin clinoforms after Leeder et al. (2005) (deviation in interpretation of Isotope Stages 11 and 13 horizons).The interpretation has been extended south to higher resolution sparker data. C1^C5 are end-lowstand clinoform slope breaks. Sea-level curve (Siddall et al., 2003) inset shows age correlation of lowstand clinoforms back to 400 ka. (b) Vertically exaggerated seismic stratigraphy on the southern Alkyonides coast adjacent to the West Alkyonides fault. Letters M and L refer to seismic facies types. M represents interpreted highstand marine deposition and L lowstand lacustrine deposition. Horizons 1, 3, 5 and 9 refer to Oxygen Isotope Stages. Detailed explanation of stratigraphy of units IÎV can be found in Table 4.

surface to be ca. 0.5 Ma; however, our more thorough stratigraphic interpretation of the same dataset supports an age of ca. 0.4 Ma. Cyclicity within unit A can be extrapo lated from the western to the central rift using the 1982 RRS Shackleton (Brooks & Ferentinos, 1984; Higgs, 1988; Stefatos etal., 2002) and SEIS-GRECE datasets (Clement, 2000; Sachpazi et al., 2003; Clement et al., 2004). By tracing our interpreted end-lowstand horizons along the E^W SEIS-GRECE pro¢le (Fig. 5, Clement, 2000; Clement et al., 2004) to other across-rift pro¢les (Fig. 6, Sachpazi et al., 2003, note this section is depth converted), we have obtained a consistent interpretation of unit A 100 ka stratigraphic cycles throughout the basin. Because the dis-

tinction between high-amplitude (highstand) and lowamplitude (lowstand) packages within unit A becomes less clear eastward, this interpretation is less reliable in places than the correlation of the clear A^B unconformity.

Alkyonides Gulf The stratigraphy of the Alkyonides Gulf has been extensively studied (Leeder et al., 1991, 2002, 2005; Collier et al., 1998, 2000). At least ¢ve buried lowstand clinoform packages exist on the subsiding northern margin, with the shoreline position shifting progressively northward with time (Fig. 7, Leeder et al., 2005). This stratigraphy

Fig. 8. (a) Depth to the basement^ sediment contact in two-way travel time (TWTT) based on seismic re£ection data.Western Gulf basement interpretation follows Bell et al. (2008) using M.V.Vasilios 2003 data, supplemented by additional interpretations from R/V Maurice Ewing EW0108/2001 data (Goodli¡e et al., 2003; Zelt et al., 2004, additional low-resolution images accessed through the Marine Seismic Data Center at http//www.igutexas.edu.sdc). Central Gulf interpretation includes data from the 1982 RRS Shackleton survey, two Hellenic Petroleum pro¢les (Sachpazi et al., 2003), SEIS^GRECE tie line (Clement, 2000; Clement et al., 2004), and additional control from R/V Maurice Ewing pro¢les (Goodli¡e etal., 2003; Zelt etal., 2004). Alkyonides Gulf interpretation from analysis of 1996 M.V.Vasilios data and additional control from R/V Maurice Ewing pro¢les (Goodli¡e et al., 2003; Zelt et al., 2004) and pro¢les within Sakellariou et al. (2007). (b) Total sediment isopach (TWTT) within the Gulf of Corinth Basin based on basement depth and bathymetry (Figs 1 and 8a). Fault plane widths and lengths are derived from seismic data and bathymetry. Fault width is de¢ned as the horizontal distance between a fault’s intersection with the sea £oor and basement. Lines indicate seismic pro¢les which penetrate r 2009 The Authors Journal Compilation r Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists

11

R. E. Bell et al.

0

500

1000

1500

0

400

800

1600

2000

2400

1200

38°00′

38°12′

(b)

38°00′

38°12′

22°00′ 38°24′

(a)

22°12′

22°24′

22°36′

22°48′

23°00′

23°12′

2800

TWTT depth to basement (ms)

TWTT sediment thickness (ms)

12


IV

III

II

I

Unit

Illustration

Few very strong re£ections Mostly low amplitude, high frequency. Re£ections that can be traced into clinoforms on the northern margin Few strong re£ections Transparent layer

Few very strong re£ections Lower amplitude, higher frequency re£ections are traced into clinoforms on the north margin Narrow band of strong re£ections Lower amplitude, higher frequency re£ections

Seismic Character Description

38

40

73

13

Average Unit thickness (m)

L. Lowstand/lacustrine

M. Highstand/marine

M. Highstand/marine L. Lowstand/lacustrine

M. Highstand/marine L. Lowstand/lacustrine M. Short lived marine? L. Lowstand/lacustrine

0

Geological Signi¢cance (M or L on Fig. 7)

The extrapolated horizon ages come from the correlation of lowstand clinoforms C1^C5 with the sealevel curve of Siddall et al. (2003) (see Fig. 7).

7

5

3?

1

Sea bed

Sequence OIS stage Boundary

Table 4. Summary of the interpretation scheme used to interpret Alkyonides basin stratigraphy (discussed in text and Fig. 7)

Ca. 240

ca. 130

ca. 50 ?

ca. 12

Extrapolated Horizon age (ka)

0.4

0.4

0.6

1.1

Sed. Rate (mm year 1)

Evolution of the Gulf of Corinth Rift, Central Greece


13

R. E. Bell et al.

0

200

400

600

1000

800

38°00′

38°12′

Unit A: Postca. 0.4 Ma 38°24′

(b)

38°00′

38°12′

(a)

22°12′ 22°00′ 38°24′ Unit B: Pre ca. 0.4 Ma

22°24′

22°36′

22°48′

23°00′

23°12′

1200

TWTT sediment thickness (ms)

14


Evolution of the Gulf of Corinth Rift, Central Greece AIG

SEF

N

NEF WCF AKR

EEF

0 –0.5 –1.0 (s) TWTT

Fig. 10. Interpretation of threedimensional (3D) subsurface structure of the western (zone 1) and central-western (zone 2) Gulf of Corinth showing variations in thickness of stratigraphic units A and B from M.V.Vasilios cruise data (after Bell et al., 2008). Line A is a N^ S-trending pro¢le which crosses the West Channel, South Eratini and North Eratini faults. Line B, to the east, crosses the East channel and Akrata faults. Line C (see also Fig. 4) is a tie line connecting lines A and B. Arrows indicate dominant stratigraphic tilt direction.

–1.5

5 km

–2.0 –2.5 –3.0

is similar to that of the North Eratini hangingwall basin, western Gulf of Corinth. The correlation of the clino form packages with eustatic sea level by Leeder et al. (2005) (which we support) suggests the oldest sediments on the northern Alkyonides margin are at least ca. 550 kyr old. Sakellariou et al. (2007) state that when the lowstand shoreline interpretation of Leeder et al. (2005) is extended southward into the main Alkyonides Gulf, the endlowstand horizons do not separate distinguishable cyclical packages. To test this hypothesis, we have extended endlowstand horizons from the clearly imaged clinoform packages (Fig. 7a) to other seismic pro¢les within the basin (including those of Sakellariou et al., 2007 Fig. 13).We ¢nd that the deeper basin stratigraphy can be separated into packages with a thin high-amplitude sequence at the base (Facies M, Fig. 7) overlain by a thicker lower-amplitude sequence (Facies L, Fig. 7), similar to that observed in the western rift, where thin high-amplitude units are interpreted as highstand marine deposits and thicker lowamplitude units are interpreted as lowstand lacustrine deposits. We therefore concur with the stratigraphy de¢ned by Leeder et al. (2005) and can extrapolate this stratigraphic interpretation to the deeper basin in a way which is consistent with our stratigraphy of the Gulf of Corinth Basin, but not consistent with Sakellariou et al. (2007)’s interpretation. Our age model predicts increased sedimentation rates during lowstands, in line with the results

N

Unit A: ca. 0.4 – 0 Ma Unit B: ca. 1.5 – 0.4 Ma

of Collier et al. (2000). A summary of our interpretation scheme and implications for sediment deposition pro cesses and environments is shown in Table 4. Deeper pro¢les in the Alkyonides basin show that sediments remain well strati¢ed down to basement and, although our interpretation implies sediments older than ca. 0.4 Ma are present, the seismic character is more similar to unit A than unit B (Sakellariou et al., 2007). Assuming sedimentation rates over multiple 100 ka cycles have remained on average constant in the Alkyonides basin, a reasonable assumption given the consistent seismic character of the sediments to basement, we estimate an age of ca. 1^1.5 Ma for the oldest sediments. This is comparable with the results of Leeder et al. (2005) but contrasts with the much younger ages of 0.5^0.7 Ma suggested by Sakellariou et al. (2007).

BASIN GEOMETRY Basement structure We have integrated basement depth from all seismic pro ¢les described in Fig. 2 and Table 3 to produce, for the ¢rst time, an approximate basement surface for the Gulf of Corinth Basin (Fig. 8a). This information is vital to establish the total displacement on o¡shore faults, the subsidence component of coastal faults, net extension across the o¡shore part of the Corinth rift and broad patterns

Fig. 9. (a) Unit B (ca. 1.5^0.4 Ma) two-way travel time (TWTT) sediment isopach. Sediments with unit B character are not identi¢ed in the Gulf of Alkyonides, however sediments of equivalent age have been interpreted and are shown here.Width of interpreted fault plane corresponds to horizontal extent of fault as seen within unit B. Length of fault during the deposition of unit B is estimated from the extent of major sediment depocentres. (b) Unit A (ca. 0.4^0 Ma) TWTTsediment isopach.Width of interpreted fault plane corresponds to horizontal extent of fault as seen within unit A. See Fig. 1 for full fault names.White areas outside dotted lines are regions of no interpretation due to data availability. r 2009 The Authors Journal Compilation r Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists

15

R. E. Bell et al. N

ECF AKR (possible EEF splay) DER

N LIK XY PER

~10 km

Unit A: ca. 0.4 – 0 Ma Unit B: ca. 1.5 – 0.4 Ma

Fig. 11. Interpretation of three-dimensional (3D) subsurface structure of the central-western (zone 2), central (zone 3) and eastern (zone 4) Gulf of Corinth showing variations in thickness of stratigraphic units A and B. Line D: our interpretation of WNWÊSE SEIS^ GRECE pro¢le (see also Clement, 2000; Clement et al., 2004); Line E and G, our interpretations of Hellenic Petroleum lines (see also Clement, 2000; Sachpazi et al., 2003); Line F, interpretation of line EW0108^35 from R/V Maurice Ewing dataset (Goodli¡e et al., 2003); Line H, interpretation of1982 RRS Shackleton survey pro¢le (Higgs, 1988; Stefatos et al., 2002). Arrows indicate dominant stratigraphic tilt direction.

of subsidence within the rift. The TWTT depth to basement map (Fig. 8a) con¢rms that the central Gulf of Corinth (between Akrata and the Perachora peninsula) is the deepest part of the basin, with basement at a depth of 2600^3000 ms ( 2700 to 3300 m). The east and west Gulf of Corinth and the Alkyonides Gulf are shallower; reaching a maximum depth of 1600 ms ( 1600 m) with localised basement highs in the form of fault- controlled horst blocks (Fig. 8a). The south margin of the central depocentre (Akrata to the eastern tip of the Perachora Peninsula) is tectonically controlled by the N-dipping enechelon East Eliki^ Akrata (eastern tip), Derveni, Likoporia, Xylokastro and Perachora faults (Fig. 8a). The northern margin of this depocentre is dominated by the East Channel fault in the west, but is more di¡use and in£uenced by a number of smaller S-dipping faults in the east (e.g. East and West Antikyra faults, Fig. 8a). Basement depth in the western Gulf of Corinth (Aigion to Akrata) is predominantly controlled by the S-dipping West Channel and South Eratini faults and the N-dipping

16

Eliki fault, although the seismic pro¢les available do not image basement depth in the immediate hangingwall of the Eliki fault (Fig. 8a; described in detail by Bell et al., 2008). McNeill & Collier (2004) suggest that there has been 2^4.5 km subsidence on the Eliki fault, which would indicate basement deepens signi¢cantly into the hangingwall of the East Eliki fault. The transition between the shallower western and deeper central Gulf of Corinth (variation in basement depth of 500 m between the west and centre, Fig. 8a) coincides with the western tips of the East Channel and Akrata faults, in a line between Eratini and Diakopto (Fig. 8a; Bell et al., 2008). A transition between shallow basement in the west and deeper basement in the centre of the Gulf of Corinth has also been suggested by Ghisetti & Vezzani (2004) and Papanikolaou & Royden (2007), but they place the transition farther east between Galaxidi and Akrata and propose that it is the result of inherited N^S structures.The compiled dataset used here shows no evidence of a change in basement depth or major N^S structural features at this location.


Evolution of the Gulf of Corinth Rift, Central Greece In the Alkyonides Gulf, applying the basement projection method of Leeder et al. (2005) to the 1996 M.V.Vasilios data, and additional basement depth constraints from the 2001 R/V Ewing data (Goodli¡e et al., 2003; Zelt et al., 2004) and pro¢les from Sakellariou et al. (2007), yields a maximum depth to basement in the hangingwall of the Alkyonides south coast faults of 1100^1350 ms ( 1300^ 1500 m) (Fig. 8a). The divide between the Alkyonides Gulf and the Gulf of Corinth involves a rapid deepening of the basement westward, although the exact nature of the boundary is unclear. A major tectonic boundary in this area has been proposed by McNeill & Collier (2004) based on a marked change in onshore uplift rates, and we concur that the Gulf of Corinth and Alkyonides Gulf are distinct basins. Major basement subsidence within the Gulf of Corinth Basin is con¢ned to the region between the N- and Sdipping faults de¢ning the southern and northern basin boundaries, respectively. The southern boundary faults (Aigion, Eliki, Derveni, Likoporia, Xylokastro, Perachora, West and East Alkyonides faults) are generally coincident with the modern coastline and the coastline follows the linear trend of the faults. In contrast, the E^W-trending major faults de¢ning the northern boundary of the main basin (e.g. North and South Eratini, West and East Channel, West and East Antikyra faults) are o¡shore and their footwalls form wide shallow platforms which have been £ooded to produce the sinuous trace of the northern margin.This structure suggests that in general hangingwall subsidence of south margin faults overwhelms uplift on north margin faults.

Syn-rift sediment thickness Isopach maps of syn-rift sediment thicknesses (Figs 8b, 9a and b), derived from depth to basement (Fig. 8a), sea£oor bathymetry and the unit A/B unconformity, record variations in basin subsidence and permit an analysis of fault evolution. Fault plane widths and lengths shown in the to tal sediment thickness isopach map (Fig. 8b) are derived from seismic data and bathymetry, with fault width de¢ned as the horizontal distance between a fault’s intersection with the sea£oor and basement.The length and width of faults given in the pre- and post-unconformity isopach maps have been estimated from the length and thickness of sediment depocentres controlled by these faults during each time period (Fig. 9a and b). Total sediment thickness is greatest in the central Gulf of Corinth Basin between the western tip of the East Channel fault and Perachora fault (Fig. 8b), and is coincident with the deepest basement (Fig. 8a). Sediment thickness in this area is relatively uniform between the north and south border faults, with a thickness of 1100^1500 ms (1300^1900 m), thickening locally into the hangingwalls of faults up to 1900 ms (2400 m) (Fig. 8b). Greatest sediment thickness in the western Gulf of Corinth is in the hangingwall of the West Channel fault (800 ms, 1000 m), although we do not have data to image sediment thickness adjacent to the Eliki fault. Sediment

thickness in the Alkyonides Gulf increases southward towards the East and West Alkyonides faults to a maximum thickness of 900 ms (1000 m), similar to that observed in the western Gulf of Corinth (Fig. 8b). Unit B in the central Gulf of Corinth Basin has accumulated within two main depocentres; one to the north of Akrata in the hangingwall of the East Channel Fault (41200 ms, 41500 m) and the other north of Xylokastro (4700 ms, 4900 m thick) (Fig. 9a). Between these two depocentres (Zone 3, Central in Fig. 9a), unit B is thin ( 200 ms, 250 m thick). There are two end member modes of origin for such a bimodal distribution of major sedimentation. Either: (1) the two depocentres are tecto nically controlled by activity on two separated fault systems creating accommodation space for sediments; or (2) di¡erences in sediment input form two lobes of positive sediment relief with minimal deposition between them. We favour the interpretation that unit B depocentres are controlled by fault activity because the depocentres lie directly within the hangingwalls of particular faults (the West and East Channel and East ElikiÂkrata fault system in the west, and the Xylokastro fault in the east). In further support of tectonic control, unit B (and rare stratigraphic re£ections within) thickens and dips towards the adjacent faults, with no evidence of deltaic patterns of sediment accumulation, which might be expected if the depocentres were controlled by sediment input at the rift margin. In addition, it is unlikely that 41-km-thick sediments could build up solely by depositional processes given likely water depths consistent with the present day. We, therefore, favour tectonics as the dominant control of depocentre formation here; however, we acknowledge that di¡erences in sediment input will result in di¡erences in sediment loading induced subsidence which will be superimposed on the tectonic subsidence. The Alkyonides Gulf has a thin ( 200 ms, 250 m thick) pre- ca. 0.4 Ma sediment thickness, although these sediments show a di¡erent seismic character to unit B sediments within the Gulf of Corinth (see ‘Stratigraphy’). The thickest deposits of the post-unconformity unit A (0^ca. 0.4 Ma) in the central rift lie directly in the hangingwalls of the Derveni, Likoporia and Xylokastro faults, again suggesting that deposition is structurally controlled. The stratigraphy of unit A also shows a strong tilt and thickening towards south coast boundary faults. Structural control is therefore favoured over a dramatic focused change in the location of major sediment input.

3D SPATIAL AND TEMPORAL EVOLUTION OF THE GULF OF CORINTH BASIN Syn-rift sediment isopachs (Fig. 9a and b) suggest strong spatial and temporal variations in fault activity within the Gulf of Corinth basin. Isopachs and seismic pro¢les (Figs 9^11) have been used together to determine relative fault activity along the rift and throughout its history. The o¡shore rift can be divided into ¢ve structural zones based


17

R. E. Bell et al. Method (a)

Parameters Determined

Problems/Assumptions

Total slip and slip rate during different time periods on a single fault

Hanging wall- footwall stratigraphic offset only seen for moderate faults

Footwall - hangingwall offset of stratigraphy θ

t1

h2

Assume no existing topography at time of horizon formation - i.e. the basin is fully-filled.

t0 h1

(b)

Both subsidence and uplift indicators Paleoshoreline of age t1

Estimate of slip and slip rate for time period t1

t1 t0

t1

U

S*

S

Estimate of slip may be based on a vertical displacement decay curve away from the fault plane (e.g. Armijo et al., 1996)

Assume vertical motion due to one fault only θ Total slip since t1 =

S* + U sinθ

(c) Uplifted or subsided shoreline only Uplift to subsidence ratio 1:1.2–2.2

Estimate of slip and slip rate for period t1

t1 U

t0

Estimate uplift based on U:S ratio (McNeill et al. 2005a) Assume vertical motion due to one fault only

θ U + 1.2U sinθ

< Total slip since t1
>0.45–1.2 mm/yr

–2000 –2500 –3000

(c) 1500 1000

Schematic relief

Depth (m)

500 0 –500

reduced unit B deposition

–1000 –1500

Likoporia fault

–2000 –2500 –3000

(d) 1500 Basement limestone

1000

Depth (m)

500

Unit B (ca. 1.5 – ca. 0.4 Ma)

0 –500 –1000 –1500 –2000

South coast fault system East/West Alkyonides/ Pstaha/Skinos

North coast fault systemLivodostros/ Germeno/Kaparelli ~0.8 –1.1 mm/yr

Onshore sediments (unknown correlation with offshore sediment) Area of subsidence Area of uplift

~0.8 –1.1 mm/yr

–2500

Fig. 14. (a) Map of major faults within the Corinth rift responsible for basement subsidence and unit B deposition in the Pliocene to Early Pleistocene (ca. 1.5 to ca. 0.4 Ma). Possible positions of the coastline and areas of shallow water are indicated. Arrows indicate directions of dominant stratigraphic tilt and stippled regions are sediment depocentres. Schematic Early Pleistocene cross sections across the; (b) western rift, (c) central rift, and d. Alkyonides Gulf (see A. for locations).

22


Evolution of the Gulf of Corinth Rift, Central Greece (a) Middle/Late Pleistocene to Recent ca. 0.4 - 0 Ma 22°00′ 38°30′

22°30′

23°00′

C South Era

tini

North Eratini

Aigion

We Cha st nne

D l

38°00′

Eliki

Derve

ni fau

lt Lik

opor

E. Alkyonides Psatha faults ia fa

ult Xylokas tro

ch or a

East

Skinos fault

a Pe r

Pisia fault

B

(b)

Distance along profile (km) 0

2

4

6

8 10 12 14 16 18 20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50

1500 1000

Depth (m)

500 0 −500

?

−1000

?

−1500

North Eratini fault 0.9 – 6.7 mm/yr

South Eratini fault E. Eliki fault/Akrata ~1.4 mm/yr West Channel ~3–5 mm/yr

−2000 −2500

fault ~0.4 mm/yr

−3000

(c) 1500 1000

Depth (m)

500 0

?

–500 –1000 –1500 –2000

West Antikyra fault Likoporia fault ~4 – 5 mm/yr

–2500 –3000

(d) 1500 1000

Basement limestone

Depth (m)

500

?

0 –500

North coast fault system Livodostros/Germeno/ Kaparelli

–1000 –1500 –2000 –2500

South coast fault system East/West Alkyonides/ Pstaha/Skinos ~1 – 3.3 mm/yr

Unit A (ca. 0.4 – 0 Ma) Unit B (ca. 1.5 – ca. 0.4 Ma) Onshore sediments (unknown correlation with offshore sediment) Area of subsidence Area of uplift

–3000

Fig. 15. (a) Map of major faults within the Corinth rift responsible for basement subsidence and unit A deposition from the Middle/ Late Pleistocene to recent (ca. 0.4^0 Ma). Arrows indicate directions of dominant stratigraphic tilt and stippled regions are sediment depocentres. Schematic recent cross section across the; (b) western rift, (c) central rift and (d) Alkyonides Gulf (see A. for locations).


23

R. E. Bell et al. In the western Gulf, McNeill & Collier (2004) and McNeill et al. (2005a) apply their footwall uplift rates to the elevation of the East Eliki fault footwall and estimate its age at ca. 0.7^1 Ma. A similar age is expected for the West Eliki Fault to the west with similar uplift rates and footwall elevation. Armijo et al. (1996) have derived an age of 1 Ma for the Xylokastro fault in the eastern Gulf based on the same factors. An age of ca. 0.8^2.2 Ma for initiation of the East Alkyonides fault, Alkyonides Gulf, has been proposed by Leeder et al. (2008), based on the age of the oldest sediments (2.2 Ma) and an uplifted calcrete (0.8 Ma) in the uplifted Megara basin footwall. The age of the unconformity within the Gulf of Corinth syn-rift sequence has been estimated to be ca. 0.4 Ma based on the interpretation of 100 kyr cycles (e.g. Figs 3 and 4). There are no consistent re£ectors below the unconformity, but Bell et al. (2008) extrapolated average sedimentation rates to estimate the age of the oldest basin sediments at ca.1.2^1.6 Ma in the western Gulf.We have not applied this method in the central Gulf to obtain a basin initiation age because fault slip rates have more clearly varied over time here. In the Alkyonides Gulf, correlation with 100 kyr cycles can be extended further to ca. 0.55 Ma (Leeder et al., 2005). Applying the average sedimentation rate from 0 to 0.55 Ma sediments to those below (sediments have similar character down to basement) yields an age of ca. 1^1.5 Ma for basin and fault initiation. This is similar to the results of Leeder et al. (2005), but contrasting with the much younger ages (0.5^0.7 Ma) of Sakellariou et al. (2007). The 1^1.5 Ma age is broadly consistent with estimates of fault initiation timing derived onshore in the Alkyonides Gulf. We recognize that a method of extrapolating average sedimentation rates is over- simpli¢ed in the Gulf of Corinth where the units above and below the unconformity may have been deposited at di¡erent rates, but the estimated age of oldest sediments in both the Gulf of Corinth and Alkyonides Gulf is comparable.

FAULT SLIP RATES Slip rate determination methods The study of depocentre evolution in the Gulf of Corinth basin has allowed us to make reliable assessments of relative fault activity in space and time. In the following section, we review quanti¢ed slip rates of Corinth faults from the literature, and use our new analysis to derive new estimates of slip for other faults. These slip rates are all reliant on the age model developed in this study and therefore a¡ected by errors in the chronology, but provide, in our opinion, reasonable estimates. The most accurate method of determining fault slip is to directly measure stratigraphic o¡set across the fault (Method A; Fig. 12a). Unfortunately, we can rarely apply this method to major active faults as they commonly straddle the coastline, with footwall onshore and hangingwall o¡shore, and individual horizons are di⁄cult to correlate.

24

Total basement o¡set can be measured in some cases, but slip rates rely on the estimated age of the oldest sediments above basement and the assumptions of no footwall erosion or pre- existing basement topography (Method D; Fig.12d). Similarly, with knowledge of the surface uplift and subsidence component of palaeoshorelines of equivalent age either side of a fault, we can estimate slip rates by extrapolating the far- ¢eld displacements to the fault plane using displacement decay laws (e.g. Armijo et al., 1996, Method B; Fig.12b). If only one component of displacement is known (footwall uplift or hangingwall subsidence) slip rates can be estimated by applying a locally veri¢ed uplift to subsidence ratio to estimate the other component or by modelling (Armijo et al., 1996; De Martini et al., 2004; McNeill & Collier, 2004, Method C; Fig. 12c). An uplift to subsidence ratio of 1 : 1.2^2.2 is suggested by McNeill et al. (2005a) for Corinth faults. This is subsequently supported by the ratio of onshore Late Quaternary uplift ( 1^1.2 mm year 1) to Holocene subsidence ( 1.3^ 2.5 mm year 1) across the Aigon fault (McNeill et al., 2007).

Fault slip rates within the Gulf of Corinth Basin Table 5 and Fig. 13 summarise slip rates within the Gulf of Corinth Basin from the literature and those derived in this study. Table 5 also states which of the four methods presented in Fig. 12 (Methods A^D) was used in the estimation of each slip rate. Some of these results are discussed in more detail below. Onshore and coastal faults Late Quaternary averaged slip rates for major southern coast faults are better constrained in the western and eastern Gulf of Corinth than in the centre, due to the presence of dated and correlated footwall marine terraces (e.g. Armijo et al., 1996; De Martini et al., 2004; McNeill & Collier, 2004). Using the methods and parameters described above (Fig.12c), uplift rates for faults in the western Gulf (East and West Eliki and Aigion faults) produce slip rates in the range 2.4^5 mm year 1 (De Martini et al., 2004; McNeill & Collier, 2004; McNeill et al., 2005a, 2007). Further west, uplift rates have been quanti¢ed for the Psathopyrgos fault footwall (Houghton et al., 2003). Using method C (uplift: subsidence ratio 5 1 : 1.2^2.2, fault dip 5 501; Fig. 12c), we estimate a slip rate of 2^ 3.5 mm year 1 for this fault (Table 5, Fig.13). In the eastern Gulf of Corinth, Armijo et al. (1996) have determined Late Quaternary uplift rates for the Xylokastro fault footwall of 1.3 mm year 1 from dated marine terraces. We have used our method C (Fig. 12c) to obtain a slip rate of 3.5^ 5.5 mm year 1 over the last 0.35 Ma for this fault. In the central Gulf, uplifted terraces in the footwall of the Derveni fault con¢rm Late Quaternary activity, but no uplift rates have been estimated due to reduced terrace preservation. The Derveni fault coastline has slightly higher Holocene uplift rates (1.3^2.2 mm year 1, Stewart & Vita-Finzi, 1996) than the East Eliki fault coastline (1^2 mm year 1, Stewart & Vita-Finzi, 1996). This sug-


Evolution of the Gulf of Corinth Rift, Central Greece gests that Late Quaternary slip rates for the Derveni fault may exceed those of the East Eliki fault (therefore 43^ 5 mm year 1). Moving eastward, in the footwall of the Likoporia fault, a 200 m shoreline is estimated at 125 ka by Flotte et al. (2001). Assuming this age is correct, we calculate an uplift rate of 1.6 mm year 1 equating to a slip rate of 4^5 mm year 1 using method C (Figs 12c and 13). In general, there appears to be a slight increase in Late Quaternary slip rates of coastal faults from the western Gulf of Corinth towards the central and eastern Gulf (Fig. 13). Late Quaternary slip rates for faults bordering the southern Alkyonides Gulf (e.g. Pisia, Skinos,West and East Alkyonides Gulf faults) are signi¢cantly lower than their equivalents in the Gulf of Corinth, and are in the range 0.3^2 mm year 1 derived from uplifted footwall terraces and deposits (Table 5, Fig. 13). Applying method C (Fig. 12c) to submerged hangingwall shorelines from ca. 0 to 0.34 Ma (Fig. 7 and Table 6), we have estimated the slip rate on the East Alkyonides fault at 1^2 mm year 1 (presuming this fault dominates subsidence o¡shore during this period). Before 0.34 Ma (in our age model), Sakellariou et al. (2007) suggest that subsidence was equally controlled by faults on the northern and southern basin margins. A total subsidence of 70 m between 0 and 0.34 Ma (Fig. 7 and Table 6) at a distance of 10 km from the south coast faults and 5 km from the north coast faults requires equivalent slip rates of 0.8^1.1mm year 1 on both faults.

O¡shore faults In the o¡shore Gulf, slip rates for the North Eratini fault on the northern margin of the western Gulf have been derived from submerged shorelines (Bell et al., 2008). O¡sets of stratigraphic horizons and the sediment-basement contact have also been used to estimate slip rates on other o¡shore faults (Table 5, Fig. 13). Slip rates averaged over the Quaternary for north margin western Gulf o¡shore faults (North and South Eratini, West and East Channel) range between 0.4 and 41.4 mm year 1.Where latest Quaternary (Holocene) subsidence and hence slip rates can be determined (e.g. North Eratini fault), rates are signi¢cantly higher than those averaged over longer periods (Bell et al., 2008). We have argued from the seismic stratigraphy that the western rift is locally a north-thickening half graben controlled by dominant S-dipping faults (McNeill et al., 2005b; Bell et al., 2008). However, the slip rates estimated for faults within the o¡shore western rift appear to be much lower than the rather better constrained slip rates for faults on its southern margin. Considering the basin subsidence and depocentre patterns (Figs 8 and 9) which provide evidence of relative fault activity, we suggest that the o¡shore fault slip rates we have determined here are likely underestimated, possibly due to basement erosion a¡ecting the calculation of slip rates using Method D (Fig. 12d), and are much more comparable with southern margin faults.

Holocene slip rates of faults throughout the Gulf, primarily derived from uplifted or submerged shorelines, are commonly higher than slip rates averaged over the Quaternary (Table 5, Fig.13). In the Alkyonides Gulf, Holocene fault slip rates are also higher than averaged Quaternary rates (Leeder et al., 2003); however, they are still signi¢cantly lower than Holocene slip rates elsewhere in the Gulf.

DISCUSSION Corinth rift evolution summary We have combined our results from the o¡shore rift with other observations o¡shore and onshore to produce, for the ¢rst time, an integrated fault evolution model for Corinth rift development based on all data currently available (Figs 14 and 15).We propose a multi- stage evolution of the rift summarised below. Stage 1: pre 2 Ma The earliest known syn-rift sediments (ca. 3.6^4 Myr old) within the Corinth rift are located onshore (Collier & Dart, 1991; Rohais et al., 2007).These sediments suggest that the original western rift was wide with activity distributed over a number of faults (Collier & Jones, 2003; Malartre et al., 2004; Ford et al., 2007; Rohais et al., 2007), although the location of the northern rift boundary is unknown. No deposition appears to have occurred during this Stage in the modern o¡shore basin, so we infer that this part of the rift underwent little subsidence. In the easternmost part of the rift system, extension was focused onshore at this time on the NW^SE-trending Megara basin south of the modern Alkyonides Gulf (Fig. 1) (Collier et al., 1992; Leeder et al., 2008). Stage 2: ca. 1^2 to ca. 0.4 Ma The period of distributed deformation in the western part of the rift ended when activity focused on the Mamoussia^ Pirgaki fault (Fig 14a and b) leading to deposition of hangingwall Gilbert fan deltas 1.5 km thick (Collier & Jones, 2003). We propose the timing of this transition, between Stages 1 and 2, to be during the Early Pleistocene (Ford et al., 2007; Rohais et al., 2007). Extension within the Megara basin, in the east, ceased ca. 0.8^2.2 Ma when activity transferred to the Alkyonides Gulf to the north (Fig. 14a and b, Leeder et al., 2008).The maximum age of sediments in the modern Gulf of Corinth basin is also estimated at ca. 1^2 Ma. To test this apparent coincidence of timing of the Stage 1^2 transition throughout the rift will require a detailed chronology of syn-rift stratigraphy, primarily in the o¡shore basin. The oldest sediments (Unit B^Stage 2) within the o¡shore Gulf of Corinth are deposited in three distinct depocentres (Figs 9a and 14a). In the west, sediments are deposited in a generally N-thickening half graben, but in£uenced by N-dipping faults to the south (Mamoussia^


25

R. E. Bell et al. Pirgaki and East Eliki; Fig. 14a and b). In the central rift, a second depocentre thickens southward towards the Xylo kastro fault (Fig. 9a).The lengths of these two depocentres are shorter than the modern fault traces controlling them, suggesting that the faults have increased in length (Fig. 9). The easternmost depocentre is in the Alkyonides Gulf where sediments thicken towards the south (Fig. 14d). These three depocentres are separated by basement highs with reduced sedimentation (Figs 8a, 9a, 14a and c). Stage 3: ca. 0.4 Ma to recent The unconformity separating Stages 2 and 3 sediments (units B and A, respectively, with distinct seismic character and depositional geometry) has an estimated age of ca. 0.4^ 0.5 Ma (this study; Sachpazi et al., 2003; Bell et al., 2008). Similarly, a change in the character of syn-rift sediments now uplifted onshore occurred at least ca. 0.4 Ma (Rohais et al’s 2007 upper group); however, given the uncertainty in chronology a link between these events cannot be con¢rmed. The change in o¡shore sediment character may be related to changes in the intensity of glacial and interglacial cycles after the mid-Pleistocene transition (e.g. Mudelsee & Schulz, 1997; Droxler & Farrell, 2000). Deposition during Stage 3 in the Gulf of Corinth Basin created a single depocentre with its thickest accumulation coincident with the thinnest Stage 2 deposition (Figs 9b and 15a). Faults controlling the region of greatest deposition (Derveni and Likoporia) therefore experienced a signi¢cant increase in displacement (Fig. 15a and c). Such an increase in late Quaternary displacement in the central rift has also been proposed by Pirazzoli et al. (2004). Sediment isopachs of this period suggest that slip rates of dominant faults in the central rift (Derveni, Likoporia) should exceed those of faults to the west and east (Eliki, Xylokastro) providing sediment £uxes into di¡erent parts of the basin were similar. The increased signi¢cance of the N-dipping Derveni fault relative to the S-dipping East Channel fault during Stage 3 (Fig.15a) produces a southward-thickening half graben, in contrast to the northward thickening half graben during Stage 2.The net product of subsidence and deposition during these two stages is a graben (Fig. 6, Sachpazi et al., 2003). In the western rift, subsidence within the o¡shore basin is controlled by both N- and Sdipping faults. In the Alkyonides basin, activity becomes focused on the N-dipping south margin faults (East Alkyonides, Pisia, Skinos and Psatha), creating a strongly south-thickening depocentre (Figs 7 and 15d). Stage 4: recent activity (latest Holocene) Recent activity is indicated by seismicity and geodetic extension rates (e.g. Ambraseys & Jackson, 1990; Bernard et al., 1997; Davies et al., 1997; Clarke et al., 1998; Briole et al., 2000; Gautier et al., 2006). Microseismicity is concentrated in the Aigion and Rion graben regions of the western Gulf (Doutsos et al., 1988; Hatzfeld et al., 2000; Gautier et al., 2006). It may be signi¢cant that the small

26

Rion graben (between the Straits of Rion and Nafpaktos; Fig. 1) separating the Gulf of Patras from the Gulf of Corinth is the site of seismicity clustering today (Doutsos & Poulimenos, 1992) and highest geodetic extension rates over the last 100 years (e.g. Clarke et al., 1998; Briole et al., 2000). The shallower basin depth between these basins is analogous to what we interpret to have occurred in the central Gulf of Corinth (Fig. 9a). Highs between two grabens are tectonically unstable (Doutsos & Poulimenos, 1992). Such instability may explain why short-term geodetic extension rates are high here, in a bid to resume equilibrium between the two graben and reduce the intervening topography.

Rift and fault evolution The proposed model for the Corinth rift highlights the following evolutionary trends: (1) Rifting initiates in iso lated depocentres which later become connected by enhanced fault activity/displacement between them; (2) The major faults responsible for controlling these initially iso lated basins dip both north and south, however as the rift evolves north-dipping faults become dominant; (3) Faulting generally migrates northward across the rifted region; (4) There is no evidence for along-axis propagation of the rift on a million year timescale. Similar trends have been proposed from numerical models of fault growth and linkage (Gupta et al., 1998; e.g. Cowie et al., 2000; McClay et al., 2002), and have been reported at other rifts through the analysis of seismic re£ection data and syn-rift sediments (Ebinger,1989; Hayward & Ebinger,1996; Gupta etal.,1999; Morley,1999; Contreras et al., 2000; Dawers & Underhill, 2000; McLeod etal., 2000;Taylor etal., 2004). Below we compare our interpretations with existing models of fault development and with other rift zones. 1. Fault and Rift Segmentation Continental rift basins often show slow initial subsidence followed by an accelerated second stage (Gupta et al., 1998). Numerical modelling suggests this is due to growth and linkage of isolated faults to form larger compound segments (Gupta et al., 1998; Cowie et al., 2000; McClay et al., 2002). One of the key questions concerns typical patterns of fault linkage in early fault development and on what timescale they occur (Morley, 2002). Comparison of model predictions with natural examples from rifts worldwide suggests wide variation in rate and timing of fault development. Morley (1999) suggests that fault linkage and propagation can occur either before signi¢cant basin deformation, after minor faulting has created a region of extensive shallow subsidence, or during the main phase of basin subsidence. Owing to the absence of isolated depo centres in the sedimentary record of the Lokichar Basin, East African rift, Morley (1999) proposes that fault linkage occurred before major basin deformation. In the Gulf of Suez, thin (200 m) isolated early basin ¢lls indicate linkage


Evolution of the Gulf of Corinth Rift, Central Greece during the earliest stages of basin development (Gupta et al., 1998). Similarly, in the o¡shore Whakatane graben, New Zealand, fault linkage around 1 Myr after rift initialisation resulted in a three-fold increase in displacement rate on the linked faults (Taylor et al., 2004). In contrast, the Usuanga £ats of the East African rift (Morley, 2002) and the Statfjord East ¢eld of the North Sea (Dawers & Underhill, 2000) have thick depocentres that are unconnected until much later in rift history. Within the Gulf of Corinth, isolated thick depocentres (41km of sediment) existed during Stage 2 between ca.1^2 and 0.4 Ma (i.e. during the ¢rst 0.6^1.6 Myr of rifting) before becoming linked during Stage 3 (post 0.4 Ma; Fig.9). Even with large potential errors in absolute age, the relative timings are su⁄ ciently constrained in the Gulf of Corinth to determine that segment linkage has occurred after individual depo centres became established and rift development was well underway, in agreement with examples from the East African Rift and North Sea. The central Gulf of Corinth appears to have evolved in a similar way to part of the Lake Malawi rift basin where three individual fault segments have dominated at di¡erent times (Contreras et al., 2000). In Lake Malawi, one fault dominates over the ¢rst ca. 6 Myr of rift history with activity switching to another fault for the next ca. 1 Myr, and, most recently, displacement has focused on a fault between the two earlier structures. In the Gulf of Corinth, enhancement of fault activity has also occurred in an area of limited displacement (Derveni and Likoporia faults at 22.51E). Contreras et al. (2000) also observe that the Malawi fault system has evolved to produce a ‘bell- shaped’ total fault displacement pro¢le as might be expected for a single fault. If we presume that sediment thickness is controlled largely by subsidence in the o¡shore Gulf of Corinth, cumulative sediment thickness (Fig. 8) suggests reduced total subsidence in the far west and east but a relatively consistent magnitude of subsidence along the centre of the basin. However, the most recent Stage of rifting (Fig. 9b) has a sediment thickness pattern expected from a generally bell- shaped and kinematically coherent rift subsidence and displacement pro¢le (e.g. Cartwright et al., 1995; Dawers & Underhill, 2000; Taylor et al., 2004). Faults along the south coast of the central Corinth rift are enechelon and do not appear to be hard linked at the surface. However, similar total displacement along and between individual fault segments and similar uplift pro¢les (Stewart & Vita-Finzi, 1996; McNeill & Collier, 2004), coupled with our ¢ndings of enhanced displacement in a region of earlier displacement de¢cit suggests that they may be linked at depth. 2. Fault and Rift Polarity The dip direction of dominant faults in the Corinth rift during Stages 1 and 2 is variable, with examples both to the N and S.This polarity varies both spatially and temporally. For example, during Stage 2, a large N-thickening depo centre controlled by S-dipping faults dominates the western rift whereas N-dipping faults produce two smaller S-thickening depocentres in the eastern rift (Fig. 9a).

Throughout rift development, there is an overall trend to wards N-dipping faults on the south coast becoming structurally dominant. The S-dipping East Channel fault (a prominent example) was instrumental in the formation of the large Stage 2 depocentre but was overwhelmed by activity on the N-dipping Derveni fault to the south in Stage 3 resulting in a net graben structure created by activity on both faults through time. Di¡erences in the polarity of half graben basins spatially and temporally are also observed in the East African rift (Rosendahl, 1987) and Gulf of Suez (Gupta et al., 1999). The emergence of a dominant fault dip direction in evolving fault systems is proposed to result from interacting fault planes at depth leading to unfavourably orientated slip vectors and the death of antithetic/conjugate faults (Jackson & McKenzie, 1983; Scholz & Contreras, 1998; Jackson, 1999). Numerical models also show that, in the presence of a lateral strain gradient, faults will dip to wards the direction of highest strain and others will cease to be active (Hardacre & Cowie, 2003). Our results in the Gulf of Corinth support these models, but highlight that, locally and over shorter timescales, alternative dip directions may be signi¢cant. The primary example is the western rift where S-dipping faults (some of which have initiated as recently as 0.5 Ma) are signi¢cant in recent stages of rifting. 3. Across-rift Migration The western part of the Corinth rift (west of 22.51) has experienced signi¢cant across-rift migration of fault activity (Ori, 1989; Goldsworthy & Jackson, 2001; Collier & Jones, 2003; Rohais et al., 2007).These studies discussed evidence for hangingwall migration only, as proposed elsewhere in the Aegean (Goldsworthy & Jackson, 2001). However, Bell et al. (2008) also presented evidence for footwall migration (i.e. from the West Channel to North and South Eratini faults). Incorporating these observations, we conclude that there is a general trend of northward fault migration during rift history; but that this includes both hangingwall and footwall migration in a system of N- and S-dipping faults. In the western rift, the northward migration of south coast faults has occurred over a wider acrossrift area than the northward fault migration o¡shore, resulting in overall rift narrowing (Bell et al., 2008). In the easternmost rift, the Megara basin faults, Kenchreai fault and other N-dipping faults bordering the NW Saronic Gulf may be considered ancestors to the currently active Skinos and Pisia faults of the Perachora peninsula (Collier et al., 1992; Goldsworthy & Jackson, 2001; Leeder et al., 2008). In the case of the Kenchreai fault, migration has apparently occurred in a single step 15 km to the north, resulting in uplift of the Corinth Isthmus area, rather than the more gradual 5 km stepping in the west (Fig. 1). 4. Along-axis rift development Although the chronology of rift evolution proposed here has signi¢cant error bars, to a general approximation, the rift system appears to have developed synchronously along-axis. For example, the oldest sediments in the o¡shore Gulf of Corinth and the Alkyonides Gulf are both


27

R. E. Bell et al. ca.1^2 Ma.There is therefore no evidence of lateral propagation of faulting along the rift axis on a million year timescale. This interpretation contrasts with models of both east to west (Armijo et al., 1996; Clarke et al., 1998; Nyst & Thatcher, 2004) and west to east (Ori, 1989) propagation for this rift.This result has strong implications for models of regional deformation in the Aegean, speci¢cally what drives rifting in the Corinth area. Importantly, our ¢ndings indicate that di¡erences in basement depth along the rift (Fig. 8a) are largely due to along-axis variations in displacement rate and not due to major di¡erences in the timing of fault initiation.

CONCLUSIONS This study has synthesised existing data for the entire Gulf of Corinth Basin (the currently active o¡shore rift basin) to investigate fault activity history and the spatio temporal evolution of the basin. The rift has undergone major changes in relative fault activity and basin geometry during its relatively short 1^2 Ma history. The major conclusions of this study are: (1) A major unconformity with an estimated age of ca. 0.4 Ma can be traced throughout much of the o¡shore basin, and separates two distinct seismic stratigraphic units. (2) Basement depth, which is expected to correlate broadly with net subsidence, is greatest in the centre of the rift ( 22120 0 to 221400 E) at 3 km, decreasing to the west and east. (3) We estimate the currently active Gulf of Corinth Basin to have formed ca. 1^2 Ma. This age, within error, appears to be consistent along the rift and we see no clear evidence to support growth or propagation of the currently active rift along axis on a million year timescale. Fault activity has migrated northward across the larger Corinth rift system during its history. (4) Throughout rift history, strain is distributed across several active faults in the westernmost basin (Zone 1, Fig. 9) producing a more complex basin structure, in contrast to other parts of the basin where one fault has typically dominated. (5) Both N- and S-dipping faults were signi¢cant along the rift before ca. 0.4 Ma, whereas N-dipping faults have dominated basin subsidence patterns since ca. 0.4 Ma (with the exception of the westernmost basin). (6) Before 0.4 Ma, two separated 20^50-km-long half graben depocentres were created, controlled by a Sdipping (western depocentre) and N-dipping (eastern depocentre) fault system (Fig. 9a). Post 0.4 Ma, a single depocentre 80-km-long formed controlled by several connected N-dipping faults on the southern margin, with maximum subsidence focused between the two older depocentres (Fig. 9b). (7) Sediment isopachs indicate major changes in slip rate on certain faults during rift history which dramatically

28

a¡ect depocentre development. Slip rates of Ndipping south coast faults in the Gulf of Corinth are 3^6 mm year 1 averaged over their 1 Myr history. Calculated slip rates on the N- and S-dipping o¡shore faults in the Gulf are apparently lower ( 1.4 mm year 1) but are likely underestimated. Individual fault slip rates are signi¢cantly lower in the Alkyonides Gulf ( 1^3 mm year 1). (8) Depocentre and fault evolution within the Gulf of Corinth indicate that the isolated fault stage of rift development, often present in numerical fault growth models, can persist for timescales of ca. 1 Myr and form major basins. This is comparable with observations from parts of the East African and North Sea rift systems. In more recent rift history (0^0.4 Ma), Gulf of Corinth basin subsidence is dominated by the connected N-dipping south coast fault systems.

ACKNOWLEDGMENTS We thank the crew of the MV Vassilios for their expertise during the western Gulf of Corinth survey and Carol Cotterill, Aris Stefatos, George Ferentinos, George Papatheoderou, Alfred Hirn, Chris Malzone, John Davis and Aggeliki Georgiopolou for technical and scienti¢c contributions. We also thank the Greek authorities for permission to conduct this work. Research was funded by Natural Environment Research Council grants NER/B/S/ 2001/00269 and NER/S/A/2004/12634, the University of Southampton, and the Royal Society. We thank P. Cowie, two anonymous reviewers and Editor P. van der Beek for constructive comments on an earlier version.

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Manuscript received 25 June 2008; Manuscript accepted 23 January 2009


Fault architecture, basin structure and evolution of the

Fault architecture, basin structure and evolution of the

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